What controls how long a sedimentary basin is active

Strike-slip faults can accommodate localized compression or extension at continental margins, in island arcs, and also within continents. Sedimentary basins commonly develop where the fault kinematics are divergent with respect to the plate vector along strike-slip faults. Since the s, various classifications of strike-slip basins have been formulated [4, 11, 36—40].

Common characteristics of strike-slip basins [ 4 , 39 ] include: 1 elongated geometry, 2 asymmetry of both sediment thickness and facies pattern, 3 dominance of axial infilling, 4 coarser-grained marginal facies along the active master fault, 5 finer-grained main facies, 6 depocenter migration opposite to the direction of axial sediment transport, 7 very thick strata relative to the burial depth, 8 high sedimentation rate, 9 abrupt lateral and vertical facies changes and unconformities, 10 compositional changes that reflect horizontal movement of the provenance, 11 abundant syn-sedimentary slumping and deformation, and 12 rapid subsidence in the initial stage of basin formation.

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.

Numbers correspond to those in Figure 5. Strike-slip basins at plate convergent margins. Red triangles, trench-linked; black squares, indent-linked; purple circles, plate-boundary transform faults. Numbers correspond to those in Table 1. Geometrical models of A a spindle-shaped fault-bend basin and B a rhomb-shaped stepover strike-slip basin.

Colored areas indicate subsiding basins. C Multistage evolution of stepover basins. Diagrams are modified from [ 42 ]. 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 6 A.

The basin geometry is generally spindle-shaped 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 ].

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.

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. 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.

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. 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. Fault-termination basins are developed in transtensional stress domains at the ends of strike-slip faults where normal or oblique slip faults diffuse or splay off to terminate the deformation field [ 44 ].

Basins formed by such subsidence are referred to as fault-termination basins or transtensional fault-termination basins [ 44 ]. 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—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.

Typical geometry of termination areas of strike-slip faults. A If the block collides with a rigid continental crust, it shortens and is uplifted, accompanied by thrusts.

At the opposite side to the uplift, upper crust is mechanically pulled away, leading to subsidence. B If the block extrudes with a rotational component into a weak oceanic crust a transtensional setting , a sedimentary basin forms at the end of the strike-slip fault.

Examples include the Yinggehai Basin, which is related to the extrusion of the Indochina Block, and the Gulf of California, which is related to the transtensional movement of the Baja California Peninsula.

C The strike-slip fault diffuses its displacement through splayed extensional normal faults at its end [ 44 ]. An example is the Cerdanya clastic basin formed by late Miocene normal faulting at the termination of the La Tet strike-slip fault, Spain [ 51 ]. The wide variability of strike-slip faults makes it difficult to develop a simple model of the formation of strike-slip basins and their sedimentary facies.

Although the geometries of such basins depend on the amount of fault displacement, the angle and distance between overstepped faults, and the depth of detachment of the faults, the basins are generally elongate, narrow, and deep.

Several representative examples of the strike-slip basins described in this section show a range of basin evolutionary paths and filling processes. 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. A Simplified geological map showing formations in the Ridge Basin [ 58 ]. B Conceptual basin-filling process for the Ridge Basin [ 8 ]. C Cross-sectional profiles showing continuous axial sediment supply and migration of sediments with relatively fixed depocenters [ 61 ].

The basin originated in the late Miocene as a narrow depression within the broad San Andreas transform belt in southern California.

The strata are exposed as a northwest-dipping homoclinal sequence that becomes younger to the northwest. Sedimentation began in the late Miocene ca. The younger Castaic Formation is interfingered with the older Violin Breccia, which consists of conglomerates adjacent to the San Gabriel Fault scarp [ 57 ]. The Peace Valley Formation consists mainly of sandstone and mudstone of lacustrine, fluvial, deltaic, and alluvial facies, accompanied by minor carbonaceous deposits.

The Ridge Route Formation, which crops out in the northeastern part of the Basin, is composed of alluvial sandstone and conglomerate, and is interfingered with the Peace Valley Formation.

