Faculty of Science Masaryk University

Department of Geological Sciences

Evolution of Extensional Systems

Literature Thesis in Requirement for Masters

Author: B.Sc. Cameron Sheya Supervisors: Assoc. Prof. RNDr. Rostislav Melichar, Dr., Prof Michal Nemˇcok, DrSc.

Brno, Czech Republic 2010 Contents

Table of Contentsi

Preface ii

1 Classification of rift systems: Basics1 1.1 Continental Rifts...... 1 1.1.1 Narrow Rifts...... 2 1.1.2 Wide rifts...... 2 1.2 Pull-apart basins...... 5 1.3 Back-arc Basins...... 10 1.4 Cratonic Sags...... 11

2 Rift system linkages and geometries 16 2.1 Fault geometries and linkages...... 16 2.2 Fault linkages...... 18 2.3 Sub-basin interactions and the role of extension geometry in the process.. 22

3 Pull-apart basin development and geometries 28 3.1 Basic basin evolution and resulting geometries...... 28

Bibliography 38

i Preface

The purpose of this literature thesis is to create a concise report on the extension and associated basin types and their subsequent history in the orogenic foreland setting. This thesis was written with the intent of first classifying the types of extensional basins and then describing their evolution.

The understanding of the basin dynamics throughout their history will help in interpreta- tions of past and present day tectonic regimes. Final chapters discuss characteristics and dynamics of foreland basins including potential inversion of pre-existing extensional basins.

ii Chapter 1

Classification of rift systems: Basics

1.1 Continental Rifts

Continental rifts are areas where extension of the continental lithosphere leads to a subsid- ing basin.It is broken by several normal faults or several sets of normal faults, which can lead to either passive or active upwelling of the hot asthenospheric material underneath the area of the rift (Allen & Allen, 1990, 2005; Brun, 1999). Good examples of continental rifts come from the Basin and Range Province in the Western United States (Hamilton, 1987) and East African rift (Rosendahl, 1987).

Extension within these continental rift systems continuously evolves until subsequent re- gional tectonic process either changes the existing local tectonic regime into another, (e.g. change from rifting to thrusting) or some process halts the extension, (e.g. the lack of sufficient stresses to control faulting) (Withjack et al, 2002).

Continental extension has several associated factors such as:

• high heat flow in the extended basin (Allen & Allen, 1990; Behn et al, 2002); • high temperature gradient from the basin to the un-extended margins (Allen & Allen 1990); • negative gravity anomaly (Prieto, 1996, Allen & Allen 2005); • internal structural highs, which may sometimes become inverted sedimentary prove- nances for deposition in the basin (Allen & Allen, 1990); • possibility associated magmatism due to the upwelling of hotter material from the asthenosphere (Behn et al, 2002); 1 1.1 Continental Rifts 2

• if the basin itself has emplaced basalt or granite, it will result in a positive magnetic anomaly (Prieto, 1996) and; • changes in isostacy.

Continental rifts can be classified into narrow and wide rifts (Allen & Allen, 1990; Brun, 1999). Narrow continental rifts are areas of extension in the lithosphere that result in a stretching instability when the geotherm is higher than normal and the crustal thickness is in the range of 30-40 km whereas wide rifts are wider than the thickness of the lithosphere in which they occur (figure 1.1 ; Brun, 1999).

1.1.1 Narrow Rifts

Narrow rifts are typically found in areas that have thermal anomalies which are associated with:

• previous thinning of the lithosphere; • in areas with high mantle strength (Gueydan et al., 2008); • localized magmatism either from tectonic reasons or the influence of a hotspot; and/or • regions of localized strain weakening that affects the stronger layers of the continental lithosphere (Buck, 1991).

Passive margins are incorporated into the classification of narrow rifts due to the thickness of lithosphere in which they generally evolve. Pre-rift rheology and thermal conditions of the crust play an integral role in determining the type and geometry of the rift. Passive margins are similar to the aforementioned continental rift systems, however; the upwelling of the Moho and asthenosphere is more pronounced in the formation of a passive margin. This upwelling of the Moho and asthenosphere usually has a longer wavelength than the width of the rift itself. Passive margins usually have a width range from around 100 km to 400 km (Brun, 1999). Examples of this rift type are found on the entire coastal margin of Africa, eastern margin of North America and the coastal margins of Australia among other coastal regions throughout the world (figure 1.2).

1.1.2 Wide rifts

Wide rifts are wider than the thickness of the lithosphere in which they occur (Allen & Allen, 1990; Brun, 1999). The extensional event in the wide rift settings take place 1.1 Continental Rifts 3

(a)

(b)

Figure 1.1: (a) Idealized narrow rift with examples of passive margins and oceanic rift stages; (b) wide rift system, note how the brittle lithosphere has been thickened due to a previous compressional event (Brun, 1999) 1.1 Continental Rifts 4

Figure 1.2: Global distribution of passive margins. Compiled from Melluso et al., 2002; Geoffroy, 2005; Watkeys, 2002; Leroy, 2008. 1.2 Pull-apart basins 5 after a period or multiple periods of lithospheric thickening that can be associated with a subduction or area of collision setting. Wide rifts are known to reach widths of up to 1000 km (Brun, 1999) and are usually separated into sub-basins. An example of a wide rift that has been separated into sub-basins is the Basin and Range Province in the Western United States. This rift contains several tilted horst blocks that separate the individual sub-basins (Hamilton, 1987). Wide rifts are usually characterized by a horst-and-graben architecture and the frequent occurrence of core complexes (Hamilton, 1987; Brun, 1999). The models of Gueyadan et al., (2008) have shown that wide rifts occur in areas where the lower crust is very ductile and weak which leads to a more wide spread fracturing of the brittle crust.

