Continental Rifted Margins 1. Gwenn Peron-Pinvidic
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extending – this one is called the “aulacogen” (Burke 1977; Şengör and Burke 1978) (Figure 1.6). Aulacogens are often reactivated in the post-rift stages under renewed extension or compression. A well-studied example is the Benue Trough, that extends from the equatorial Atlantic Ocean into the continental hinterland of Nigeria (Figure 1.7). It formed in the Early Cretaceous as a major rift arm accommodating the separation of South America from Africa (Olade 1975; Robert 2008).
Figure 1.5. Topographic map of the East African Rift system with indication of some of the major segments, basins and structures regularly mentioned in the literature. Elevation and bathymetric map is based on the NOAA Etopo1 (source: Amante and Eakins 2009)
Figure 1.6. Schematic representation of an Aulacogen: a typically narrow intracontinental rift in a triple junction that ceased activity
Figure 1.7. Maps locating the Benue Trough aulacogen in the South Atlantic plate tectonic evolution. The maps have been created with the GPlates software (Müller et al. 2018), using the Matthews et al. (2011) sample data sets. Elevation and bathymetric map is based on the NOAA Etopo1 (source: Amante and Eakins 2009)
Wide intracontinental rifts
Wide/diffuse rifts are characterized by uniform thinning of the crust and lithospheric mantle over a width greater than the thickness of the lithosphere (Figure 1.8). The resulting geometry typically corresponds to isolated basins that are distributed over broad regions (Brun Choukroune 1983), as in the Basin and Range Province (e.g. Dickinson 2002; Parsons 2006). Unlike narrow rifts, wide/diffuse rifts are characterized by relatively small lateral gradients in crustal thickness and topography variations, with a relatively flat Moho.
Case example: The Basin and Range Province
The Basin and Range Province is a region covering a vast area in the western United States and northwestern Mexico (Figure 1.9). The physiography is defined by changes in elevation, alternating between north–south oriented, narrow-faulted mountain chains (15–20 km wide) and flat arid valleys (~30 km wide). The subparallel mountain ranges correspond to the crests of fault-blocks bounded by normal faults developed from east–west-directed extension. The blocks are exposed to erosion, leading to sediment transport to the adjacent valleys. However, the climate is so arid in most of the province that the runoff water is not significant enough to proceed to long-distance transport of the sediments. This special setting results in the building of thick alluvial fans, which led to the definition of the famous “Gilbert deltas” (Gilbert 1928).
Figure 1.8. Illustration of the “wide rift mode” (source: Buck 1991)
The tectonic reason for the formation of the province is interpreted to be due to crustal buoyancy forces and stress changes at plate boundaries (Sonder and Jones 1999). The entire region is characterized by high heat flow, slightly thinned continental crust (27–35 km) and a relatively flat Moho with reduced upper mantle P wave velocities (7.5–8.0 km s-1) (Holbrook 1990; Gilbert 2012). Structurally, the region displays distinct modes of deformation (high-angle faulting, low-angle detachment and metamorphic core complexes) and the rocks have experienced syn-deformational high-grade metamorphism (amphibolite, granulite, eclogite facies) (Axen et al. 1993). The observations led to the development of the simple shear model (Wernicke 1981, 1985) (see Chapter 2) where a detachment surface separates an upper plate setting (hanging wall) from a lower plate setting (footwall), which are mostly made of upper crust dissected into blocks by high-angle normal faults and highly-deformed lower crust, respectively (Davis and Lister 1988; Lister and Davis 1989).
Figure 1.9. Topographic map of the Basin and Range Province. Elevation and bathymetric map is based on the NOAA Etopo1 (Amante and Eakins 2009). Cross-sections modified from Brun et al. (2018), based on Miller et al. (1999) and Spencer and Reynolds (1990) to illustrate the combined effects of detachment, decollement and normal faults into the formation of the Basin and Range physiography
Intracratonic rifts
Cratons are regions of the Earth where the continental lithosphere is much thicker than the average thickness of 35–40 km and are considered to be tectonically stable. The lithospheric mantle roots of cratons have unusually high thicknesses, reaching 200–400 km – more than double the standard values observed at the non-cratonic continental lithosphere (Artemieva Mooney 2002). Cratons are usually divided into two categories: the shields, where crystalline basement rocks outcrop and platforms, where the basement is overlaid by younger sediments. An intracratonic rift is thus a rift that develops on a specific lithospheric context away from active tectonic boundaries (Figure 1.10).
The driving forces creating intracratonic rifts remain extremely poorly understood. Their characteristics are not explained by standard rift models, as the extensive basement faulting and graben and/or half-graben basins appear to be absent (Klein and Hsui 1987; Hartley and Allen 1994). Uniform stretching models could explain the basic configuration of intracratonic basins, but these models fail at explaining the subsidence characteristics. The subsidence is extremely slow, and subsidence curves typically lack the characteristic two-step structure usually observed for rifted margins with the distinction between the first phase of tectonic-related subsidence (the effect of the faulting events) and the thermal subsidence (see Chapter 2). Various hypotheses have been posited to explain these peculiar subsidence curves of intracratonic basins. The main theories include subsidence caused by density changes within the lower crust and lithospheric mantle due to mineral phase changes (Gac et al. 2012), thermal relaxation and subsidence due to (negative) dynamic topography (Heine et al. 2008), thermal relaxation of a thick lithosphere undergoing low strain rates (Armitage and Allen 2010), tectonic reactivation (Braun and Shaw 2001) and uplift followed by subaerial erosion (Burov and Cloetingh 1997). Cacace and Scheck-Wenderoth (2016) reviewed all these hypotheses and determined that the long life span and slow subsidence of intracratonic basins are the result of feedback effects between sedimentation and thermal re-equilibrium at deeper crustal and mantle depths.
The basins of intracratonic rifts