Chapter 2 Strike-Slip Basin – Its Configuration and Sedimentary Facies

Plate convergent margins are areas of concentrated lithospheric stress. They include areas of compression accompanied by thrusting, mountain building, and related foreland/forearc basin development, and also areas of strike-slip movement associated with transpressional uplift, transtensional subsidence, or pure strike-slip displacement. The degree of shortening and uplift or extension and subsidence depends on the modes of convergence between oceanic plates, island arcs, and continental crusts. Strike-slip faulting is one of the most important mechanisms of sedimentary basin formation at plate convergent margins, where localized extension can cause topographic depressions. Sedimentary strata deposited in these basins record the history of lithospheric response to the convergence.


Introduction
Plate convergent margins are areas of concentrated lithospheric stress. They include areas of compression accompanied by thrusting, mountain building, and related foreland/forearc basin development, and also areas of strike-slip movement associated with transpressional uplift, transtensional subsidence, or pure strike-slip displacement. The degree of shortening and uplift or extension and subsidence depends on the modes of convergence between oceanic plates, island arcs, and continental crusts. Strike-slip faulting is one of the most important mechanisms of sedimentary basin formation at plate convergent margins, where localized extension can cause topographic depressions. Sedimentary strata deposited in these basins record the history of lithospheric response to the convergence.
Sedimentary successions of archetypal examples of strike-slip basins, such as the Ridge Basin in California, have been characterized in terms of the dominance of axial sediment supply and continuous depocenter migration in a direction opposite to that of the sediment supply [1]. Their basin lengths are typically about three times longer than the basin widths [2]. The Izumi Group of the Cretaceous turbidite successions in southwestern Japan [e.g., 3 and references therein] has the same characteristics in the sedimentary succession as the Ridge Basin, and is therefore considered to contain strata that were deposited in a strike-slip basin. However, the basin geometry is quite different from that of the Ridge Basin, whose shape is more elongated with a length of more than 300 km and a width of less than 20 km, and whose southern margin has been truncated by post-depositional strike-slip fault displacement. Strike-slip basins thus present a wide diversity in terms of their geometry, evolution, and filling processes.
Since many local examples were collected after the 1980s [4], advanced research techniques including subsurface exploration including seismic survey and borehole drilling [e.g., 5], analog experiment [e.g., 6], and numerical simulation [e.g., 7] revealed that basin formation and filling processes were not simple but variable. In this paper, I try to review some of representative strike-slip basins along convergent margins, especially focusing on basin formation and filling processes, as the first step for comprehensive understandings of the tectono-sedimentary evolution in strike-slip basins.

Strike-slip faults
Much research has been published regarding the classification and terminology of strike-slip faults [e.g., [8][9][10][11][12]. Here, I use three tectonic settings as a basis for classifying strike-slip faults along convergent plate margins: subduction, continental collision, and plate-boundary transform zones.

