4
Active Faults Related to Folding

ROBERT S.YEATS

Oregon State University

ABSTRACT

Active convergent zones contain fold-and-thrust belts that deform a sedimentary wedge by low-angle thrusting and flexural-slip folding over a subjacent, more rigid basement. Flexural-slip faults form by bedding slip during flexural-slip folding and are observed when unconformably overlying deposits are deformed by renewed folding. These faults are upthrown toward synclinal axes and die out in fold hinges. Bending-moment faults are produced because the convex side of a folded layer is lengthened normal to the fold axis and placed in tension, forming normal faults and extension fractures, whereas the concave side is shortened and placed in compression, forming reverse faults. Neither class of fault is likely to extend downward to rocks of such high strength that enough elastic strain energy could be stored to produce a large earthquake when released suddenly. However, all known historical examples are coseismic, and age relations on such second-order faults may apply to subjacent first-order seismogenic faults. Some folds such as Anticline Ridge at Coalinga, California, apparently are surface expressions of buried seismogenic faults. Regionally, fold-and-thrust belts may be modeled by the snow plow model of Davis et al. (1983) if the slope of the upper and lower boundaries of the forward-tapering wedge, the coefficients of basal and internal friction, and the ratio of pore pressure to hydrostatic pressure are known. This model implies that the deformation migrates toward the margin of the thrust belt such that youngest structures are at the edge; and pore pressures within individual folds, even those back from the edge, probably exceed hydrostatic.

INTRODUCTION

Folds and low-angle thrust faults are an important component of ancient mountain belts. Thus it is surprising that they are not more widely reported in the literature on active tectonics. Part of the reason for this is that mountain belts characterize zones of plate convergence, and most of these zones comprise the accretionary wedges confronting island arcs and are, accordingly, offshore. The southern margin of the convergence zone of southern Asia extending west from the Himalaya to the Zagros Mountains is perhaps the most notable exception on land.

Most of the literature on active tectonics deals with zones of strike slip: the San Andreas Fault system of California, the Alpine Fault system of New Zealand, and the North Anatolian Fault system of Turkey, among others. Yet even in these zones, there are smaller regions dominated by tectonic convergence: the Transverse Ranges of California and the ranges and basins of the



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Active Tectonics: Studies in Geophysics 4 Active Faults Related to Folding ROBERT S.YEATS Oregon State University ABSTRACT Active convergent zones contain fold-and-thrust belts that deform a sedimentary wedge by low-angle thrusting and flexural-slip folding over a subjacent, more rigid basement. Flexural-slip faults form by bedding slip during flexural-slip folding and are observed when unconformably overlying deposits are deformed by renewed folding. These faults are upthrown toward synclinal axes and die out in fold hinges. Bending-moment faults are produced because the convex side of a folded layer is lengthened normal to the fold axis and placed in tension, forming normal faults and extension fractures, whereas the concave side is shortened and placed in compression, forming reverse faults. Neither class of fault is likely to extend downward to rocks of such high strength that enough elastic strain energy could be stored to produce a large earthquake when released suddenly. However, all known historical examples are coseismic, and age relations on such second-order faults may apply to subjacent first-order seismogenic faults. Some folds such as Anticline Ridge at Coalinga, California, apparently are surface expressions of buried seismogenic faults. Regionally, fold-and-thrust belts may be modeled by the snow plow model of Davis et al. (1983) if the slope of the upper and lower boundaries of the forward-tapering wedge, the coefficients of basal and internal friction, and the ratio of pore pressure to hydrostatic pressure are known. This model implies that the deformation migrates toward the margin of the thrust belt such that youngest structures are at the edge; and pore pressures within individual folds, even those back from the edge, probably exceed hydrostatic. INTRODUCTION Folds and low-angle thrust faults are an important component of ancient mountain belts. Thus it is surprising that they are not more widely reported in the literature on active tectonics. Part of the reason for this is that mountain belts characterize zones of plate convergence, and most of these zones comprise the accretionary wedges confronting island arcs and are, accordingly, offshore. The southern margin of the convergence zone of southern Asia extending west from the Himalaya to the Zagros Mountains is perhaps the most notable exception on land. Most of the literature on active tectonics deals with zones of strike slip: the San Andreas Fault system of California, the Alpine Fault system of New Zealand, and the North Anatolian Fault system of Turkey, among others. Yet even in these zones, there are smaller regions dominated by tectonic convergence: the Transverse Ranges of California and the ranges and basins of the

