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 are often described as bowl-shaped depressions hundreds of kilometers across. The sedimentary successions correspond to terrestrial and shallow-water sediments (carbonates, shales, sandstones), typically exhibiting a layer cake infill architecture, supposedly thicker and more complete than in adjacent areas of the craton but still relatively thin. The accumulation rates are extremely slow (2–20 m/Ma), and sedimentation usually keeps pace with basin subsidence throughout the rift history. A key characteristic of intracratonic basins is that they preserve their stratigraphy over very long time periods.
Other than the Congo Basin, other examples of intracratonic basins include the Michigan basin, the East Barents Sea and the Chad basin.
Figure 1.10. Illustration of the intracratonic rift basin case
Case example: The Congo Basin
The Congo Basin is a large, circular intracratonic sedimentary basin located in Central Africa (Figure 1.11) (Giresse 2005; Kadima et al. 2011; Watts et al. 2018). The depth of the lithosphere–asthenosphere boundary is modeled at 220 ± 30 km (Crosby et al. 2010). The basin is composed of at least three sub-basins separated by ~100 km wide basement ridges. The dimensions of the tilted basement blocks are unusually wide, theorized to be a result of the cold thermal gradients and thus higher strength of the brittle lithosphere (Foster and Nimmo 1996). The average sedimentary thickness is about 4 km, but up to 9 km of sedimentary rocks have been observed locally. The stratigraphic record spans 0.55 Ga, starting from the late Pre-Cambrian, and the basin is still subsiding today (Daly et al. 1992). The drilled stratigraphic sequences show consistently lagoonal and shallow-marine environments, with some truncation events that are correlated to far-field tectonic events (Daly et al. 1992). No magmatism has been reported, except at Early Cenozoic times with the emplacement of kimberlites at the borders of the basin.
1.2.2.2. Rifts at convergent plate boundaries
Sedimentary basins can also develop in close spatial and temporal relation to an active orogen in a convergent tectonic setting. These include basins in trenches, fore-arcs and foreland, intra-arcs, back-arcs/hinterland, retro-arcs/-foreland, pro-foreland and wedge-top regions. The list of basins in convergent settings is very long, often confusing and sometimes misleading. The various descriptors can identify the different structural settings in which the basins form (e.g. if the basin develops over the subducting plate, on the edge of the orogen, far from the orogen, within the orogen, etc.), and/or the different types of substratum (e.g. continental or oceanic crust). However, given the regular structural complexity of orogenic settings, the description of the basin into one category or the other is often not straightforward.
In general, basins in convergent plate boundaries can be produced in areas under local extension or in areas under compression where accommodation space is generated. For instance, extension can occur in a convergent tectonic boundary in the back-arc region (Figure 1.12). Typical examples of basins related to orogenesis include the Upper Rhine Graben, the North Alpine foreland basin, the Ganges basin and the Andean basins.
In a post-orogenic context, the overthickened lithosphere tends to a re-equilibration of its gravitational and thermal status by lithosphere thinning and, hence, extension (e.g. Dewey 1988; Rey et al. 2001). This is the so-called orogenic collapse, a specific tectonic mechanism that can lead to the formation of extensional basins inside the orogen (see Chapter 2).
Figure 1.11. a) Topographic map of central Africa showing the outline of the main cuvette of the Congo Basin (blue line). The inset globe shows worldwide location; b) sediment thickness map of the Congo Basin. The red line shows the location of the c) seismic profile, and the thick black line shows the d) section; c) seismic profile, without (left) and with (right) vertical exaggeration. BCBS: base of the cratonic basin sequence and d) geological cross-section showing the stratigraphic architecture of the Congo Basin (source: Watts et al. 2018)
Back-arc basins can be further highlighted, as these can evolve into actual rifts and rifted margins. Back-arc basins form above subduction zones in response to tensional forces generated by the down-going slab and related dynamics, and/or the collapse of the adjacent edge of the continent (Figure 1.12) (Stern 2002). Back-arc basins have peculiar geometries; they are usually very long (several hundreds to thousands of kilometers) and relatively narrow (a few hundred kilometers) and are often associated with magmatism and significant hydrothermal activity. Basalts in these basins have geochemical signatures similar to mid-ocean ridge basalts (MORBs), although they are significantly richer in magmatic water. In extreme cases, the back-arc rifting can be very pronounced, with a significant rise of the asthenosphere leading to a lithospheric breakup and the genesis of new oceanic crust. Typical examples of active back-arc basins include the Tyrrhenian Sea, the Mariana Trough, the Black Sea, Tonga-Kermadec, the south Fiji basin, with most being found along the western margin of the Pacific Ocean.
Figure 1.12. Illustration of the back arc, forearc and trench basins related to a subduction zone. The figure is a schematic section through the upper 140 km of a subduction zone, showing the principal structural components (source: modified from Stern 2002)
1.2.2.3. Rifts at divergent-plate boundaries
Rifted margins – also called “continental margins”, “passive margins” or “Atlantic-type margins” – were defined by Ruppel (1995) as regions “where continental lithosphere meets oceanic lithosphere without an intervening plate boundary”. This rift category can be considered as the successful version of the intracontinental rifts, in the sense that the extensional processes continued on to the ultimate separation of the continental lithosphere and the emplacement of an oceanic spreading system. Rifted margins are then an archive of the entire spectrum