The deposition of this formation, including alluvial conglomerate, sandstone, and mudstone, ended at ca. The releasing bend may have a paired restraining bend on the northwestern side of the fault. Within the restraining bend, highlands were formed, which in turn provided sediment to be transported into the basin.

Most of the sediment filling the basin was carried by rivers draining source areas located to the northeast. The sediments forming the Ridge Basin Group were deposited at a rate of about 2 m kyr The right-lateral displacement of the San Gabriel Fault carried the basin, together with the sediments, southeastward, resulting in a northwestward migration of the depocenter and successively younger beds onlapping onto the basin floor Figure 8C [ 58 , 59 ].

More than 45 km of lateral displacement is estimated, based on the distribution of the Violin Breccia. This displacement, and basin migration, ended in the early Pliocene. 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 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.

On the northern side of the basin, the Lebanon and Anti-Lebanon ranges were uplifted in a transpressional domain. The locations of faults are taken from [ 12 ]. The basin is segmented into sequential sub-basins by deep transverse normal faults rather than by listric faults. 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 — mGal [ 64 ]. 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 sub-basins 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. 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 , , 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.

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. 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.

Since fracture, flow and other non-linear phenomena seen in the general solid material are not considered in the model, it is difficult to directly compare the amount of displacement between the modeled structure and the actual structure in long-time scale modeling. In general, dislocation modeling including visco-elasticity effects is often employed in discussions on crustal movements over a long time-scale such as on a geological time scale e.

As mentioned above, dislocation modeling defined using a range of the linear elasticity has disadvantages in the ability to directly compare the amount of displacement between the modeled structures and the actual structures. However, when we simply discuss the essential aspects of tectonics from the distribution pattern of structures caused by fault motions, dislocation modeling is a very useful tool because it provides the pattern of displacement using easy calculations.

The aims of this study are to simplify the complex formation processes of sedimentary basins through numerical simulations and to show that the simplification enables us to estimate deductively which processes cause materialization. Central Hokkaido was selected as the field in which to achieve these aims, as many sedimentary basins are distributed in this area. As will be described later in the paper, many sedimentary basins were formed from 48 to 12 Ma in central Hokkaido, and it is difficult to discuss their formation processes using only observational data because of their complex distribution both at and below the surface.

In order to simplify the complex formation processes of sedimentary basins, we attempted to restore sedimentary basins using the advanced technique already mentioned, and we evaluated the fault type and the amount of movement required to form these sedimentary basins. In the following sections, we describe the basic background and gravity anomaly in central Hokkaido, and we attempt the restoration of the sedimentary basins.

Hokkaido is located on the North American plate, at a junction of the Northeast Japan arc and the Kurile arc Figure 3. Using recent GPS observations, an east-west compressive strain field has been observed in the northern part of Hokkaido, and this strain field is considered to be caused by the convergence of the Eurasia plate with the northern part of Hokkaido functioning as a part of the plate boundary e.

Location map of our study area. Hokkaido is located on the North American plate, at a junction of the Northeastern Japan arc and the Kurile arc. Gray dashed line indicates the old plate boundary between the Eurasian and North American Plates.

Rectangular area by gray thin line indicates the study area of Itoh and Tsuru [ 58 ]. Although the present plate boundary between the North American plate and the Eurasian plate exists in the Sea of Japan, it is known that the plate boundary was located in central Hokkaido at around 13 Ma. This period of time corresponds to the stage when the uplifting of the Hidaka Mountains began e. This tectonic framework is controlled by the dextral oblique collision between the Eurasian and North American Plates and the oblique subduction of the Pacific Plate beneath the Kurile Trench.

It is considered that the Kurile arc migrated into the southwestward as a forearc sliver by the oblique subduction of the Pacific Plate and that the Hidaka Mountains would be formed by collision of the Kurile arc and the Northeast Japan arc e.

Since, consequently, it is an important area for understanding characteristics and mechanism of collision zone, numerous geophysical surveys e. Using these surveys, subsurface structures which indicate a collision between the Northeast Japan arc and the Kurile arc have been obtained, and have contributed to tectonic discussions. However, in contrast, geophysical studies in the sedimentary basin area in central Hokkaido are limited in number.