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Active Tectonics: Studies in Geophysics northwestern South Island and central Otago, New Zealand, for example. Both the full-scale convergence zones, such as the Himalaya, and the zones subordinate to transform faulting, such as the Transverse Ranges, contain examples of active folds and low-angle thrust faults, and future research is likely to discover that active folds and thrust faults are as widespread in the active convergent zones as they are in extinct mountain belts. This chapter reviews the state of knowledge of structures related to folding in light of their impact on society. MECHANICAL BACKGROUND The mechanical properties of fold-and-thrust belts were considered by Elliott (1976) and Chapple (1978), building on the earlier work of Hubbert and Rubey (1959). Chapple (1978) noted that fold-and-thrust belts, whether on land or offshore, should show (1) a basal surface of décollement below which there is little or no deformation; (2) an overall shape in cross section of a wedge tapering toward the edge of the mountain belt with its base, the basal décollement, sloping toward the interior of the mountain belt; and (3) extensive horizontal contraction in the tapered wedge above the basal décollement. Davis et al. (1983) considered the mechanics of a fold-and-thrust wedge to be analogous to that of the wedge of snow that forms ahead of the blade of a moving snow plow. The snow deforms until the wedge attains a critical taper, then slides stably, growing as new snow is accreted at the front of the wedge. Parameters essential to an understanding of the mechanics of a fold-and-thrust wedge include the angle of topographic slope of the wedge toward the frontal edge of the deformed belt, the angle of rearward slope of the basal décollement, the coefficient of internal friction within the wedge, the coefficient of sliding friction on the base (about 0.85 according to Byerlee, 1978), and the ratio of pore fluid pressure to the vertical stress imposed by overburden (λ). Davis et al. (1983) applied their mechanical model to the active fold-and-thrust belt of western Taiwan, where extensive subsurface information is available, and they determined the critical parameters to be angle of forward topographic slope 2.9±0.3°, rearward slope of the décollement 6°, and λ equal to 0.7. The fold-and-thrust wedge is above sea level, and the topography is at steady state: thickening of the wedge by contractile tectonics is balanced by erosion, which proceeds at a rate of 5 to 6 mm/yr (Li, 1976; Suppe, 1981). The model is sensitive to the nature of material comprising the décollement, where it is assumed that essentially pure frictional sliding occurs. However, evaporites characterize the fold-and-thrust wedges of the Zagros Mountains (Stöcklin, 1968) and the Salt Range of Pakistan (Seeber et al., 1981), and these may yield plastically rather than by pressure-dependent Coulomb friction. The snowplow model has two important implications for active tectonics: (1) the age of deformation should migrate outward toward the front of the wedge, and (2) rocks within the wedge are near the point of critical failure and are likely to exhibit pore pressures greater than hydrostatic. Yeats et al. (1981) pointed out that faults that do not extend downward into rocks of high strength will not be expected to produce large-amplitude ground acceleration that is due to seismic shaking because such rocks under near-surface confining pressures are not capable of storing enough elastic strain energy to generate a large earthquake when that strain energy is released instantaneously. Because fold-and-thrust belts terminate downward at a basal décollement over an undeformed rigid basement, the question of their seismotectonic signature is an important one. Where the décollement contains rocks of such low strength that deformation may occur plastically under low confining pressure, as in the Zagros Mountains and the Pakistan Salt Range, internal deformation is probably not accompanied by large earthquakes (Berberian, 1981; Seeber et al., 1981; Seeber, 1983). In the thinner portions of fold-and-thrust wedges, the rocks are probably not under sufficient confining pressure to possess much shear strength, even though they behave according to Coulomb friction laws. However, in the thicker portions of wedges, the earthquake potential is not so clear. Seeber et al. (1981) suggested that the greatest earthquakes of the Indian Himalaya occur in front of the range in the Ganges foredeep, even though this area has been characterized by low instrumental seismicity in recent years. The large isoseismal areas of these earthquakes and lack of surface rupture in great historical earthquakes of the Ganges flood plain suggest that a large part of the décollement surface moves as a blind thrust, producing a detachment earthquake. The 1964 Prince William Sound, Alaska, earthquake also may be of detachment type (Seeber et al., 1981). FLEXURAL-SLIP FAULTS General Statement Most folds characteristic of foreland fold-and-thrust belts form by flexural slip. A stack of stiff beds alternating with thin, less stiff layers is end-loaded, and these beds buckle by slip on inherently weaker bedding surfaces separating stiffer beds (Ramsay, 1967, pp. 392–393). Slickensides on these weak surfaces are perpendic-

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Active Tectonics: Studies in Geophysics ular to the fold axis. Thickness of individual competent beds remains constant from hinge to limbs of folds. Rock layers are flexed, and upper layers slip over underlying layers toward anticlinal hinges and away from synclinal hinges. Slip is zero at fold hinges. On the limbs of the fold, the amount of slip on bedding faults depends on the maximum dip of the limbs and the thickness of each flexed layer. Because such bedding faults, called flexural-slip faults by Yeats et al. (1981), do not produce stratigraphic separation, they are not likely to be recognized unless there is a sequence overlying the flexural-slip fold with angular unconformity, and continued folding causes displacement of these younger deposits. Such deposits are cut by faults that are parallel to bedding in the folded sequence and upthrown toward the subjacent synclinal axis. The younger deposits are likely to be tilted toward the synclinal axis so that their dip is in the same direction as that in the subjacent, more strongly folded beds. Because flexural-slip faults remain in bedding, they do not extend to depths greater than the amplitude of the flexural-slip fold. Furthermore, displacement vanishes in the axis of the syncline where the faulted bedding plane is deepest. Beds at these depths, which would in most cases not exceed a few kilometers, are likely to have shear strength too low under such low confining pressures for bedding faults to generate large earthquakes, particularly in view of the fact that such faults form along the structurally weakest layers. Grey-Inangahua Basin, New Zealand Flexural-slip faulting was first described (although not named) by Suggate (1957) and Young (1963) in the Grey-Inangahua Basin of the northwestern South Island, New Zealand (Figure 4.1). The Grey-Inangahua Basin is a northeast-trending structural depression filled with Cenozoic strata resting with angular unconformity on a terrane consisting of Mesozoic and older rocks, including granitic basement. The Cenozoic strata are folded asymmetrically on the west side of the basin against the Paparoa Range. The folded strata are overlain with angular unconformity by glacial outwash gravels that form prominent surfaces above present river level. South of Giles Creek, one of these surfaces is cut by faults parallel to bedding in subjacent strata and upthrown toward the subjacent synclinal axis (Suggate, 1957). The surface is also tilted toward the synclinal axis (shown diagrammatically in Figure 4.2). A terrace riser cuts across the faults, and the fault scarps are higher on the more elevated terrace surface (Suggate, 1957; Yeats, in press; Figure 4.2). The changes in scarp height may be explained using Lensen’s FIGURE 4.1 Tectonic map of the Buller region, South Island, New Zealand. Heavy lines are faults with bar and ball on downthrown side. Arrows on monoclines point to downwarped side. Short heavy lines show flexural-slip fault sets in Grey-Inangahua Basin: BB, Blackball; BR, Big River; GC, Giles Creek; RC, Rough Creek (active in 1968); RO, Rotokohu (active in 1968); IF, Inangahua (active in 1968). (1968) reasoning at Branch River, New Zealand. Faulting occurred after the outwash stream had abandoned the terrace surface southwest of the terrace riser and while it still flowed across the lower terrace surface to the northeast. Further erosion removed the fault scarps northeast of the terrace riser where the stream still flowed but preserved the scarps on the abandoned, higher surface to the southwest. Renewed faulting occurred after the stream had abandoned the outwash surface altogether. The scarps northeast of the terrace riser reflect only the later faulting whereas the scarps southwest of the terrace riser reflect both episodes of faulting. Transverse Ranges, California Flexural-slip faulting is also documented in the Ventura Basin, California, northeast of Santa Paula, where Pliocene-Pleistocene strata are strongly folded and locally overturned on the north flank of the Santa Clara syncline and are overlain by alluvial fans derived from mountains to the north (Figure 4.3) (Keller et al., 1982; Rockwell, 1983). Surfaces of these fans are tilted toward the Santa Clara synclinal axis and are cut by faults that are parallel to bedding in the subjacent, strongly folded