Subduction zone
Subduction zones, where oceanic plates obliquely subduct underneath continental or island arc crusts, are sometimes accompanied by strike-slip faults separating elongate forearc slivers from continental margins or island arcs. This type of strike-slip fault is referred to as a trenchlinked [9] or trench-parallel [13] strike-slip fault ( Figure 1).
Trench-linked strike-slip faults lie parallel to the trench in the accommodating part of the trench-parallel component of oblique convergence of subducting plates [14][15][16]. The basic principle is considered to be that a trench-linked strike-slip fault is able to concentrate shear in a much more efficient way than distributing shear across the much larger and more gently dipping interface of the subducting plate. Therefore, the low dip angles and high friction These strike-slip faults sometimes separate a narrow sliver plate from the remainder of the over-riding forearc plate. Modern examples include the Median Tectonic Line in Japan, the Sumatra Fault in Indonesia, the Denali Fault in Alaska and the Philippine Fault in Philippine ( Figure 2). An ancient example can be observed in the Bering Sea [17]. The strike-slip faults in these settings are typically long (hundreds of kilometers) but occasionally segmented [9].
The Median Tectonic Line (MTL) active fault system is the longest and most active arc-parallel, right-lateral strike-slip fault system in Japan [e.g. 18,19]. The MTL extends over a distance of 500 km and accommodates the trench-parallel component of oblique subduction of the Philippine Sea Plate (Figure 2A).
The Great Sumatra Fault is a right-lateral strike-slip fault more than 1900 km in length ( Figure  2B). It is related to the northward subduction of the Australian Plate beneath the Sundaland Plate along the Java-Sumatra Trench ( Figure 2B) [20][21][22].
Strike-slip faulting in Alaska has involved several widely spaced major faults, with an overall seaward migration of activity: the Denali Fault was initiated in the Late Cretaceous or Paleocene and the Fairweather Fault in the Pleistocene ( Figure 2C) [23,24]. These right-lateral strike-slip faults are associated with the Alaskan subduction zone where the Pacific Plate subducts northwestward. Although current strike-slip movement takes place predominantly on the Fairweather Fault, the Denali Fault also shows some Holocene movement.
The Philippine Fault is a left-lateral strike-slip fault sandwiched between the Manila and Philippine trenches [25][26][27]. It completely traverses the Philippine archipelago and extends for more than 1000 km ( Figure 2D). Several strike-slip basins are developed along releasing bends or overstepped faults in this fault zone [28,29].   [33]. Plate names: EU, Eurasian; AR, Arabian; NU, Nubia (Africa); AT, Anatolian; SU, Sundaland; BU, Burma; YZ, Yangtze. Maps were drawn using SRTM and GEBCO with plate boundary data [30]. Black, red, and purple lines are plate convergent margins, plate-boundary transform faults, and indent-linked strike-slip faults, respectively. Blue arrows indicate the direction and velocity of plate motion (mm yr -1 ) relative to the Eurasian Plate based on [31]

Continental collision zones
Continental collision can cause crustal shortening and thickening by thrusting and escape or by extruding crustal blocks along conjugate strike-slip faults within the plate. These types of collision-related strike-slip faults between continental blocks are classified as indent-linked strike-slip faults ( Figure 1) [9].
Modern examples include several strike-slip faults in Turkey where the Arabian Plate is converging with the Eurasian Plate, and in southern China where the Indian Plate is colliding with the Eurasian Plate ( Figure 3). In the latter example, the collision originally formed the left-lateral Red River Fault associated with the southeastward extrusion of the Indochina Block [32]. After the propagation of the indent, the South China Block was extruded along the preexisting Red River Fault as a block boundary with right-lateral movement. . All maps were drawn by using SRTM and GEBCO with plate boundary data [30]. Black, red, and purple lines are subduction zones, oceanic spreading ridges, and plate-boundary transform faults, respectively. Blue arrows indicate the direction and velocity of relative plate motion (mm yr -1 ) based on [31]

Plate-boundary transform zones
Plate-boundary transform faults develop between two plates rotating around the poles that define the relative motion between them [9]. The San Andreas Fault, Dead Sea Fault, Alpine Fault (New Zealand), and the northern and southern margins of the Caribbean Plate are modern examples of this type (Figure 4).
There are many strike-slip basins along plate convergent margins ( Figure 5 and Table 1). Here I classify strike-slip basins into four types, discussed in turn below.

Fault-bend basins
Fault-bend basins result from vertical displacement of normal faults in front of releasing bends corresponding to gentle transverse R (synthetic Riedel) faults connected to stepped master Y (principal displacement) faults (Figures 1 and 6A). The basin geometry is generally spindleshaped or lazy-Z-shaped in plan view [38]. This type is considered to represent an early stage of the evolution of a pull-apart basin [12].