Core complexes occur in areas with a very low angle normal detachment fault that allows the hotter and more ductile material to protrude through the tilted blocks of a wide rift sys- tem (Allmendinger and Platt, 1983; Brun et al., 1994; Brun, 1999). When a core complex is forming, the extension of the upper crust is largely localized whereas the extension of the lower crust is much more broadly distributed (Allmendinger and Platt, 1983). The ductile material of the lower crust is then exhumed and exposed by a system of low angle normal faults. The exhumed material is then eroded and uplifted which leads to the exhumation of highly metamorphosed rocks (Brun, 1999). These core complexes tend to occur in areas that have larger than normal heterogeneities in the crustal lithosphere (Brun, 1999).

Wide rifts occur by definition over a large area, which accounts for large differences in the amount of strain throughout the whole basin (Hamilton, 1987). Therefore, there can be a multitude of different fault expressions, including low angle detachments of core complexes, and high-angle normal faults between the horsts and grabens, scattered throughout the basin as a whole (Allmendinger and Platt, 1983; Brun et al., 1994).

1.2 Pull-apart basins

Pull-apart basins are structural depressions that are formed by localized zones of extension within an area of strike-slip deformation (figure 1.3 (a) ;(b); 1.4; Carey, 1958, Rahe et al., 1998). The areas of extension occur in the accommodation zone between two strike-slip branches that are nearing each other (Rahe et al., 1998; Sims, 1999). As the area between these two main controlling faults continues to become fractured, the subsequent linking and accommodation space created by this movement will create a subsiding extensional basin (Allen & Allen 1990). The resulting shape from the localized extension is usually rhombic (figure 1.3 (a) ;(b); 1.4; Price and Cosgrove, 1990; Rahe et al., 1998; Sims et al., 1999).

The process of fault linking and subsequent extension and subsidence evolves as follows:

• R shears begin to form throughout the accommodation zone between the two linking 1.2 Pull-apart basins 6

(b) (a)

Figure 1.3: (a) Model set up from Rahe (1999) showing the initial parameters of a pull- apart system; (b) after deformation of the model in (a) showing the relative movements and area of accommodation (Rahe, 1999).

strike-slip faults. The R shears have the same sense as the overall strike-slip regime (figure 1.5 (a); (b), Allen & Allen, 1990; Twiss & Moores, 1992).

• as the basin continues to evolve, R’ shears begin to propagate through the accom- modation zone. They have a displacement that is opposite to that of the overall strike-slip sense (figure 1.5 (a) ;(b); Allen & Allen, 1990; Sims et al., 1999).

• subsequently, P shears begin to develop through the accommodation zone, having the same sense as the overall movement of the strike- slip system (figure 1.5 (a) ;(b); Allen & Allen, 1990). Tension fractures and normal faults are being developed inside the pull-apart basin apart from the mentioned R, R’ and P shears (Allen & Allen 1990).

• the final stage of linkages in a pull-apart basin is when the various fractures and faults hard-link the two controlling strike-slip faults together (Allen & Allen 1990).

The evolution of a pull-apart can be divided into three stages (Rahe et al., 1998):

1. incipient;

2. early; and

3. mature.

The incipient stage of a pull-apart is characterized by the initiation of graben or half graben which is oriented parallel to the oblique step located between the controlling strike-slip faults (figure 1.5 (a) ;(b);1.6; Rahe et al., 1998; Sims et al., 1999). 1.2 Pull-apart basins 7

Figure 1.4: Sandbox modeling of a pull- apart basin; (a) 0.75 centimeters of deformation in map view; (b) 5.0 centimeters of deformation with white sand powdered on top to show contrasts; (c) cross section detailing the fault structures (Sims et al., 1999) 1.2 Pull-apart basins 8

(b) (a)

Figure 1.5: (a) Riedel Shears along with the associated P and R shears within a pull apart basin (b) manifestations of the R’, R and P shears within a pull-apart basin.

The initial internal geometry of the pull-apart basin as well as its overall shape are dictated both by the displacement and relative displacement rates across the strike-slip faults (Rahe et al., 1998; Sims et al., 1999).

The evolution of this graben or half graben accumulates a large amount of additional normal faults that are forming more towards the basin center rather than on the basin margins. This is analogous to a rift system that forms additional normal faults towards the basin center accommodating the amount of extension. Normal faults of the pull-apart basin are dipping mostly toward the basin center. However, in certain circumstances with differing rheologies as well as brittle/ductile coupling and extension rates, antithetic faults have been observed (Rahe et al., 1998).