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Active Tectonics: Studies in Geophysics FIGURE 4.2 Block diagram showing relations at Giles Creek, South Island, New Zealand. Outwash gravels overlie steeply dipping strata on limb of syncline in which synclinal axis is southeast of diagram. Gravels are tilted toward synclinal axis. Bedding-plane (flexural-slip) faults propagate upward through gravels and appear as scarps facing away from synclinal axis. Fault scarps are higher on high side of terrace riser, indicating that some faulting occurred after outwash drainage had abandoned the high side of the terrace riser but still occupied the low side. Additional faulting occurred after the drainage had abandoned the terrace altogether. Fault scarp A faces up the depositional slope and is ungullied, indicating that it formed so suddenly that the rills occupying old outwash channels were ponded at base of scarp. Fault scarp B also formed instantaneously, but it is small enough that rill drainage was maintained across it, thereby gullying the scarp. strata and are upthrown toward the synclinal axis. The faults show normal separation where the subjacent bedding is overturned and reverse separation where bedding is upright (Figure 4.3). Fan emplacement during tilting resulted in a cycle starting with fan deposition, then tilting followed by fan entrenchment, then deposition of a new fan surface, which was itself subsequently tilted. Seven geomorphic surfaces ranging in age from present day to an estimated 120,000 yr are recognized by Keller et al. (1982) and Rockwell (1983) on the basis of degree of soil development calibrated in part by 14C dates and on the amount of tectonic deformation. Four of the older surfaces show increasing tilt and scarp height with increasing age. Other examples of flexural-slip faulting in the Transverse Ranges are recognized in the Oakview area of the Ventura Basin (Keller et al., 1982) and Point Conception, west of Santa Barbara, California (Cluff et al., 1981). Shinano River, Niigata Prefecture, Japan Active folding has been recognized in Japan for many years (Otuka, 1942), and attempts have been made to compare rates of folding of river terraces with rates of folding of subjacent strata (Sugimura, 1967; Kaizuka, 1967; Nakamura and Ota, 1969). The most extensively documented active folds occur in central Japan along the Shinano River (Ota, 1969), which flows northward across folded strata of the Pliocene-Pleistocene Uonuma Group. These folds are overlain with angular unconformity by fluvial terraces that are themselves folded. Strike faults cut the terraces at Yamamoto-yama (cross section 6 of Figure 6 of Ota, 1969) and northwest of Ojiya town (cross section 7 of Figure 6 of Ota, 1969). At Yamamoto-yama, Ota observed that the faults are upthrown in the direction of the synclinal axis of the folded Uonuma Group and of the oldest fluvial terrace. In a recent visit to the area with the author, she confirmed that the terrace remnants between the strike faults slope toward the synclinal axis, indicating that the faults are probably flexural-slip faults. The early work northwest of Ojiya town had suggested that the strike faults there are downthrown toward the synclinal axis. However, Ota and Suzuki (1979) described a new quarry exposure near Katakai, which exposed the faulted terrace and the underlying Uonuma Group (Figure 4.4). Four faults parallel to bedding in the Uonuma Group cut the overlying terrace and are upthrown toward the synclinal axis. In the quarry exposure, the steeply dipping Uonuma Group FIGURE 4.3 Diagrammatic cross section across north flank of Santa Clara syncline east of Santa Paula, Ventura Basin, California. Flexural-slip faults have normal separation where subjacent strata are overturned, reverse separation where subjacent strata are upright. Older fan has more separation and tilt than younger fan. Modified from Yeats et al. (1981) and Keller et al. (1982). No vertical exaggeration. Inset: OT, Orcutt-Timber Canyon flexural-slip faults; OV, Oakview flexural-slip faults.

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Active Tectonics: Studies in Geophysics FIGURE 4.4 Quarry exposure of tilted and faulted Shinano River stream terrace unconformably overlying steeply dipping Uonuma Group near Katakai, Niigata Prefecture, Japan. Two faults in gravel (F1 and F2) are propagated upward into terrace material from contacts between conglomerate and fine-grained strata of Uonuma group (from Ota and Suzuki, 1979). consists of fine-grained nonmarine strata with two interbeds of conglomerate. Two of the four faults occur at the contact between a gravel interbed and the finer grained strata, suggesting that the faults are localized by lithology contrasts in the underlying, folded strata. Examples of Coseismic Flexural-Slip Faulting There are no known examples of flexural-slip faults formed by aseismic creep. However, there are examples of such faults accompanying earthquakes at Lompoc, California; El Asnam, Algeria; and possibly Inangahua, New Zealand. On April 7, 1981, an earthquake of ML= 2.5 in a diatomite quarry near Lompoc, California, produced a zone of reverse fault scarps at least 575 m long, with the fault plane formed in clay interbeds in diatomite and diatomaceous shale on the north flank of a syncline (Yerkes et al., 1983). Maximum net slip was 25 cm, with maximum dip slip of 23 cm and right-lateral strike slip of 9 cm. The earthquake is best explained FIGURE 4.5 Diagrammatic cross section of coseismic flexural-slip faulting at Kef el Mes, El Asnam earthquake, 1980, after Philip and Meghraoui (1983). Earthquake-generating reverse fault at A; earthquake-generated faults at B produced by renewed folding of syncline. by removal of diatomite during quarrying under conditions in which maximum principal compressive stress is horizontal and normal to the synclinal axis. Removal of overburden increased the shear stress on the bedding plane enough to cause the earthquake 2 yr after the end of quarrying (Yerkes et al., 1983). The El Asnam, Algeria, earthquake reactivated a reverse fault and was accompanied by extensive internal deformation of the rocks adjacent to the surface trace of this fault (Philip and Meghraoui, 1983). At Kef el Mes, near the northeastern limit of 1980 surface rupture, the reverse fault was accompanied by renewed flexural-slip folding of Pliocene strata in an adjacent syncline (Figure 4.5). Flexural-slip faulting on bedding planes produced fault scarps that produced displacement in an opposite sense to that on the main fault. The Inangahua, New Zealand, earthquake of May 24, 1968, was apparently generated on the Inangahua reverse fault, which joins the Glasgow Fault to the northeast (Figure 4.1; Lensen and Suggate, 1968; Lensen and Otway, 1971; Nathan, 1978a). The Rotokohu and Rough Creek Faults were formed during this earthquake by slip parallel to bedding on the north flank of a syncline in folded Neogene strata, producing scarps in late Pleistocene outwash gravels that overlie the folded strata unconformably (Lensen and Otway, 1971). Vertical leveling profiles across the Rotokohu Fault also showed this displacement (Boyes, 1971). Lensen (1976) later used the surface faulting accompanying the Inangahua earthquake to classify faults as earthquake-generating (Inangahua) and earthquake-generated (Rotokohu and Rough Creek). However, the Rotokohu and Rough Creek Faults are parallel to a sharp linear gradient in isostatic gravity anomalies (Anderson, 1979), and they may be the surface expression of a structure involv-