Stepover basins
As the master faults continue to propagate, they overlap and pull the crustal blocks farther apart, with lengthening geometries that gradually change from lazy-Z-shaped fault-bend basins to rhomboid-shaped stepover basins ( Figure 6B). The basins subside by extension along strike-slip fault systems where the sense of en échelon segment stepping coincides with the sense of the slip (i.e., right-stepping faults have dextral displacement). The term 'pull-apart basin' was originally introduced to explain a depression in the Death Valley whose sides were pulled apart along releasing bends or oversteps of faults [41]. According to the pull-apart mechanism, two sides of the basin are bounded by faults with primarily horizontal displacement, and the other two sides are bounded by listric or transverse faults.
Stepover basins generally maintain their length/width ratio [2], as expressed by the following relationship between the length (l) and width (w) of a pull-apart basin based on the dimensions of natural pull-apart basins ( Figure 6): The best fitting constants have been found to be c 1 = 1.0 and c 2 = 3.2, which yield l/w ≈ 3.2 with a 95% confidence interval about the ratio of 2.4 < l / w < 4.3.
In sandbox experiments [42], a spindle-shaped basin appears in the first stage of basin evolution and is bounded by master Y faults and their synthetic Riedel (R) faults. Subsequently, antithetic Riedel (R') faults replace R faults, leading to a rhomb-shaped basin. The l/w ratio depends on the angles α and β ( Figure 6): The mean angle between R and Y faults in the experiments is a = β = 30 o ; that is, l/w=3.5. This value is consistent with those of natural basins.
As overlapped offsets of the master strike-slip faults propagate, basins elongate and finally become long pull-apart basins. The Dead Sea Basin, with a length of 132 km and a width of 18 km (l/w=7.2), is considered to have been formed by the coalescence of three successive and adjacent sedimentary basins whose depocenters migrated northward with time [43]. Although each sub-basin has a l/w ratio typical of a pull-apart basin (2.4, 3.3, and 2.6 from south to north), propagation of the master fault, accompanied by the creation of a new stepover basin, has resulted in the basins high l/w ratio ( Figure 6C) [42].
Very high l/w ratios can also result from continuous transtension leading to extreme thinning of the crust and rupture. This process induces magmatic activity, high heat flow, and then the generation of new oceanic crust that may be younger than the overlying sedimentary succession (e.g., the Cayman Trough or the Gulf of California).

Fault-termination basins
Fault-termination basins are developed in transtensional stress domains at the ends of strikeslip faults where normal or oblique slip faults diffuse or splay off to terminate the deformation field [44]. If a part of a crustal block undergoes translation within the block, it results in shortening/uplift at one end and extension/subsidence at the other ( Figure 7). Basins formed by such subsidence are referred to as fault-termination basins or transtensional fault-termination basins [44].
Modern examples include the Yinggehai Basin (Song Hong Basin) along the Red River Fault zone [45], the Malay and Pattani basins in the Gulf of Thailand [46], several segmented basins in the Gulf of California [44,47], the northern Aegean Sea [48,49], and Beppu Bay along the Median Tectonic Line ( Figure 5 and Table 1) [50].

Transpressional basins
Transpressional basins tend to develop along oblique convergent margins whose subsidence results from flexural loading of the hanging-wall crust, similar to foreland basins adjacent to uplifted blocks [52][53][54]. Such basins are usually long, narrow structural depressions that lie parallel to the master faults.
The Sumatra forearc basins are modern examples of this type. Uplift of outer arc highs bounded by trench-linked strike-slip faults may cause flexural subsidence on the forearc side and generate elongate wedge-shaped sedimentary basins.

Examples of strike-slip basins
The

Geology
The Ridge Basin, which is one of the best-studied examples of a strike-slip basin [37,55], is situated along the San Andreas Fault, a right-lateral plate-boundary transform fault between the Pacific and the North American plates, and along the San Gabriel Fault, a major strand of the San Andreas Fault ( Figure 4B).
The San Gabriel Fault is a listric, ESE-dipping, oblique-slip fault rather than a subvertical, strike-slip fault [56]. The Ridge Basin is a type of fault-bend basin developed in front of a releasing bend on the San Gabriel Fault, along which the upper crust stretched and subsided to form a space in which sediments could be accommodated. The bottom of the basin is bounded by the subhorizontal San Gabriel Fault at a depth of ∼4 km.
The basin originated in the late Miocene as a narrow depression within the broad San Andreas transform belt in southern California. The basin has a length of 45 km and a width of 15 km; the length/width ratio of 3 is a typical value for pull-apart basins [2]. The strata are exposed as a northwest-dipping homoclinal sequence that becomes younger to the northwest. The exposed sediment thickness reaches ∼14 km, somewhat larger than the basin depth (∼4 km) [56].

Basin filling processes
The strata within the Ridge Basin are assigned to the Ridge Basin Group, which includes five formations: Castaic, Peace Valley, Violin Breccia, Ridge Route, and Hungry Valley ( Figure  8A). Sedimentation began in the late Miocene (ca. 11 Ma) with deposition of the marine Castaic Formation. The younger Castaic Formation is interfingered with the older Violin Breccia, which consists of conglomerates adjacent to the San Gabriel Fault scarp [57].