The mature stage of a pull-apart basin is characterized by the hard linking of the bounding normal faults of the extensional basin with the controlling strike-slip faults (Rahe et al., 1998). Such basins are fully developed and undergo formation of normal and antithetic faults as the extension progresses. Detachments also play a role in the evolution of a pull- apart basin. They can be located either in the ductile or brittle environments (Sims et al., 1999). The location of the detachment may result in drastically differing basin geometries. The detachment and basic fracture pattern create a flower structure geometry throughout the basin (Allen & Allen, 1990; Rahe et al., 1998; Sims et al., 1999; Twiss & Moores, 1992). As it is seen in figure 1.6, the main strike slip fault has several branches that spur off in several sub-parallel directions. These branches further fracture the basin as well as creating a very complicated fault pattern and tectonic regime.

If the basin has a detachment in the brittle deformation zone, the basin tends to have one bounding normal fault and is comprised of a single extending basin that is fractured by a series of synthetic and antithetic faults (Rahe et al., 1998; Sims et al., 1999). However, if the pull-apart basin has a detachment that is located within a ductile zone, the system is comprised of individual sub-basins, or it is composed of smaller basins that eventually coalesce. These basins, although extensional in the main depocenters, have a strike-slip- dominated geometry. Depending on the thickness of the ductile zone and how far the 1.2 Pull-apart basins 9

Figure 1.6: Sandbox-models of an asymmetric pull-apart system; (a) after 5.00 centimeters of asymmetric extension; (b) Cross section detailing the structure of the faults; (c) fault schematic of the cross section (Rahe 1998). 1.3 Back-arc Basins 10 detachment penetrates into this zone, it will determine the shape and tectonic character of the pull-apart basin (Sims et al., 1999). Examples of pull-apart basins come from Death Valley in California United States (Burchfiel et al., 1987) and North Anatolian Fault in Turkey (Sengor et al., 1985).

1.3 Back-arc Basins

Back-arc basins are extensional basins that are usually located behind a volcanic arc that has formed due to the influences of a subducting oceanic plate under the associated active continental margin (Allen & Allen 1990). These basins can either be underlined by thinly stretched continental crust, or, if the extension is intense enough, oceanic crust can be generated (Allen & Allen, 1990; Sdrolian & Muller, 2006). After the initial rifting and extension, back-arc basins are subject to subsidence either from thermal relaxation or coeval with folding and thrusting, which occur along the arc boundary.

The most accepted tectonic reason for the extension of a back-arc basin is the influence of slab roll-back at the subducting hinge, which occurs when the vertical velocity of the down-going oceanic crust exceeds the rate of convergence. This basically means that the subducting slab is pulling the continental crust along the zone of interaction, thus creating an area of extension that is surrounded by a compressional regime (Hall et al., 2003; Sdrolian & Muller, 2006).

According to Sdrolian & Muller (2006), the formation of a back-arc extensional regime is most successful when the oceanic lithosphere that is being subducted under the more buoyant continental lithosphere is older than 55 my. They argue that the oceanic litho- sphere should have had enough time to cool and become dense enough to be subducted in a fashion that the interaction between the oceanic plate and the continental plate can create the required extension behind the arc. In addition, the angle of incidence that the subducting plate has with the continental margin is important. The old dense oceanic crust is most successful at creating a roll-back hinge when the angle of subduction is above 30 ◦, thus, giving the oceanic and continental plates the right orientation and interplaying dynamics to create the regime of extension behind the arc (Sdrolian & Muller, 2006).

Back-arc basins can be some of the most rapidly extending regions on the modern crust due to their mechanism of extension and the near proximity of the extending basin to the area of stress generation thus giving it a pseudo-focusing effect on the extension (Allen & Allen, 1990).

The Aegean Sea and the Black Sea are good examples of extension zones that have been or are currently being created and deformed by extensional forces that are a result of a subduction process. 1.4 Cratonic Sags 11

The Black Sea back-arc basin was formed in association with the subduction events of the PaleoTethys and NeoTethys oceanic crusts under the Eurasian craton during the Creta- ceous and Paleocene, respectively (figure 1.7; Nikishin, 2003). As the oceanic plates were subducted under the craton, an associated volcanic arc formed in present-day Turkey. The roll-back forces of the PaleoTethys and the later NeoTethys oceanic plates being subducted under the Eurasian cratonic continental crust created a regime that was able to create ex- tending basin in response to the subduction forces (Robinson et al., 1995; Banks et al., 1997; Gorur and Tuysuz, 1997; Yilmaz et al., 1997).

Back-arc basins have a stretching direction, which usually mimics the shape and strike of the oceanic subducting hinge (Allen & Allen, 1990; Sdrolian & Muller, 2006). These basins usually posses main bounding normal faults on each basin margin. There is also a number of associated normal faults in the interior portions of the basins that allow for accommodation space from the incurred extension throughout the time of active subduction (Allen & Allen, 1990).