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Active Tectonics: Studies in Geophysics ing basement rocks rather than faults produced by flexural-slip folding. Field evidence suggests that at least the larger flexural-slip faults at Giles Creek and Blackball in the Grey-Inangahua Basin are coseismic (Yeats, in press). The flexural-slip faults at both these localities face west, up the depositional slope of the outwash surfaces that they cut (Figure 4.2). The surface at Giles Creek is grooved by braided channels that formerly carried outwash gravels from glaciers on the east side of the Paparoa Range. After abandonment of the outwash surface, the braided channels remained, but they now carry only the runoff derived from the surface itself whereas formerly they carried detritus from the Paparoa Range. Most of the west-facing fault scarps are ungullied, even though they are very old; the faults cut surfaces more than 100,000 yr old (Suggate, 1965; Nathan, 1978b; R.P. Suggate, New Zealand Geological Survey, personal communication, 1984), but they do not cut an outwash surface at Blackball considered to be 18,000 to 23,000 yr old (Nathan, 1978b; Suggate, 1965; Suggate and Moar, 1970). The ungullied west-facing scarps at Giles Creek are in contrast to extensively gullied east-facing scarps on the same surface, scarps presumably related to range-front faulting (Yeats, in press). Rill drainage down the old outwash channels dissects the scarps that face east, down the depositional slope, but collects at the base of the west-facing scarps, forming small bogs or small streams that drain parallel to the scarp and off the surface. The only west-facing scarps that are gullied are those in which scarp heights are relatively low, such as the southern parts of scarps B and D of Suggate (1957) and scarp C of Young (1963). If the west-facing scarps had formed by aseismic creep on flexural-slip faults, the rill drainage should have been able to maintain itself across the slowly rising fault scarps (Figure 4.2), just as it does when scarp height is very low. However, the drainage is blocked by the scarps, suggesting that the scarps formed suddenly, accompanying an earthquake. To be ungullied by the drainage, a scarp must appear instantaneously and must be high enough that the former drainage, even at maximum stream flow, is unable to overtop the new scarp. The flexural-slip fault scarps northeast of Santa Paula, California, face up the depositional slope of the fans that they cut (Figure 4.3), but the high initial slope of the fan surface and the high-volume stream flow under the occasional torrential rainfall experienced in southern California result in gullying of the newly formed scarps, even if they do form suddenly. BENDING-MOMENT FAULTS Deformation of a flexed layer can be treated as bending an elastic plate around a fold axis. If the plate is bent by equal and opposite moments applied at its ends, the convex side is lengthened and placed in tension, and the concave side is shortened and placed in compression (Figure 4.6; cf. photoelastic experiments of Currie et al., 1962). The compression and tension of the sides are produced by a couple or bending moment (Johnson, 1970, pp. 41–50). Between that portion of the plate in compression and that portion in tension, there is a neutral surface on which there is neither compression nor tension (Figure 4.6). The strain within a bent plate is approximately proportional to the distance from the neutral surface and inversely proportional to the radius of curvature of the plate. On the convex surface, the minimum principal compressive stress (σ3) is tangent to the plate surface but perpendicular to the axis of bending, whereas on the concave surface, the maximum principal compressive stress (σ1) has this orientation. If the neutral surface is located at the center of the plate, the deviatoric stress (σ1−σ3) is zero at the neutral surface and maximum at the convex and concave surfaces. If the beam is structurally isotropic and homogeneous, and folding is upright, initial rupture above the yield stress will be by extension fractures or normal faults at the convex surface and by reverse faults at the concave surface (Figure 4.6). These faults are small scale because they extend to no greater depth than the neutral surface of the flexed plate. Such faults should not be capable of large earthquakes if the flexed plate is a sedimentary bed of low strength and the radius of curvature of the flexure is relatively small. The importance of bending-moment faults in seismic-risk evaluations is in the fact that these faults are likely to FIGURE 4.6 Bending moment faults developed on the concave and convex sides of a flexed layer. Faults should occur on both upper and lower surfaces, but they have only been observed on upper surface.