Geology
The Dead Sea Fault system is located along a plate-boundary transform zone that separates the Arabian Plate from the African Plate ( Figures 4C and 9) [12]. Movement along the Dead Sea Fault commenced in the Miocene in response to the opening of the Red Sea. The very low rate of relative plate motion between Arabia and Africa (6-8 mm yr -1 ) has yielded only 30 km of displacement during the past 5 Myr, and about 105 km of total offset during the past 18 Myr.
The Dead Sea Fault system includes both transpressional and transtensional domains ( Figure  9). Several strike-slip basins are developed along the steps of segmented faults in the transtensional domain, while the Lebanon and Anti-Lebanon ranges have been uplifted in the transpressional domain related to the restraining bend. The Dead Sea Basin is the largest strike-slip basin in the system, and is partly overlain by a deep hypersaline lake located at Earth's lowest continental elevation (418 m below sea level at the lake surface) [43,[62][63][64].
The Dead Sea Basin is 132 km long and 7-18 km wide, yielding a high length/width ratio (> 7). The basin is segmented into sequential sub-basins by deep transverse normal faults rather than by listric faults. The length of the basin is greater than the total offset length (∼105 km) of the fault system, which is atypical of pull-apart basins [65].
The basin has a cross-sectional asymmetry, with a steep eastern slope and a gentle western slope. Seismic refraction and gravity data indicate that the southern Dead Sea Basin is unusually deep, containing about 14 km of sedimentary fill [66]. Geophysical data suggest that the deep basin is probably bordered on all sides by vertical faults that cut deep into the basement [67]. The thick sediment accumulation yields a large negative Bouguer gravity anomaly (lower than -100 mGal) [64]. Given the depth of the basin, ductile deformation in the lower crust would be expected; however, the present-day heat flow in the Dead Sea Basin is low (∼40 mW m -2 ) [68], suggesting that the lower crust may still be cool and brittle, and that the Moho is not elevated beneath the basin. These inferences are consistent with seismic activity at depths of 20-32 km.
The Dead Sea Basin has traditionally been considered a classic example of a stepover basin [2], but other interpretations have been proposed, including propagating basins [67], stretching basins [64], and sequential basins [63]. The sequential basin model, in which several active subbasins are delimited by boundary master faults and transverse faults, and simultaneously become larger and deeper as the master faults propagate, could explain why the Dead Sea Basin is longer than the total amount of slip along the Dead Sea Fault.

Basin-filling processes
The depositional environments of the Dead Sea Basin are affected by the arid climatic conditions, with the area having an average annual rainfall of 50-75 mm. The modern sediments are transported to the basin mainly from the north by the Jordan River and from other directions by marginal tributaries. The mean annual discharges from the north, east, west, and south are 1100, 203, 4-5, and 4 mm, respectively [71].
In the middle to late Miocene, fluvial clastics of the Hazeva Formation were deposited in the southern sector of the basin (Figure 9). The formation consists of fluvial sandstones and conglomerates, including pre-Cretaceous components, transported from distant sources south and southeast of the Dead Sea Basin [43,64]. During the Pliocene, the evaporitic Sedom Formation accumulated in estuarine-lagoonal environments in the Dead Sea basin; the formation consists mostly of lacustrine salts, gypsum, and carbonates interbedded with some clastics, and is found in the central sector of the basin. The 2-3 km thickness of this evaporitic formation may have formed in < 1 Myr; therefore, the sedimentation rate was probably higher than 2 m kyr -1 .
In the Pleistocene and Holocene, fluvial and lacustrine deposits, alternating with evaporites and locally sourced clastics, accumulated in lakes that post-date the formation of the Sedom lagoon. The Amora, Lisan, and younger formations consist of laminated evaporitic (gypsum) and aragonite sediments that continue to accumulate in the modern Dead Sea in the northern sector of the basin. The average sedimentation rate in this stage reached 1-1.5 m kyr -1 [64]. On the whole, the depocenters have migrated northward since the Miocene.
The margins of the Dead Sea are dominated by alluvial fans. The modern basin margin environments consist of (1) talus slopes, (2) incised and confined stream channels, and (3) coarse-grained and relatively high-gradient alluvial fans. In contrast, sediments in the offshore environment are composed of thick sequences of evaporitic salt intercalated with thin beds of laminated aragonite and detrital silt [72].