The geological characteristics in Hokkaido are that Cenozoic strata consist of island-arc-trench systems of the Northeast Japan arc and Kurile arc, and that each Cenozoic strata distributed in the Northeast Japan arc and the Kurile arc appear in the western half and eastern half area of Hokkado, respectively. This characteristic also appears in Neogene and Quaternary strata, volcanoes and their products, and topography e.

Figure 4 shows the distribution of the Paleogene strata in a north-south direction in central Hokkaido. This distribution traces the old plate boundary. This N-S elongation area is included in the Ishikari-Teshio Belt that is underlain by the Cretaceous Yezo Group, and is regarded as a typical sequence in a forearc basin setting [ 55 ].

It is known that the Paleogene sedimentary strata were deposited during the early Eocene and Oligocene in almost the entire region e. Distribution of the Paleogene strata. Green and blue areas indicate distribution areas of the Paleogene sedimentary layer under and on the surface, respectively.

After Kurita and Hoyanagi [ 56 ]. In the study area, sedimentary basins and sedimentary layers were formed during 48—12 Ma and have been complexly distributed.

The Ishikari stage is divided into early 48—45 Ma and late 45—40 Ma stages. The shape of the sedimentary basin and the distribution of sedimentary layers are shown in Figure 5. It is known that the Ishikari Group differentially subsided and was then divided into several components [ 57 ]. Based on detailed sedimentological studies, Takano and Waseda [ 57 ] also points out that the rate of subsidence accelerated during deposition of the Ishikari Group.

The Ishikari stage is the sedimentation stage of the Ishikari Group, corresponding to the Eocene and is divided into early and later stages according to the sedimentation style. Sediments in the early Ishikari stage are distributed shallowly and widely Figure 5A.

Sediments in the later Ishikari stage are distributed deeply and narrowly Figure 5B. Itoh and Tsuru [ 58 ] identified a NNW-SSE trending deformation zone bounded by large transcurrent faults including T1 and T2 later describing from seismic reflecting data in the northern part of the Northeast Japan forearc Figure 3 and their right lateral motions have been indicated by the clockwise rotation of Paleogene marine sediments and by paleogeographic reconstruction.

Since, as already mentioned, the present study area the western half of Hokkaido has same Cenozoic strata distributed in the Northeast Japan arc, study area of Itoh and Tsuru [ 58 ] and our study area are geologically continuous in the Paleogene time. Consequently, it is expected that a right lateral motion of the crust was dominant. This stage corresponds to the Eocene and the early Oligocene.

It is expected that a right lateral motion was dominant in this stage, because right lateral motion was also dominant in the Eocene and the late Oligocene see below.

This stage corresponds to the late Oligocene and the early Miocene. Kurita and Yokoi [ 60 ] also stated that lateral faulting was dominant in forming some of the tectonic structures during the late Oligocene. During the Neogene, Japan was affected by the opening event of the back-arc basin of the Sea of Japan. As mentioned above, a lateral motion of the crust was dominant during the early Neogene [ 37 , 38 , 59 ].

Although a building of the Hidaka Mountains in around 13 Ma has been pointed out e. Numerous geological and geophysical surveys have been carried out in the Hokkaido area, and each survey has played an important role in the understanding of crustal characteristics and tectonic events in the area.

In particular, seismic prospecting has proved very useful in obtaining information relating to subsurface structures. However, seismic prospecting is almost two-dimensional, and it is difficult to intuitively understand the subsurface structures as three dimensional structures, even when provided with data from more than one profile.

In contrast, the characteristics of gravity anomaly maps are easy to interpret and can be used to roughly estimate three dimensional subsurface structures from the data. Figure 6 shows the Bouguer gravity anomaly map of the study area. This map is based on the gravity mesh data by Komazawa [ 61 ]. Bouguer gravity anomaly map.