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Active Tectonics: Studies in Geophysics be encountered in trenches across scarps suspected of being faults. The trench is dug with the strategy of encountering late Quaternary deposits cut by a suspected fault. To be recognized as bending-moment faults, fault displacement must be normal on the convex side of a fold and reverse on the concave side. Two examples from trench investigations follow. The Ventura Fault in the Ventura Basin of California forms a linear, south-facing scarp that trends eastward along the foot of the hills bordering the city of Ventura on the north (Sarna-Wojcicki et al., 1976). In east Ventura, the scarp diverges from the foothills and maintains an easterly strike whereas the edge of the hills trends more to the east-northeast (Figure 4.7). In this area, the surface of the Harmon alluvial fan is cut by a south-facing scarp attributed to the Ventura Fault. The scarp forms the boundary between an uplifted bench of older deformed deposits on the north and young undeformed sediments on the south (Sarna-Wojcicki et al., 1976). The seismotectonic setting of this fault in a region of north-south horizontal contraction predicts that it should have reverse displacement. However, two trenches across the Ventura scarp (Figure 4.7) show several normal faults that offset all but latest Holocene sediments (Sarna-Wojcicki et al., 1976; Gardner and Stahl, 1977). The normal faults are concentrated in the flexed region and are interpreted as bending-moment faults. These faults are secondary features apparently produced by displacement on the subjacent Ventura Fault. Age of youngest sediments displaced by the normal faults can be assumed to date the latest movement on the Ventura Fault, but the orientation and sense of displacement on these normal faults cannot be extrapolated to infer sense of displacement on the Ventura Fault. In the city of Camarillo, south of Ventura, trenches dug by Geotechnical Consultants, Inc., across a linear, south-facing scarp (Gardner, 1982) were dug to shed light on late Quaternary movement history of the Camarillo Fault, a north-side-up member of the Simi reverse-fault system (Figure 4.8). The trenches (Figure 4.8) revealed an increase in south dip northward toward the scarp, but instead of a high-angle reverse fault dipping north, a low-angle, north-dipping reverse fault and a set of high-angle, south-dipping reverse faults were found. These faults are interpreted as being produced by bending moment on the concave side of a monoclinal bend in Quaternary sediments that are either draped over a buried fault or are part of a pressure ridge. This interpretation makes it unlikely that the faults exposed in the trenches are seismogenic, and it also implies that the trenches provide no evidence about whether the Camarillo Fault itself is seismogenic. As in the case of the Ventura Fault, the youngest sediments warped across the scarp or cut by bending-moment faults can be assumed to have been deposited prior to latest movement on the subjacent Camarillo Fault. Bending-moment faults are also found at Toppenish Ridge, an east-trending anticline in Miocene basalts and Quaternary sediments in south-central Washington State (Campbell and Bentley, 1981; Figure 4.9). On the north side of Toppenish Ridge, surface ruptures near the hinge of the overturned Satus Peak anticline in basalt cut late Quaternary alluvium and landslide material. These ruptures were interpreted by Campbell and Bentley (1981) as related to bending moment in the hinge of the fold. The hinge of the overturned syncline immediately to the north is characterized by thrust faults that cut Quaternary alluvium. Campbell and Bentley (1981) interpreted these as the surface expression of a décollement thrust (Mill Creek thrust) passing underneath the anticline. I consider it to be more likely that they are due to bending moment on the concave side of the syncline. Finally, Philip and Meghraoui (1983) present evidence that normal faulting on the hanging wall of the 1980 El Asnam thrust fault (Figure 4.10) is produced by bending moment as the strata in the hanging wall are folded. These are called “extrados fractures” by Philip and Meghraoui (1983). In addition to normal faults, tension cracks appear close to the surface with edges vertically offset and with dip-slip slickensides. Two cases are presented. Where thrusting is dip-slip (Figure 4.10A and 4.10B), normal faults are parallel to the thrust fault trace and to the anticlinal axis. Where thrusting is oblique-slip (Figure 4.10C and 4.10D), the normal-fault grabens develop at an angle to the thrust fault trace and the anticlinal axis. In plan view, they resemble pull-apart basins. The extension is due not only to anticlinal bending but also to left-oblique slip in the hanging wall. As in the case of the bedding-slip faults at Kef el Mes, these faults are low-seismic although they are coseismic—that is, they are part of the deformation accompanying a major thrust-fault earthquake, but they themselves do not deform strata that are strong enough to store much elastic strain energy. FOLDS RELATED TO FAULTING The discussion above has concentrated on faults that are by-products of folding. They are near-surface structures, hence they are small scale and unlikely to generate large earthquakes. However, the flexural-slip and bending-moment faults at El Asnam accompanied the 1980 earthquake (Philip and Meghraoui, 1983). They accompanied near-surface folding that was itself a by-product

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Active Tectonics: Studies in Geophysics FIGURE 4.7 Sketches of two trenches across Ventura Fault simplified from Sarna-Wojcicki et al. (1976) and Geotechnical Consultants, Inc. Ventura Fault is expressed as a flexure with slopes and dips steeper than areas to the north and south. Hospital trench has vertical exaggeration ×2; reservoir trench has no vertical exaggeration. In both trenches, fractures are concentrated where strata are flexed. Most fractures are normal faults or soil-filled extension fractures, as would be expected if fractures are related to bending moment.

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Active Tectonics: Studies in Geophysics FIGURE 4.8 A and B, Simplified sketches of trenches across Camarillo Fault, after D.A.Gardner (1982; unpublished technical reports by Geotechnical Consultants, Inc.). No vertical exaggeration. C, Topography of ridge in downtown Camarillo from U.S. Geological Survey Camarillo 7 1/2-minute quadrangle, locating two trenches. Camarillo Fault is presumed to control steep south flank of this ridge. D, Sketch of presumed relation between bending-moment faults in trenches and subjacent Camarillo Fault.