Geology
The Yinggehai Basin (Song Hong Basin) is an example of a fault-termination basin, and is situated at the southeastern end of the Red River Fault zone (RRFZ) [73,74]. The RRFZ, extending for some 1000 km, separates the South China Block to the north from the Indochina Block to the south (Figure 10), and is considered to be related to the continental collision between the Indian and Eurasian plates [e.g., 75]. The formation of the Yinggehai Basin was controlled by the successive clockwise extrusions of the Indochina Block and the South China Block.
The RRFZ was a sinistral strike-slip fault in the first stage of its evolution , associated with ductile deformation [76] and the creation of an unconformity in the Gulf of Tonkin (the offshore part of the Hanoi depression [77]). After a quiescent stage from 17 to 5 Ma, due to a slowdown in the clockwise rotation of the Indochina Block [78], the movement along the RRFZ became dextral [45,79,80]. The right-lateral shearing is indicated by geomorphic fault traces and large river offsets [81][82][83], as well as GPS observations [80]. This change from a sinistral to a dextral sense of movement supports the basic tenets of the two-phase extrusion model; namely, the early, collision-driven escape of Indochina towards the SE, and the subsequent change to accommodate the present-day escape of Tibet and South China [78,84,85].
The Yinggehai Basin, situated in the offshore extension of the RRFZ, is 500 km long and 50-60 km wide (l/w ≈ 10); it is oriented SE-NW and is located offshore between Hainan Island to the east and the Indosinian Peninsula to the west [73]. The basin formed originally as a sinistral strike-slip basin [77,86], but developed into a dextral strike-slip basin after the change in the sense of fault displacement of the RRFZ [87,88]. The basin subsided by simple shear on lowangle, detached normal faults of the upper crust and by pure shear of the lower crust [45,87].

Basin-filling processes
The Yinggehai Basin is filled with 10-17-km-thick clastic deposits [45]. Basin sedimentation began in the late Eocene [45,88]. During the Oligocene (> 21 Ma), clockwise rotation of the Indochina Block induced sinistral slip along the RRFZ and the basin rapidly subsided and started to fill. The depocenter was situated in the southern part of the basin during this synrift stage ( Figure 11) [45,91]. During the quiescent post-rift stage of the RRFZ, the rate of basin subsidence accordingly decreased and the depocenter gradually migrated northwestward [45]. The reactivation of the RRFZ with dextral movement triggered rapid subsidence [80] and enhanced the input of sediment [45]. The depocenter migrated from the center to the southeastern end of the basin (Figure 11).
The infill of the Yinggehai Basin varies from alluvial, fluvial, and lacustrine deposits (before 21 Ma, the syn-rift stage) to marine sediments (after 21 Ma, the post-rift stage) [80]. Almost all the sediments in the basin are considered to have been derived from the Himalayas through the Red River drainage network [92]. Large mountain belts with high rates of sediment yield and along-fault transport networks were able to effectively supply huge volumes of detritus into the basin. Thick sediments and high sedimentation rates resulted in an over-pressured condition leading to mud diapirism [73], and also to depressed surface heat flow (∼80 mW m -2 ) [93]. The orientations of mud diapirs in the basin ( Figure 13) indicate E-W extension related to right-lateral motion of the northeastern bounding fault [89,90].
The Pattani and Malay basins in the Gulf of Thailand (inset map in Figure 10) are also considered to be fault-termination basins related to the continental collision of the Indian Plate [94,95,96]. Sediment supply into the basins is dominated by rivers flowing along the strikeslip faults [97,98]. The Pattani and Malay basins contain thicknesses of sediment of more than 8 and 14 km, respectively. The subsidence of the basins was controlled at first by tectonic depression related to strike-slip deformation and then by thermal subsidence due to high surface heat flow (100-110 mW m -2 ) [46,[99][100][101].