All the faults are characterized by roughly strike-parallel orientation to the subducting hinge. Depending on pre-rift rheology and architecture of the crust, the basin may have different orientation of the normal faults and this can result in differing geometries of the basin (figure 1.7; Nikishin, 2003).

Due to the large amount of heat flow from the volcanic arc and the stretching itself, the thermal subsidence accounts for a large portion of the sedimentary accommodation space. An example comes from the Black Sea where the Maykop formation and Neogene strata, which can reach thicknesses of up to 4 and 2 kilometers respectively, were deposited in a back-arc setting which was thermally subsiding (Robinson et al., 1995; Banks et al., 1997). This is supported by the absence of any new normal faults through the system as well as the continuity of the strata.

The back-arc basins that are active in the present-day are mainly located around the perimeter of the Pacific Ocean, developed in response to the subduction of the Pacific Plate around the rim (figure 1.8). Other active basins are located around the Indian Ocean and Mediterranean Sea.

1.4 Cratonic Sags

Cratonic sag basins are sedimentary basins that are superimposed upon a previously failed rift system (figure 1.9; Middleton, 1989; Ulmeshik, 2003). They are usually located within the middle of stable continental blocks (Middleton, 1989). The striking feature of these basins that sets them apart from rifts is the nature of the sedimentary cover in the basin. The sedimentary cover is deposited in a basin mainly without the aid of faulting on the 1.4 Cratonic Sags 12

Figure 1.7: Basic tectonic development of the Black Sea since the Neocomian to Eocene. Note the horst and graben styling of the Eastern Black Sea Basin (Nikishin 2003) 1.4 Cratonic Sags 13

Figure 1.8: Active back-arc basins throughout the world (Stern, 2006) 1.4 Cratonic Sags 14 margins or within the basin interior itself (Middleton, 1989). This excludes any sort of tec- tonic extension or separation as the primary reason for subsidence and leaves the remaining equalizing forces from the failed rift system as the main instigators of the basin subsidence (Middleton, 1989). These forces include both thermal relaxation due to the cooling of the post-rift system, as well as the isostatic crustal rebound in response to sediment loading (Middleton, 1989).

Figure 1.9: Asthenospheric cooling effect on the lithosphere (Middleton, 1989).

When the rifting first initiates within the continental craton, the lithospheric stretching causes an upwelling of the hot asthenosphere. This causes a higher than normal heat flow in this area which is controlled by convection currents of the hot asthenosphere. This activity is common for many regions that undergo extension of the lithosphere (Middleton, 1989; Ulmeshik, 2003). As the extension in the rift shuts down, the previously ascending hot asthenosphere starts to cool and, as it subsides, it begins to pull the previously rifted crust down with it (Middleton, 1989). As the sediments begin to fill the depression remaining after the extension, the system continues to cool and subside further. This creates more accommodation space for the eroding sediments from the margins that are now relatively 1.4 Cratonic Sags 15 growing because of the subsidence of the topographically low depression. These cratonic sag basins, because they are formed in the middle of stable cratons, end up being the main depocenters for most of the sediments from the surrounding regions (Middleton, 1989). These sediments put more weight on the depo-center which causes more subsidence in the basin.

An example of a cratonic sag basin is the Western Siberian Basin in the north central portion of Russia between the Ural Mountains and the Yenisey River (Ulmeshik, 2003). This basin is a Mesozoic-Tertiary-aged sag, which is superimposed upon an earlier Triassic rift system (Ulmeshik, 2003). In the case of the Western Siberian Basin, the maximum sediment thickness is approaching 4 km with an average of approximately 3 km of sedi- ments, which indicates that there was a considerable amount of post-rift subsidence in this region (Ulmeshik, 2003). The sediments above the extended rift basement have virtually no brittle deformation due to faulting, and in certain areas, especially within the center of the depocenters, there are few or no faults present to account for the creation of the accommodation space (Ulmeshik, 2003). Chapter 2

Rift system linkages and geometries

Rift systems evolve as a series of growing normal faults which eventually interact with each other. The continuing linkage of these normal growing faults eventually controls subsiding sub-basins. It is the interaction and joining of these sub-basins that creates a rift system. This chapter will focus primarily on:

• nucleation and interlinking of growing normal faults to create a controlling basin margin fault; • types and evolution of the linkages between two growing normal faults; • interaction of sub-basins and the role of the extensional geometries in the process;

The goal of this chapter is to describe the evolution of a rift system from the nucleation of normal faults, through the steps of the linkages between the normal faults which lead to the creation of a sub-basin and describing the geometries and linkages of sub-basins that eventually create a rift system.

2.1 Fault geometries and linkages

Although continental rifts can be formed in a variety of tectonic settings and in areas with vastly differing rheologies, all rift systems contain certain basic intrinsic properties such as:

1. basin margin faults that are composed of several smaller normal faults that have sub-sequentially linked together and very frequently cut through crystalline basement (Gibbs, 1984; Ebinger et al., 1999; Withjack et al., 2002; Noll and Hall, 2006); 16 2.1 Fault geometries and linkages 17

2. the growth and eventual interaction of normal faults with neighboring faults (Trudgill and Cartwright, 1994; Cowie et al., 2000; Mansfield and Cartwright, 2001; Young et al., 2001; Withjack et al., 2002; McClay et al., 2002); and

3. the interaction between previously independently forming sub-basins with other sub- basins which results in a rift system (Rosendahl, 1987; McClay, 2002).