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Active Tectonics: Studies in Geophysics FIGURE 4.9 Tectonic sketch and diagrammatic cross section of Satus Peak anticline at Toppenish Ridge, Washington, redrawn and modified from Campbell and Bentley (1981). Sawteeth mark the hanging-wall side of reverse faults; ticks mark the hanging-wall side of assumed normal faults. of a deep-seated left-lateral reverse fault that generated the 1980 mainshock (Philip and Meghraoui, 1983) (Fault A of Figure 4.5). In a similar fashion, the folds generating the flexural-slip faults in the Grey-Inangahua Basin, New Zealand, may themselves be related to deep-seated reverse faulting at the margin of the Paparoa Range (Figure 4.1). The seismogenic Newport-Inglewood Fault in the Los Angeles Basin, California, is for the most part not exposed at the surface. The fault cuts basement rocks, but the fault zone is expressed at the surface as en echelon anticlines formed as pressure ridges along the fault (Barrows, 1974; Yeats et al., 1981). Near Ventura, California, the Montalvo Mounds are late Quaternary anticlinal ridges that mark the surface expression of the Oak Ridge high-angle reverse fault that does not itself reach the surface (Yeats et al., 1981). The May 1983 earthquake at Coalinga, California (MS=6.5), did not rupture the ground surface (although one aftershock did) but instead augmented a fold at Anticline Ridge, as based on geodetic data (Stein and King, 1984). Deformation of the stream bed of Los Gatos Creek, an antecedent stream that cuts across the FIGURE 4.10 Coseismic bending-moment faults associated with 1980 El Asnam earthquake, after Philip and Meghraoui (1983). Displacement on the main seismogenic thrust fault was accompanied by anticlinal folding in the hanging wall, which was itself accompanied by grabens along the anticlinal crest. These grabens are parallel to the anticlinal crest (A, B) or oblique to the anticlinal crest and stepped right where thrusting had a left-slip component (C, D). Map symbols: 1, normal fault; 2, tensile crack; 3, en echelon cracks; 4, thrust faults and pressure ridges; 5, attitude of dipping bed; 6, attitude of horizontal bed.

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Active Tectonics: Studies in Geophysics anticline, and an alluvial fan surface indicates that folding similar to the 1983 event has been taking place for the past 2500 to 10,000 yr at a rate of surface uplift of 1 to 4 mm/yr. The folding resembles that of subjacent strata and indicates that the folding may have been taking place for at least the past 2 m.y. Stein and King (1984) modeled the folding as having been generated by a reverse fault with a slip rate of 3 to 12 mm/yr based on the deformation of Los Gatos Creek and 1 to 4 mm/yr based on the deformation of underlying strata. Folding is coseismic with an average recurrence interval of earthquakes the size of the 1983 event of 200 to 1500 yr (Stein and King, 1984). In dissent, Hill (1984) noted the failure of the 1983 aftershocks to define a subsurface fault plane. Focal mechanisms of 10 earthquakes at Coalinga have nodal planes that are parallel to folded bedding at the epicenter of each event, leading Hill (1984) to suggest that the Coalinga earthquake was produced by deep-seated flexural-slip folding. ACTIVE TECTONICS OF ON-LAND FOLD-AND-THRUST BELTS Flexural-slip faults and bending-moment faults are secondary features related to flexural-slip folding, which itself characterizes foreland fold-and-thrust belts. Davis et al. (1983) developed their mechanical model from a study of western Taiwan (cf. Suppe and Wittke, 1977; Suppe and Namson, 1979; Suppe, 1980a,b, 1981; Namson, 1982). The active-tectonic setting of fold-and-thrust belts is illustrated in this paper by a discussion of the active foreland thrust belt of northern Pakistan and India, part of the Himalayan convergence zone between the Indian and Eurasian plates, and the central Ventura Basin of California, where the convergence zone is limited in length, apparently controlled by the big bend of the San Andreas Fault. These two regions offer the possibility of determining rates of convergence directly from the geologic record. Active foreland thrusting occurs on a continental scale in the foothills of the Himalaya as the Indian shield is overridden by its own northern margin in a series of south-verging thrusts (Yeats and Lawrence, 1984). The Precambrian Indian shield slopes gently northward beneath the Indo-Ganges floodplain and is overlain by flat-lying molasse deposits (Siwalik Group), which are older versions of the modern drainage system (Figure 4.11). In India, the alluvial plain consists of a northern domain in which large rivers break out of the mountains and flow south across broad distributary fans, a central domain in which the south-flowing streams merge with the east-southeastward flow of the main Ganges River and flow parallel to the Himalayan front, and a southern domain in which small-volume rivers flow northward from the Indian shield into the Ganges (Geddes, 1961; Figure 4.11A). The depositional axes of Pleistocene and older Siwalik molasse basins are parallel to the Ganges River but north of it; the older the molasse, the farther north its depositional axis (Acharyya and Ray, 1982). The molasse basins are forced southward owing to the southward advance of Himalayan thrust sheets such that it may be possible at a given site to progress upsection from southerly, shield-derived sediments to northerly, Himalaya-derived sediments (Figure 4.11B; Tandon, 1976; Parkesh et al., 1980) and finally to the appearance of the thrust sheets themselves. Lyon-Caen and Molnar (1985) suggested that the age of the basal molasse sediments overlying the Indian Shield decreases southward at a rate of 10–15 mm/yr. Active faults are found in the interior of the Himalaya of Nepal and India as well as the southern mountain front, but only the mountain-front fault (Himalayan Frontal thrust) shows north-over-south thrusting (Nakata, 1972, 1982; Nakata et al., 1984). Major thrusting in the Kathmandu intermontane basin predates deposition of the Lukundol Formation, which is Pliocene in age as based on magnetostratigraphy and vertebrate fossils (Yoshida and Igarashi, 1984; M.Yoshida, personal communication, 1984). The southward advance of décollement thrusting may be better calibrated in northern Pakistan (Figure 4.12) where the molasse basin is wider and subsurface evidence more abundant. New thrusts appear within the previously undeformed foreland sediments resting on top of the northward-sloping Indian shield. The Salt Range overrides its own fan material and alluvium along the newest thrust—the Salt Range thrust of late Quaternary age—and most of the major deformation is younger than 0.4 m.y. ago (Ma) (Yeats et al., 1984; Point A, Figure 4.12). Farther north, near Rawalpindi, major deformation is dated as between 1.9 and 2.1 Ma on the basis of molasse deposits calibrated by magnetostratigraphy and tephrochronology above and below the major angular unconformity (Raynolds and Johnson, in press) (point B, Figure 4.12), and the Soan syncline began to control deposition about 3 Ma (Raynolds, 1980). The distance between points A and B projected perpendicular to strike is 100 km, and the age difference between the end of deposition of subsequently deformed molasse deposits at the two points (2.1–0.4 Ma) is 1.7 m.y., a rate of southward migration of deformation of 60 mm/yr. Farther north, at the southern edge of the Peshawar Basin (point C, Figure 4.12), major thrusting predates deposition of nonmarine deposits as old as 2.8 Ma as dated by magnetostratigraphy and tephrochronology (Burbank, 1983; Burbank and Tahirkheli, 1985; Yeats and Hussain, 1985). The distance between points B and C projected perpendicular to strike is 70 km, and

