Border faults that bound asymmetric rift basins with uplifted flanks are among the most common features in continental rifts (Fig. 7.25). Some aspects of this characteristic morphology can be explained by the elastic response of the lithosphere to regional loads caused by normal faulting.
Plate flexure (Section 2.11.4) describes how the lithosphere responds to long-term (>105 years) geologic loads. By comparing the flexure in the vicinity of
4 Rift flank uplift
4 Rift flank uplift
different types of load it has been possible to estimate the effective long-term elastic thickness (Te) of continental lithosphere (Section 2.12) using forward models of topography and gravity anomaly profiles (Weissel & Karner, 1989; Petit & Ebinger, 2000). The value of Te in many rifts, such as the Basin and Range, is low (4 km) due to the weakening effects of high geothermal gradients. However, in other rifts, including those in East Africa and in the Baikal Rift, the value of Te exceeds 30 km in lithosphere that is relatively strong (Ebinger et al., 1999). The physical meaning of Te, and its relationship to the thickness (Ts) of the seismogenic layer, is the subject of much discussion. Rheological considerations based on data from experimental rock mechanics suggest that Te reflects the integrated brittle, elastic, and ductile strength of the lithosphere. It, therefore, is expected to differ from the seismogenic layer thickness, which is indicative of the depth to which short term (periods of years) anelastic deformation occurs as unstable frictional sliding (Watts & Burov, 2003). For these reasons, Te typically is larger than Ts in stable continental cratons and in many continental rifts.
The deflection of the crust by slip on normal faults generates several types of vertical loads. A mechanical unloading of the footwall occurs as crustal material in the overlying hanging wall is displaced downward and the crust is thinned. This process creates a buoyancy force that promotes surface uplift. Loading of the hanging wall may occur as sediment and volcanic material are deposited into the rift basin. These loads combine with those that are generated during lithospheric stretching (Section 7.6.2). Loads promoting surface uplift are generated by increases in the geothermal gradient beneath a rift, which leads to density contrasts. Loads promoting subsidence may be generated by the replacement of thinned crust by dense upper mantle and by conductive cooling of the lithosphere if thermal diffusion outpaces heating.
Weissel & Karner (1989) showed that flexural iso-static compensation (Section 2.11.4) following the mechanical unloading of the lithosphere by normal faulting and crustal thinning leads to uplift of the rift flanks. The width and height of the uplift depend upon the strength of the elastic lithosphere and, to a lesser extent, on the stretching factor (P) and the density of the basin infill. Other factors may moderate the degree and pattern of the uplift, including the effects of erosion, variations in depth of lithospheric necking (van der Beek & Cloetingh, 1992; van der Beek, 1997) and, possibly, small-scale convection in the underlying mantle
(Steckler, 1985). Ebinger et al. (1999) showed that increases in the both Te and Ts in several rift basins in East Africa and elsewhere systematically correspond to increases in the length of border faults and rift basin width. As the border faults grow in size, small faults form to accommodate the monoclinal bending of the plate into the depression created by slip on the border fault (Fig. 7.25). The radius of curvature of this bend is a measure of flexural rigidity. Strong plates result in a narrow deformation zone with long, wide basins and long border faults that penetrate deeper into the crust. Weak plates result in a very broad zone of deformation with many short, narrow basins and border faults that do not penetrate very deeply. These studies suggest that the rheology and flexural rigidity of the upper part of the lithosphere control several primary features of rift structure and morphology, especially during the first few million years of rifting. They also suggest that the crust and upper mantle may retain considerable strength in extension (Petit & Ebinger, 2000).
Lithospheric flexure also plays an important role during the formation of large-magnitude normal faults (Section 7.3). Large displacements on both high- and low-angle fault surfaces cause isostatic uplift of the foot-wall as extension proceeds, resulting in dome-shaped fault surfaces (Buck et al., 1988; Axen & Bartley, 1997; Lavier et al., 1999; Lavier & Manatschal, 2006). Lavier & Manatschal (2006) showed that listric fault surfaces whose dip angle decreases with depth (i.e. concave upward faults) are unable to accommodate displacements large enough (>10 km) to unroof the deep crust. By contrast, low-angle normal faults whose dips increase with depth (i.e. concave downward faults) may unroof the deep crust efficiently and over short periods of time if faulting is accompanied by a thinning of the middle crust and by the formation of serpentinite in the lower crust and upper mantle. The thinning and serpentini-zation weaken the crust and minimize the force required to bend the lithosphere upward during faulting, allowing large magnitudes of slip.
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