Rift basins are elongate in shape and can be several kilometers deep, tens of kilometers wide, and hundreds of kilometers long (Scholz, 1998; Withjack et al., 2002).

As a certain region begins to undergo extensional stress loading, normal faults begin to form in areas of previous anisotropies or areas of pre-existing weaknesses (Mansfield and Cartwright, 2001; Noll and Hall, 2006). Such areas serve as a nucleation point for normal faults, which may link together with neighboring faults to form large-scale normal faults that will grow over the period of extension (figure 2.1; Mansfield and Cartwright, 2001, Young et al., 2001).

Figure 2.1: Stretching models that illustrate the nucleation of a fault through the eventual linkage with neighboring normal faults which results in the development of a boundary fault (Mansfield and Cartwright, 2001).

However, many times the small fault segments will not manage to link with any other neighboring faults and will end up being abandoned due to the larger fault segments ac- commodating most of the stress and displacement within in the extensional area (figure 2.2 Fault linkages 18

2.2 ; Mansfield and Cartwright, 2001; Noll and Hall, 2006). As the individual fault seg- ments begin to create longer and more continuous faults, these faults will have a larger displacement than the smaller faults surrounding it (Mansfield et al, 2001; McClay et al., 2002). This linking of fault segments into larger segments eventually leads to the creation of a basin controlling boundary fault on the margin of the rift (Mansfield and Cartwright, 2001; McClay et al, 2002; Withjack et al., 2002; Noll and Hall, 2006; Young et al., 2001).

The boundary faults almost cut into the crystalline basement (Gibbs, 1984; Ebinger et al., 1999; Withjack et al., 2002). The controlling boundary fault, however, does not ac- commodate all of the extensional forces in the area, so as the basin continues to extend, more normal faults in the basin center begin to link together and create accessory faults throughout the basin (Gibbs, 1984; Rosendahl, 1987; Trudgill and Cartwright 1994; Mans- field and Cartwright 2001; McClay et al, 2002; Withjack et al, 2002). The footwall regions of the border faults are usually uplifted which produces a small flank on the rift basin. Conversely the hanging wall usually contains a low-lying trough like feature. The trough features on the border fault serve as very effective conduits for sediments that are being shed by the new topographic highs from recent fault movements (Withjack et al., 2002).

Figure 2.2: Evolution and linking of a boundary normal fault (Mansfield and Cartwright, 2001).

2.2 Fault linkages

Over a period of time, many of the smaller growth fault segments will create linkages with neighbors, contributing to the development of the large scale faults (figure 2.1; Trudgill and Cartwright, 1994; Cowie et al., 2000; Mansfield and Cartwright, 2001; Young et al., 2001; McClay et al., 2002; Withjack et al., 2002; Noll and Hall, 2006). As more of these isolated small normal faults begin to link together, the joining of these smaller faults will be more and more likely to rupture, which alleviates some of the extensional load stresses in the basin. This process of accommodating the stresses in a basin will cause the isolation 2.2 Fault linkages 19 of some of the other, lesser developed, nucleated faults surrounding that linked fault. This process of isolation causes many of these faults to become abandoned (Noll and Hall, 2006).

The linking of two proximal normal faults that share the same dip direction starts with a damage zone that is located between the tips of two normal faults within proximity of each other (Rosendahl 1987; Peacock, 2002; Withjack et al, 2002). This fracture zone is an area that contains numerous small-scale deformational features. As the normal faults continue to evolve with ongoing extension, they will eventually form a relay ramp structure in a stepover between their propagating tips (figure 2.3; 2.4; Trudgill and Cartwright 1994; Walsh et al., 1999; Peacock, 2002).

Figure 2.3: Basic structure and evolution of a relay ramp structure from conception to breach (Peacock, 2002).

The relay ramp functions as a soft linkage between the two faults, which means it is a zone of accommodation for the differing rates of displacement between the two fault tips rather than a direct linkage between the two. The term ’soft linkage’ merely refers to the fact that there is no through-going connection that directly joins the two main normal faults together, rather there exists an area of strain compensation between the two (Rosendahl, 1987; Trudgill and Cartwright, 1994; Cowie et al., 2000; Peacock, 2002).

The relay ramp itself contains a series of strike-parallel normal faults and joints that also act as an accommodation within the relay ramp itself and divide it longitudinally. There are also strike-perpendicular joints that are created due to the difference in movement of the two main faults (figure 2.4; Trudgill and Cartwright, 1994). As the two main faults 2.2 Fault linkages 20

Figure 2.4: Strike oriented view of a relay ramp structure. Note the parallel and perpen- dicular fractures within the ramps themselves that are serving as joints to accommodate the differing movements between the two faults (Trudgill and Cartwright, 1994). continue to grow, the relay ramp will eventually become breached on each of its margins (figure 2.4; Peacock, 2002). As the main faults continue to respond to extension, the relay structure will continue to serve as a soft linkage that serves as an area of fracturing and connection. However, as soon as the two main fault tips propagate to the point that the relay ramp becomes breached, the two faults will join at the tips and become essentially a through-going fault (figure 2.5; Rosendahl, 1987; Trudgill and Cartwright, 1994; Withjack et al., 2002). The remnant relay ramp structure is then abandoned and erosion eventually destroys the structure totally (figure 2.5; Trudgill and Cartwright, 1994).

