phases or crustal uplift parameters. However, modern observations show that these models, although useful, are too simplistic. Various rifts show tectonic evolution and structural settings that combine “passive” and “active” characteristics (Ruppel 1995). These passive or active rifting classifications rather appear as end-member types in a continuous spectrum of continental extension processes.
Figure 1.3. Schematic illustration of the two main theoretical categories of rifting: active and passive. Active rifting is controlled by the warm asthenospheric mantle that rises beneath the base of the continental lithosphere, causing thinning of the plate and uplift at the rift zone. Passive rifting is controlled by distant extensional stresses, causing stretching and lithospheric thinning and passive rise of the asthenosphere below the rift zone
1.2.2. Plate tectonic setting classification
Rifts develop under extensional stress regimes. Even though the structural context sounds straightforward, the regional tectonic setting can span a wide range of situations between the end-member convergent (with fore-arc, back-arc, retro-arc and intra-arc basins) and divergent settings (with rift and rifted margin basins). Hybrid structural contexts are also encountered. Within each setting, several types of basins can be categorized, depending on the tectonic stress regime of the basin itself (divergent, strike-slip, transform, transtensional, transpressional), the composition of the basement (continental, cratonic, oceanic), the sediment supply (sediment-rich, sediment-starved), the magma supply (magma-rich, magma-poor) and the presence of salt. Numerous other parameters are also recognized as strongly influencing the development and final geometries of the rift, such as the thermal structure, lithospheric rheology, inherited structures, crustal composition, opening geometries and extension rate. The interplay between these parameters is still poorly understood.
We review the main rift classifications and their characteristics. In order to construct a clear list of the various types of rifts, we follow Ruppel’s (1995) early approach and propose that we first classify rifts based on their regional tectonic setting: intracontinental, convergent and divergent. This results in an overarching basic classification that has the advantage of embracing the majority of rift basin types, although some hybrid and alternative cases are always possible. The influencing parameters will be summarized in the next chapter.
1.2.2.1. Intracontinental rifts
Intracontinental rifts develop in the interior of a continental plate, far from any plate boundary, thus with no apparent direct tectonic connection to the regional stress fields produced by plate tectonics. However, continental plates can deform internally in response to far-off changes in stress conditions, such as changes in the plate boundary configuration (e.g. changes in the dip/velocity/direction of the subduction slab). Intracontinental rifts are classified according to their geometry (narrow vs. wide) or their structural setting (aulacogens and intracratonic rifts).
Narrow intracontinental rifts
Narrow rifts (Buck 1991) are characterized by a thinning in the crust and lithospheric mantle over a relatively narrow zone of about 100–150 km wide (Figure 1.4). The formation of narrow rifts is attributed to local weakening factors such as thermal thinning, strain weakening and local magmatism, where deformation is focused at the weakest or thinnest region of the lithosphere (Kusznir 1987; Buck 1991, 2004). The resulting rifts developed under this extensional mode are characterized by distinct lateral gradients in crustal thickness, topography and heat flow (Bonatti 1985; Corti et al. 2003). Archetypal examples of narrow rifts are the East African Rift system (Chorowicz 2005), the Rhine Graben (Brun et al. 1992; Dèzes et al. 2004), the Baikal Rift (ten Brink Taylor 2002), the Rio Grande Rift (van Wijk et al. 2018) and the Gulf of Suez-Red Sea Rift (Ragab and El-Kaliouby 1992; Bosworth et al. 2021).
Figure 1.4. Illustration of the “narrow rift mode” introduced by Buck (1991). In this extensional case, the lithospheric thinning affects a narrow zone of up to 100–150 km wide
Case example: The East African Rift
The East African Rift (EAR) covers a broad region of East Africa and encompasses a series of discrete fault-bounded valleys organized into distinct rifts, branches and plateaus (Figure 1.5). The EAR is considered to be the archetype of active intracontinental rifts. The EAR contains several distinct sectors, from the Red Sea, Gulf of Aden and Afar region in the North to the eastern and western branches southward on the African continent. The eastern branch crosses the Ethiopian Plateau and includes two main segments: the Main Ethiopian Rift and the Kenyan rift. The western branch contours along the East African Plateau and Tanzania Craton and includes the East African Great Lakes and particularly the Lake Albert and Albertine Rifts. The western branch continues south-eastward with the Malawi Rift towards the Mozambique Channel.
The regional architecture can be described as a series of grabens of various dimensions, arranged in an en-échelon geometry within the different branches. Typically, the basins are flanked by chains of normal faults which shift polarity along-strike through transfer zones (Chorowicz 2005; Ebinger 2005; Corti 2009). The basins are usually asymmetric, up to 7 km deep, and bordered by 70–120 km long normal faults that accommodate dip-slip movements. The Kenyan and Ethiopian branches are magmatically dominated and the successive stages of rift evolution are preserved (Keranen et al. 2004; Rooney et al. 2005). The extensional geometries are so evident at the tectonic plate-scale that researchers have already named the new “plates-to-be”: the Nubian plate, which corresponds to the western lands and encompassing most of Africa, and the Somalian plate, which is pulling away eastward (Bird 2003). The plates are separating at a rate of 6–7 mm yr-1 (Corti 2009) and should reach a lithospheric rupture in about 10 Ma. These two plates meet the Arabian plate in the north at a triple junction situated in the Ethiopian Afar region (Figure 1.5).
The geological reason and exact mechanisms responsible for the formation of the EAR are debated. Dating volcanic rocks from the Main Ethiopian Rift indicates that the rifting process probably began ~30 Ma ago with the eruption of voluminous flood basalts, which resulted in the formation of the Ethiopian and Somalian plateaus (Corti 2009; Rooney 2017). The basins’ initiations are diachronous along the EAR with distinct episodes of extension followed by periods of relative tectonic quiescence. The fault systems responsible for the formation of the separate rift basins are understood to date from Oligocene and Mio-Pliocene times (Chorowicz 2005).
The observation and mapping of thick volcanic sequences apparently emplaced before the rift-faulting events favored the development of plume theories. Since the observations are not straightforward to interpret and the models multiple, several plume hypotheses have been developed. These invoke the presence of several distinct localized plumes (Kenya, Ethiopia, Arabia) (Montelli et al. 2006), a super-plume (Ebinger and Sleep 1998; Bastow et al. 2008) or a composite scenario including a deep plume originating at the core–mantle boundary and feeding multiple plume-stems in the upper mantle beneath the EAR (Hansen et al. 2012). Tectonic theories focus instead on the