Geology
Sumatra is a classic example of slip partitioning due to an obliquely subducting plate [14]. The Indian and Australian plates are subducting beneath the Sundaland Plate in southeastern Asia along the Java-Sumatra Trench (Figure 2), where oblique subduction is accompanied by trench-parallel forearc translation [21]. The subduction thrust and the trench-parallel strikeslip fault (Great Sumatra Fault) isolate a wedge of forearc in the form of a sliver plate (the Burma Plate). The Great Sumatra Fault extends along the entire length of Sumatra Island (>1900 km) [22] and finally joins the West Andaman Fault (WAF), which constitutes a series of transform faults and spreading centers in the Andaman Sea [102]. In the forearc sliver, outerarc uplift related to development of the accretionary prisms occurs on the trenchward side, and forearc basins (the Aceh and Simeulue basins) are developed on the landward side.
The Aceh Basin (Figures 12 and 13 [22,102]. The enclosed area is shown in Figure 13. Bathymetry is based on using SRTM and GEBCO with the data recently collected by [108][109][110]. 106]. The bounding fault is a right-lateral transpressional fault with accompanying topographic highs of en échelon anticlines on the western margin and compressional ridges of Tuba Ridge on the southern margin ( Figure 13). The transpressional uplift generating the outer-arc high by a thickening crustal block may be resulting in subsidence opposite to the high. This type of sedimentation is similar to that of foreland basins, where depressions are caused by the overburden pressure of the thrusted crust. Rifting and basin formation started in Sumatra during the Paleogene [107].

Basin-filling processes
The deposits in the basin are thickest along the boundary fault between the basin and the outerarc high, and gradually thin with increasing distance from the faults ( Figure 14B and C).
Regarding the recent deposits, represented by seismic units 3 and 4 ( Figure 14), unit 3 sediments in the southern part are thicker than those in the northern part, but unit 4 sediments are thicker in the northern part. Therefore, the main depocenter is considered to have migrated from the south (unit 3) to the north (unit 4). This interpretation is supported by seismic profiles of [106], who noted that the southern part of the Aceh Basin is raised above the northern part.
Most of the sediments are considered to have been supplied from Sumatra Island through small submarine channels ( Figure 13). However, little is known about axial sediment redistribution within the basin. Figure 13. Detailed bathymetry around the Aceh Basin. Red and yellow lines are strike-slip faults and axes of anticlines [105], respectively. Thick solid lines in the Aceh Basin mark cross-sectional profiles shown in Figure 14. Abbreviations are the same as for Figure 12. Bathymetry is based on using SRTM and GEBCO with the data recently collected by [108][109][110].

Figure 14.
Cross-sectional profiles of the Aceh Basin. The uppermost seismic unit 4 was deposited predominantly in the northern part of the basin, suggesting a northward migration of the depocenter. Modified from [103]

Summary
This preliminary review has introduced some of the representative strike-slip basins at convergent margins from the viewpoints of basin formation and filling processes. Because strike-slip basins present a wide range of formational processes and sedimentary facies, it is difficult to establish a simple model of their evolution. To understand both modern and ancient strike-slip basins, the following factors need to be considered: • Compressional uplift along master faults is probably needed for the dominance of axial sediment supply into the basin, which has the potential to produce and distribute huge volumes of detritus ( Figure 15). Restraining bends as paired bends [12] such as in the Dead Sea Basin setting, and collision related to continental indentation such as in the Yinggehai Basin setting, are possible cases where axial sediment supply is enhanced. Conversely, a marginal high along the master fault is required for the formation of transpressional basins such as the Aceh Basin; therefore, marginal sediment supply may tend to dominate in such basins.
The continuous migration of depocenters requires that the progressive displacement of the master faults creates new accommodation space ( Figure 15). In standard models of pull-apart basins, which are bounded by steep master faults and listric transverse faults, increasing the offset leads to a widening of the fault zone, resulting in wider pull-apart basins with a l/w ratio of about 3 for each basin [2,42]. Therefore, for large l/w basins with continuous depocenter Depocenter fixes along the releasing bend result from the right-lateral migration of sediments deposited on the footwall. A transpressional component would be required to generate the sediment source, and en échelon folds may form along the master faults. Both models generate deposits with axial sediments whose thicknesses are greater than the burial depths.
migration, such as the Dead Sea Basin and the Izumi Group, other formational mechanisms should be considered instead of the pull-apart basin model.
Step-wise propagation or sequential progradation of the master fault could produce elongated stepover basins. A high l/w ratio could also potentially be produced by a transpressional basin along a trench-linked strike-slip fault. However, there is a need to establish physical models for basin formation in such settings.