The dimensions of a relay ramp depend on factors such as the rheology of an area as well as the angle of extension to the rift zone trend (McClay et al., 2002). In the case of the Murchison - Statfjord North Fault Zone in the northern North Sea, the relay ramp between two propagating fault tips was approximately 2 kilometers wide and persisted for about 10.5 million years (Young et al., 2001). During this time the relay ramp served as a sedimentary conduit between the hanging walls and footwalls (Young et al., 2001) as is the case for many of the relay ramps in Canyonlands National Park in southern Utah (Trudgill and Cartwright, 1994).

Different accommodation zone geometries develop if the two neighbor faults have opposing dip directions. It can be either a high or a low relief accommodatoin zone depending on the amount of overlapping of the main boundary faults (figure 2.6 (a) (b); Rosendahl, 1987; McClay et al, 2002; Withjack et al, 2002). The high relief accommodation zone develops in small overlap case forming an intra-basinal high. (figure 2.6 (b); Rosendahl, 1987). The low relief accommodation zone develops in large overlap case (Rosendahl, 1987) It forms 2.2 Fault linkages 21

Figure 2.5: Evolution of a pair of faults from the nucleation (a), creation of a relay ramp (b), breaching of the relay ramp (c), hard-linking and erosion of the relay ramp structure (d);(Trudgill and Cartwright, 1994). 2.3 Sub-basin interactions and the role of extension geometry in the process 22 an intra-basin swell neighbored by basin lows on each side (figure 2.6 (a); Rosendahl, 1987; McClay et al, 2002).

As the linked boundary faults continue to interact and join together, the process eventually creates a sub-basin within the developing rift zone. As the extension continues throughout a region, such sub-basins will eventually grow to the point that they will start interacting with other sub-basins. Depending on the obliqueness and angle of extension with respect to the overall rift zone, the way these sub-basins react with each other differs (Trudgill and Cartwright, 1994; McClay et al, 2002; Withjack et al, 2002).

2.3 Sub-basin interactions and the role of extension geometry in the process

The angle of extension to the rift zone axis plays a critical role in the normal fault for- mation and the geometry of fracturing in the basin, which has been demonstrated by the experimental modeling (McClay et al.2002). The extension oriented at 90◦ to the rift zone trend creates an area that contains long, linear, strike-parallel faults that link together through a series of small, narrow high angle relay ramps (figure 2.7 (a); Rosendahl, 1987; McClay et al., 2002; Withjack et al, 2002). The normal faults that occur in this orthogonal extension setting usually have a very steep dip. This setting is characterized by the usual lack of strike-slip and/or oblique slip accommodation faults, which are unnecessary due to the absence of any lateral movement and very little offset from the main boundary fault nucleations (Mansfield and Cartwright, 2001; McClay et al., 2002; Withjack et al., 2002).

As the angle of extension to rift zone trend is altered to 60◦ (figure 2.7 b), the border fault pattern resembles the orthogonal rift case, however; wider and more frequent relay-ramps are needed due to the normal faults forming at a greater offset from each other (McClay et al, 2002). As extension continues, the intra-basinal faults continue to interact with each other by means of relay-ramps. The development of strike-slip and oblique slip faults is not needed (McClay et al., 2002; Withjack et al., 2002).

As the extension angle to the rift zone trend is increased to 45◦, the relay ramps become wider in order to accommodate for the growing spacing between tthe neighboring boundary faults. However, in this case the need for strike-slip faults is still un-necessary (figure 2.8 (a) (b); McClay et al., 2002). It can be noticed that as the width of the relay ramps increase, the angle of extension in relation to the trend of the rift zone also increases.

Interactions between the separate sub-basins are also affected by the angle of extension (McClay et al., 2002; Withjack et al., 2002). The transfer zone between the two sub-basins is usually composed of one or more strike-slip or oblique slip faults (figure 2.9; Rosendahl, 1987; Mansfield and Cartwright, 2001; McClay et al., 2002 Withjack et al., 2002). As these 2.3 Sub-basin interactions and the role of extension geometry in the process 23 laterally offset sub-basins continue to evolve, the strike-slip or oblique-slip faults of the accommodation zone will eventually develop one or more strike-slip or oblique-slip faults that span the entire area between the two sub-basins. This creates a hard linkage between the two sub-basins and an effective means for the sub-basins to interact with each other as extension continues. 2.3 Sub-basin interactions and the role of extension geometry in the process 24

(a)

(b)

Figure 2.6: (a) Low relief accommodation zone created from a pair of overlapping, opposing dip faults. (b) High relief accommodation zone created from a pair of opposing dipping faults (Rosendahl, 1987). 2.3 Sub-basin interactions and the role of extension geometry in the process 25