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Active Tectonics: Studies in Geophysics FIGURE 4.11 A, Typical profile of the Himalayan thrust front and molasse sediments controlled by it. The Himalaya undergoes uplift as a consequence of foreland thrusting, shedding sediments into the flanking molasse basin. These side tributaries are collected by a master stream (Indus, Jhelum, Ganges) flowing parallel to mountain front. Tributaries on the south are derived from Indian shield. B, Southward migration of deformation front and sedimentary facies boundaries of molasse due to advancing foreland thrust sheets, compared to plate convergence rate. Depocenters and facies migrate south at 30 mm/yr (Raynolds, 1980), basal molasse sediments become younger southward at 10–15 mm/yr (Lyon-Caen and Molnar, 1985), and deformation as dated by angular unconformities migrates south at a poorly controlled rate of 60–78 mm/yr. FIGURE 4.12 Tectonic map of southern margin of Himalaya in northern Pakistan. Intermontane basins are shaded. Points A, B, and C are localities where age of major deformation may be dated; see text. Sawteeth mark late Quaternary thrusts; arrows mark late Quaternary strike-slip fault. the age difference between the oldest deposits after major deformation at the two points (2.8–1.9 Ma) is 0.9 m.y., a southward younging of 78 mm/yr. Those rates imply that new thrusts appear and propagate in the previously undeformed molasse basin at rates higher than the convergence between the Indian and Eurasian plates. Southward migration of the Siwalik foredeep in the Jhelum re-entrant, eastern Potwar Plateau, as based on southward progradation of conglomerate facies (Raynolds, 1980), takes place at a rate of 30 mm/yr. This progradation was influenced by increased uplift rates in the source areas as well as southward advance of Himalayan thrust sheets, so the rate of advance of thrust sheets would be less than 30 mm/yr. These rates may be compared with the northward motion of the Indian plate with respect to Eurasia of 40 mm/yr at the western end of the Himalaya to 65 mm/yr at the eastern end (Minster and Jordan, 1978). Either most of the conver-

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Active Tectonics: Studies in Geophysics gence between Eurasia and India is being accommodated at the southern margin of the thrust belt rather than being more uniformly distributed between the thrusts at the southern margin and more interior and more poorly dated thrust zones, such as the Main Central Thrust, or southward propagation of thrusts takes place faster than the rate of plate convergence. The rate of underthrusting of India beneath the Himalaya is 18 mm/yr based on data from large earthquakes since the year 1900 (Molnar and Deng, 1984), suggesting that only part of the Indian-Eurasian convergence is being taken up within the Himalaya, with the remainder being taken up by escape-block tectonics farther north. The Transverse Ranges of California are controlled by north-south contractile tectonics related to the big bend in the San Andreas Fault. In the Ventura Basin, continuous sedimentation throughout most of Quaternary time was strongly influenced by contractile tectonics. These sediments have been age-calibrated by tephrochronology, magnetostratigraphy, and radiometric dating (Izett et al., 1974; Blackie and Yeats, 1976; Lajoie et al., 1979, 1982; Liddicoat and Opdyke, 1981). Based on this calibration and on the construction of balanced cross sections across the Ventura Basin, the rate of convergence of the northern edge of the Ventura Basin against its southern edge is 23 mm/yr over the past 0.2 m.y. at Ventura (Yeats, 1983), over half the northsouth component of Pacific-American plate motion in this area, which is 42 mm/yr (Bird and Rosenstock, 1984). Most of this shortening occurred across the Ventura Avenue anticline and the adjacent syncline to the north (Yeats, 1982). The Ventura Avenue anticline contains a giant oil field with more than 1460 wells, resulting in an extensive, detailed data set on its internal structure (Figure 4.13). Prior to formation of the anticline, the Taylor low-angle thrust fault set began to form 1.3 Ma along a weak layer in the Pliocene turbidite sequence, moved up a 45° ramp, and stopped motion about 0.65 Ma. Maximum net slip rate was 2.8 mm/yr to the southeast (Yeats, 1983). Following the end of deposition of nonmarine coarse-grained strata of the San Pedro Formation about 0.2 Ma, the anticline began to buckle. The Ventura River was antecedent to this buckle, and uplift of the crest of the fold is calibrated by deformed river terraces that have been dated by 14C, with age extrapolations beyond the limits of 14C dating based on a soils chronosequence (Keller et al., 1982; Rockwell, 1983). Uplift rates on the anticlinal crest were 4.3 to 5.2 mm/yr for the last 29,600 yr, 10.5 to 11.5 mm/yr from 80,000 to 29,600 yr ago, and 15 to 16 mm/yr from 200,000 to 80,000 yr ago (Keller et al., 1982). Rockwell (1983) demonstrated that a rootless buckle fold formed by end loading would undergo rapid displacement normal to its loading direction early in its formation, and this displacement would slow down as the fold became more fully developed [Rockwell, 1983; Yeats, 1983; cf. theoretical considerations by Currie et al. (1962) and Adams (1984)]. The fold is still under high horizontal stress because it is overpressured, with the ratio of fluid pressure to overburden pressure increasing from 0.55 to 0.8 as radius of curvature decreases from 300 to 30 m in the core of the fold (Yeats, 1983). If the fold is still undergoing contraction as uplift of its crest continues at a decreasing rate, how is this contraction being accommodated? The limbs may still be steepening, but an alternative way to accommodate contraction may be the Ventura Fault (Figure 4.7). The south-facing fault scarp occurs at the sharp boundary between flat-lying strata of the Santa Clara syncline and south-dipping strata of the south flank of the Ventura Avenue anticline. This sharp boundary dips 45° north to depths of several kilometers and is planar, and it is assumed that this boundary is the Ventura Fault at depth (Figures 4.7 and 4.13). It is younger than most of the folding because it is planar, in contrast to other reverse faults that curve over the crest of the anticline and, therefore, formed as low-angle thrusts during the early stages of folding (Figure 4.13). Although the fault scarp is linear and well documented, subsurface well correlations indicate little or no stratigraphic separation across the fault. As shown in Figures 8, 10, and 11 of Yeats (1982), horizons can be correlated across the fault with a sharp change in dip, but no displacement. This led me to propose earlier that the Ventura Fault formed by bending moment (Yeats, 1982), a conclusion that is probably wrong. My present view is that the fault formed so recently that it has not had time to accumulate enough displacement to be documented without ambiguity in the subsurface. The evidence for this is the planar nature of the boundary between flat-lying and steeply dipping strata. The Ventura Avenue fold may have acquired a configuration that is stable to continued horizontal stress if there is a zone south of the fold along which further displacement may take place. If displacement occurs by southward sliding of the Pliocene-Pleistocene turbidite sequence over a subjacent Miocene ductile sequence, this displacement may break through to the surface as the Ventura Fault, as illustrated in the inset to Figure 4.13. SOCIETAL IMPLICATIONS OF ACTIVE FAULTS AND FOLDS Yeats et al. (1981) suggested that flexural-slip faults of the Ventura Basin would cause ground-rupture prob-