Contour interval is 10 mGal. There are negative gravity anomalies in the southern and northern parts of central Hokkaido. The southern negative gravity anomaly located at the curved subduction zone is the lowest in the country. From seismic prospecting, it is known that this negative gravity anomaly consists of a very thick sedimentary layer 5—8 km , with a velocity of 2.

The sedimentary layer was formed by imbrications associated with the collision process of the Northeast Japan arc and the Kurile arc e. In contrast, there is a positive gravity anomaly in the area of the mountains, and the mountain elevations are roughly less than m.

It would not be necessary to consider isostasy for the mountains, because the mountain elevations are not very high and the gravity anomaly in this area is positive. Gray indicates gravity low area less than 20mGal. A flat gravity anomaly of less than 20 mGal is distributed like a belt between the northern and southern gravity anomalies Figure 6 and Area II in Figure 7. Figure 7 is the gravity anomaly map that the area less than 20 mGal was painted by gray.

This painted area corresponds to the area where the Paleogene strata distribute under the surface Figure 4. From these profiles, the gravity anomalies in the region are shown to have the characteristics as follows:. Gravity anomalies in the west-east direction have a regional trend which tilts toward the east. This could indicate the regional gravity field in Hokkaido. Gravity anomalies less than 20 mGal have a steep gradient on the east side, while those on the west side vary gently.

Gravity anomalies in the north-south direction are relatively high and flat at the center of Hokkaido. It is possible that the high density of metamorphic belts near this region affect the observed gravity anomalies.

Another cause to be considered could be the effect of subsurface structures such as a reduction of low density materials e. In general, gravity anomalies are caused by spatial variations of subsurface structures, and indicate a deficiency or an excess of mass under the surface. In general, high gravity indicates the existence of a mass excess or of high density materials, and low gravity indicates the existence of a mass deficiency or of low density materials. These deficiencies or excesses, of mass can be evaluated quantitatively using Gauss's theorem e.

Here, g x, y is the gravity anomaly data given on xy mesh with a constant interval. Equation 1 is described by an infinite integration and it is difficult to perform an infinite integration with actual field data. Consequently, we understand this as being an approximate calculation and perform a numerical integration within a finite area S as follows:. We applied equation 2 to three areas, I, II and III, and we attempted to estimate the magnitude of mass deficiency for the formation of a gravity anomaly less than 20 mGal in each area.

In the calculations, we employed the Gauss-Legendre numerical integral formula e. As a result, mass deficiencies of 4. There are large mass deficiencies in areas I and III, where the negative gravity anomalies observed are very large and a small mass deficiency in area II. In central Hokkaido, the amount of mass deficiency is different by about two digits in both the maximum and the minimum values. It is well known that if there is a tectonic line or a fault with a large gap in the vertical direction, the spatial distribution of the gravity anomaly varies steeply around these structures.

In general, the first derivative of the gravity anomaly is more practical, because the calculation used is very simple and the geophysical and geological interpretations for the calculated results are straightforward. We employed the first derivative of the gravity anomaly defined by the following equation 4 , and calculated the horizontal gradient of the gravity anomaly Figure 8 :.

Although there are no continuous horizontal gradient anomalies within the area where the gravity anomaly is less than 20 mGal, the continuous horizontal gradient anomalies appear around this area. If the area for modeling is small, or if the tectonics and faults assumed for modeling are clear, such a modeling procedure is useful and practical e. The faults and their displacements appropriately assumed can then be considered as an initial model and can then be corrected by trial and error, so that the calculated results fit to the actual structures or their distribution pattern.

There are numerous small faults in central Hokkaido. As mentioned above, the details of tectonics and faulting in this area are unclear. It would, therefore, be impossible to attempt to model each fault for restoring the distribution of the sedimentary basins by trial and error.

Consequently, in this study, we assumed that the dislocation plane used for the modeling was not a fault plane, but a typical or average plane of a fault zone. After trial and error, we defined the nine fault zones as shown in Table 1 and Figure 9 , and employed them for numerical simulations. Each fault included in these fault zones is listed in Table 1 with literature. Fault zones defined in this study.

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