(a)

(b)

Figure 2.7: Sandbox-models for a basin extending at 90◦ (a) and 60◦ (b) to the rift zone trend (McClay et al., 2002). 2.3 Sub-basin interactions and the role of extension geometry in the process 26

(a)

(b)

Figure 2.8: Sandbox-model for a basin extending at 45◦ to the rift zone trend. (a-c) map view after increments of extension, (d) cross sections as located in (c); (McClay et al., 2002). 2.3 Sub-basin interactions and the role of extension geometry in the process 27

(a)

(b)

Figure 2.9: (a) Sand-box model of a transfer zone between two offset rifts (b) note the need for oblique-slip and strike-slip faults for accommodation between the two basins (McClay et al., 2002). Chapter 3

Pull-apart basin development and geometries

Pull-apart basins are localized structural lows that occur in basement strike-slip systems. These basins are usually rhombic in shape and are bound by obliquely extending normal faults. Pull-apart basins can have one or numerous depositional centers (figure 3.1; Allen and Allen, 1990; McClay and Dooley, 1995; Rahe et al., 1998; Nilsen and Sylvester, 1999; Sims et al., 1999; Wu et al., 2009). The geometry, number of contained depositional centers and characteristics of a pull-apart basin depend on:

• angle of extension (pure strike-slip or various degree of a normal slip component), (McClay and Dooley, 1995; Rahe et al., 1998; Wu et al., 2009); • occurrence of a brittle or ductile d´ecollement (Sims et al., 1999); • thickness of the d´ecollement(Sims et al., 1999); • rate of overall extension (McClay and Dooley, 1995; Rahe et al., 1998; Nilsen and Sylvester, 1999; Atmaoui et al., 2006; Wu et al., 2009); and • relative fault movement rate along the master strike-slip/oblique slip faults (McClay and Dooley, 1995; Rahe et al., 1998; Atmaoui et al., 2006; Wu et al., 2009).

3.1 Basic basin evolution and resulting geometries

The inception of a pull-apart basin begins in the stress zone that is located between two under lapping strike-slip faults or in a bend on a single strike-slip fault (figure 3.2 (a) ;(b)). 28 3.1 Basic basin evolution and resulting geometries 29

Figure 3.1: Block model of a pull-apart basin. Note the asymmetric flower structure as well as the ramps that are utilized for sediment transport into the basin (Wu et al., 2009). 3.1 Basic basin evolution and resulting geometries 30

The under lapping strike-slip faults can either be perfectly parallel to each other, or they can be at an angle to each other (Allen and Allen, 1990; McClay and Dooley, 1995; Rahe et al., 1998; Nilsen and Sylvester, 1999 a,b; Sims et al., 1999; Wu et al., 2009).

Figure 3.2: (a) pull-apart basin forming in a bend of a strike-slip fault; (b) pull-apart basin forming in the zone between two strike-slip faults; (c) basic shears that form in an area of strike-slip deformation (Rahe et al., 1998).

The accommodation zone is this aforementioned stress zone between the two bounding strike-slip faults. The fracturing pattern in the accommodation zone starts with a series of R-shears (figure 3.2c). These R-shears are at low angles to the strike of the main strike slip faults and cut through the accommodation zone in a pattern that links the tips of the two main bounding strike-slip faults (figure 3.3 Rahe et al., 1998; Sims et al., 1999; Wu et al., 2009).

The ratio of strike-slip movement versus dip-slip movement of these cross cutting R-shears as well as their number depends on the angle, distance between and movement rate of the main bounding strike-slip faults. If the main boundary strike-slip faults are parallel to each other, R-shears do not usually create a hard-linking zone across the basin (figure 3.4; Wu et al., 2009). However, given enough extension in the basin, the R-shears will eventually be linked by a series of other fractures that can create a through-going hard-link. As the angle of divergence between the two main bounding strike-slip faults increases, the likelyhood that the R-shears, which will later be linked together by a series of P-shears, will create a ’hard-linking’ fracture zone through the basin increases as well as the creation of an internal 3.1 Basic basin evolution and resulting geometries 31

Figure 3.3: En-echelon R-shear pattern developed in the accommodation zone between two strike-slip faults (Atmaoui et al., 2006). 3.1 Basic basin evolution and resulting geometries 32 horst and graben structure that can create multiple depositional centers and through-going basin faults (figures 3.5, 3.6G ¨olke et al., 1994; Nilsen and Sylvester, 1999; Waldron, 2004; Atmaoui et al., 2006; Wu et al., 2009).

As the angle between the two bounding normal faults increases, the amount of dip-slip on the R-shears will increase, whereas if the main bounding strike-slip faults are parallel to each other, the amount of dip-slip is reduced (Wu et al., 2009).