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Active Tectonics: Studies in Geophysics FIGURE 4.13 North-south cross section across Ventura Avenue anticline near Ventura River, showing the 6553 m deep test well, Shell-Taylor 653 (T-653). In inset: ORF, Oak Ridge Fault; RMF, Red Mountain Fault; VA, Ventura Avenue anticline; VF, Ventura Fault. Oak View terrace dated as 32,000 yr before present. lems, but would not result in large earthquakes with major seismic shaking. These faults curve into bedding in flexural-slip folds, and they do not extend downward into rocks of such high strength that sudden release of stored elastic strain energy would cause a large earthquake. The 1981 Lompoc earthquake showed that flexural-slip faults are not aseismic, and Yeats (1982) referred to them as low-shake faults. Bending-moment faults are also low-shake faults, as are at least some thrust faults involving only the sedimentary cover. But all known historical examples of bending-moment and flexural-slip faults accompany earthquakes, so if they are low-seismic, they are also coseismic. There are no known examples of flexural-slip faults forming by creep. The Giles Creek, New Zealand, flexural-slip faults show evidence of at least two episodes of faulting within a few thousand years, then no evidence of additional faulting in the last 20,000 yr, despite the fact that the Grey-Inangahua Basin where they occur produced a large earthquake in 1968. Flexural-slip faults (and perhaps bending-moment faults as well) that cut late Quaternary deposits may be used to monitor the recurrence of stick-slip faulting on subjacent seismogenic faults that may not cut late Quaternary deposits. The relations between coseismic flexural-slip and bending-moment faults to the seismogenic El Asnam thrust are instructive to this point. The flexural-slip faults northeast of Santa Paula, California, may provide information about movement on the nearby San Cayetano Fault (Figure 4.3), and the flexural-slip faults in the Grey-Inangahua Basin may document displacement on seismogenic reverse faults bounding the basin on the west. Finally, the coseismic growth of the Anticline Ridge anticline during the 1983 Coalinga, California, earthquake may provide information on the recurrence interval of a subjacent seismogenic reverse fault. An implication of the mechanical model of Davis et al. (1983) is that folds in active fold-and-thrust belts are likely to be overpressured, and Davis et al. (1983) summarize evidence that overpressured folds exist in subaerial fold-and-thrust belts in Taiwan and the Himalaya and in submarine accretionary wedges adjacent to the Middle America, Aleutian, and eastern Caribbean trenches and off the coasts of Oregon and the Pakistani Makran. The Ventura Avenue anticline, in a somewhat similar contractile environment, is perhaps the most extensively documented overpressured structure in the world. Future oil exploration in active fold-and-thrust belts offshore and onshore should expect overpressured reservoirs and adopt necessary techniques to prevent blowouts. Another implication of the model of Davis et al. (1983) is that folds migrate out toward the edge of the

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Active Tectonics: Studies in Geophysics fold belt, and therefore the most frontal folds and thrusts are most likely to be still growing. Accordingly, ground rupture is most likely in these frontal structures and less likely in those structures farther back. However, even these more internal structures are likely to be overpressured, and changes of fluid pressure due to water flooding may produce failure, because the tapered wedge is assumed to be on the verge of shear failure throughout (Davis et al., 1983). Where the basal décollement in fold-and-thrust wedges is composed of material that deforms plastically, displacement may not be accompanied by large earthquakes. Where the basal décollement yields by Coulomb friction, its ability to produce large earthquakes may depend on the thickness of the wedge. Where the wedge is thick, very large earthquakes may occur (Seeber et al.,1981). ACKNOWLEDGMENTS My work has been supported principally by contracts from the Earthquake Hazards Reduction Program of the U.S. Geological Survey. Work in New Zealand was supported by the U.S. Geological Survey, by Grant INT-82–19897 from the National Science Foundation, and by Oregon State University. Work in Pakistan was supported by Grant INT-81–18403 from the National Science Foundation, and my visit to Japan was sponsored by Grant EAR-83–18194 from the National Science Foundation. R.P.Suggate reviewed an early draft to the paper. REFERENCES Acharyya, S.K., and K.K.Ray (1982). Hydrocarbon possibilities of concealed Mesozoic-Paleogene sediments below Himalayan nappes—reappraisal, Am. Assoc. Petrol. Geol. Bull. 66, 57–70. Adams, J. (1984). Active deformation of the Pacific Northwest continental margin, Tectonics 3, 449–472. Anderson, H.J. (1979). 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