The geometry of the R-shears can also be controlled by the depth, thickness and deforma- tion environment of the d´ecollement zone. With an increase in d´ecollement thickness, the R-shears tend to become more sub-parallel, to the main basin bounding strike-slip faults, more widely spaced, and more numerous (Sims et al., 1999). As the R-shears continue to grow, the amount of dip slip, if any, on the main bounding faults is mainly transferred to the R-shears, thus making the R-shears in the accommodation zone the main source of subsidence at this particular time in the basin evolution (Sims et al., 1999; Wu et al., 2009). The area between the R-shears serves as a ramp conduit for sediments that are shedding from the new topographic relief being created by the subsidence of the R-shears and flanking faults. During this basin evolution stage, the largest amount of sediments is deposited. However, it is not always true.

It can be seen in places like the Sea of Marmara pull-apart system in Turkey, Vienna Basin in Austria and the Stellarton pull-apart basin in Nova Scotia that the rates of subsidence are not very continuous. On the contrary, the basins tend to subside in events characterized by fast subsidence divided by periods of either slow subsidence, or tectonic quiescence (figure 3.7 Armijo et al., 2002; King et al., 2002; Ehrhardt et al., 2004; Kovac, 2004; Waldron, 2004). These basins can experience long period of quiescence until enough stress has built up on the fault planes and in the basin center that another event becomes inevitable (Waldron, 2004).

As the accommodation zone continues to extend and the R-shears have traversed much of the basin, P-shears begin to propagate throughout the basin as well as linking R-shears (McClay and Dooley, 1995; Rahe et al., 1998; Wu et al., 2009). In basins that have a thin brittle d´ecollement zone, the R-shears and P-shears usually create a single narrow basin (Wu et al., 2009). The R-shears and linking P-shears create a hard linkage through the middle of the basin which becomes the main depositional center. This narrow depositional center is also achieved by a system in which the two main boundary strike-slip faults are almost parallel to each other.

It is at this point of the basin evolution that the pull-apart basin begins to develop the characteristic flower structure. This is a structure in which many of the curved faults have converged in the basement. In the case of pull-apart basins, this flower structure is negative. On the contrary, the flower structure is positive in areas of convergence (Allen and Allen, 1990; McClay and Dooley, 1995; Wu et al., 2009). 3.1 Basic basin evolution and resulting geometries 33

Figure 3.4: Evolution model of a pull-apart basin with parallel marginal strike-slip faults (Wu et al., 2009). 3.1 Basic basin evolution and resulting geometries 34

Figure 3.5: Evolution model of a pull-apart basin with bounding strike-slip faults at an angle to each other. PDZ - principal displacement zone (Wu et al., 2009). 3.1 Basic basin evolution and resulting geometries 35

Figure 3.6: Multiple basins of the Dead Sea pull-apart system caused by strike-slip faults with differing strikes. PDZ - principal displacement zone (Wu et al., 2009). 3.1 Basic basin evolution and resulting geometries 36

Figure 3.7: Correlated stratigraphic cores from the Stellarton basin which depict how the evolution of a pull-apart basin does not contain consistent opening, but rather a series of events and periods of quiescence (Waldron, 2004).

As a pull-apart on a thin brittle d´ecollement continues extending and fracturing, there are usually two flanking oblique-slip faults that have a large amount of dip-slip that extend from the tips of the main boundary strike-slip faults (figure 3.4 Allen and Allen, 1990; Wu et al., 2009). These flanking faults essentially encapsulate the basin and contribute to the overall terraced topography of the basin. These flanking faults are important for the sediment transport into the basin as they will create a zone of low topographic relief and tilt in the basin. The margins of the flanking faults (figure 3.1) create a funneling effect of sediments into the basin. This style of extension was modeled in Wu et al., 2009 and occurs in the Strait of Sicily rift zone, Atmaoui et al., (2006).

If the geometry of the main boundary strike-slip faults is not parallel, or there is a ductile d´ecollement, the resulting pull-apart basin can have two or more depo centers (McClay and Dooley, 1995; Sims et al., 1999; Wu et al., 2009). The models of Wu et al., (2009) indicate the angle between the strike-slip faults controls the fracturing of the basin. The R-shears, which created a cross basin hard linkage in conjunction with the P-shears in the narrow basin case, have now become more dubious in their placement and do not make a single straight hard linkage through the basin. Instead, depending on the angle between the bounding strike-slip faults and the nature of the d´ecollement zone, the R-shears are scattered through the basin.They have differing amounts of dip-slip. The basin contains a horst and graben structure in its center. This is where the basin begins to separate into sub-basins (Sims et al., 1999; Wu et al., 2009). 3.1 Basic basin evolution and resulting geometries 37

With the structural high in the basin center, the discrete sub-basins will begin to evolve almost separately from each other. It can be seen in the Dead Sea Transform as well as in the Vienna Basin that there are many separate sub-basins within the main pull-apart system (Armijo et al., 2002; Ehrhardt et al., 2004; Kovac et al., 2004; Wu et al., 2009). Each sub-basin has a sedimentary record and evolution that slighly differs from that of the neighboring sub-basins. This is caused by the different times of opening, sediment sources and the subsidence rate. In the Vienna basin, there are two separate sub-basins that have differing thicknesses of Neogene sediments and each of them has a slightly different tilt than the other (Kovac et al., 2004). Bibliography

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