The East African Rift system

The East African Rift system (Fig. 7.2) is composed of several discrete rift segments that record different stages in the transition from continental rift to rifted volcanic margin (Ebinger, 2005). The Eastern Rift between northern Tanzania and southern Kenya is an example of a youthful rift that initiated in thick, cold and strong continental lithosphere. Volcanism and sedimentation began by ~5 Ma with the largest fault escarpments forming by ~3 Ma. Strain and magmatism are localized within narrow asymmetric rift basins with no detectable deformation in the broad uplifted plateau adjacent to the rifts (Foster et al., 1997). Earthquake hypocenters occur throughout the entire 35 km thickness of the crust, indicating that crustal heating is at a minimum (Foster & Jackson, 1998). The basins are shallow (~3 km deep) with 100-km-long border faults that accommodate small amounts of extension. The border faults have grown from short fault segments that propagated along their lengths to join with other nearby faults, creating linkages between adjacent basins (Foster et al., 1997). Faults that were oriented unfavorably with respect to the opening direction were abandoned as strain progressively localized onto the border faults (Ebinger, 2005). Geophysical (Green et al., 1991; Birt et al., 1997) and geochemical (Chesley et al., 1999) data show that the mantle lithosphere has been thinned to about 140 km. Elsewhere the lithosphere is at least 200 km and possibly 300-350 km thick (Ritsema et al., 1998). These patterns conform to the predictions of lithospheric stretching models (Section 7.6.2, 7.6.3) in regions of relatively thick lithosphere. They also illustrate that the cross-sectional geometry and the along-axis segmentation in youthful rifts are controlled by the flexural strength of the lithosphere (Section 7.6.4).

The effects of pre-existing weaknesses on the geometry of rifting are also illustrated in the southern segment of the Eastern Rift in Tanzania. Border faults and half graben preferentially formed in a zone of weakness created by a contrast between thick, cool lithosphere of the Archean Tanzanian craton and thin, weak Proterozoic lithosphere located to the east (Foster et al., 1997). From north to south, the axis of the rift diverges from a single ~50 km wide rift to a ~200 km wide zone composed of three narrow segments (Fig. 7.2b). This segmentation and a change in orientation of faults occurs where the rift encounters the Archean Tanzanian craton (Fig. 7.37), indicating that the thick lithosphere has deflected the orientation of the rift. These observations illustrate that lateral heterogeneities at the lithosphere-asthenosphere boundary exert a strong control on the initial location and distribution of strain at the start of rifting (Section 7.4).

An example of a rift that is slightly more evolved than the Tanzanian example occurs in central and northern Kenya where rifting began by 15 Ma. In this rift segment the crust has been thinned by up to 10 km and the thickness of the lithosphere has been reduced to about 90 km (Mechie et al., 1997). A progressive shallowing of the Moho occurs between central and northern Kenya where the rift widens from ~100 km to ~175 km (Fig. 7.2b). In northern Kenya, crustal thickness is about 20 km and the total surface extension is about 35-40 km (P = 1.55-1.65) (Hendrie et al., 1994). In the south, crustal thickness is 35 km with estimates of total extension ranging from 5 to 10 km (Strecker et al., 1990; Green et al., 1991). As the amount of crustal stretching increases, and the lithosphere-asthenosphere boundary rises beneath a rift, the amount of partial melting resulting from decompression melting also increases (Section 7.4.2). Young lavas exposed in central and northern Kenya indicate source regions that are shallower than those in Tanzania (Furman et al., 2004). High velocity, high density material is present in the upper crust and at the base of the lower crust, suggesting the presence of cooled basaltic intrusions (Mechie et al., 1997; Ibs-von Seht et al., 2001). These relationships indicate that as a continental rift enters maturity mag-matic activity increases and a significant component of the extension is accommodated by magmatic intrusion below the rift axis (Ebinger, 2005).

Figure 7.37 Structural map of the Eastern branch of the East African Rift system in Kenya and northern Tanzania showing the deflection of faults at the boundary of the Archean Tanzanian craton (after Macdonald et al., 2001, with permission from the Journal of Petrology 42,877-900. Copyright © 2001 by permission of Oxford University Press, and Smith & Mosely, 1993, by permission of the American Geophysical Union. Copyright © 1993 American Geophysical Union).

Figure 7.37 Structural map of the Eastern branch of the East African Rift system in Kenya and northern Tanzania showing the deflection of faults at the boundary of the Archean Tanzanian craton (after Macdonald et al., 2001, with permission from the Journal of Petrology 42,877-900. Copyright © 2001 by permission of Oxford University Press, and Smith & Mosely, 1993, by permission of the American Geophysical Union. Copyright © 1993 American Geophysical Union).

The increase in magmatic activity that accompanies a shallowing of the asthenosphere-lithosphere boundary beneath the Kenya Rift also results in increased crustal heating and contributes to a decrease in lithospheric strength (Section 7.6.7). This effect is indicated by a progressive decrease in the depth of earthquake hypocenters and in the depth of faulting from 35 km to 27 km (Ibs-von Seht et al., 2001). These patterns suggest a decrease in the effective elastic thickness (Te) of the lithosphere (Section 7.6.4) compared to the rift in northern Tanzania. Although both the mantle and crust have thinned, the thinning of mantle lithosphere outpaces crustal thinning. This asymmetry occurs because a sufficient amount of magma has accreted to the base of the crust, resulting in a degree of crustal thickening. It also results because the mantle lithosphere is locally weakened by interactions with hot magmatic fluids, which further localizes stretching.

Extension in the central and southern part of the Main Ethiopian Rift began between 18 and 15 Ma and, in the north, after 11 Ma (Wolfenden et al., 2004). The deformation resulted in the formation of a series of high-angle border faults that are marked by chains of volcanic centers (Fig. 7.38a). Since about 1.8 Ma the loci of magmatism and faulting have become progressively more localized, concentrating into ~20-km-wide, 60-km-long magmatic segments (Fig. 7.38b). This localization involved the formation of new, shorter and narrower rift segments that are superimposed on old long border faults in an old broad rift basin. This narrowing of the axis into short segments reflects a plate whose effective elastic thickness is less than it was when the long border faults formed (Ebinger et al., 1999). The extrusion of copious amounts of volcanic rock also has modified both the surface morphology of the rift and its internal structure. Relationships in this rift segment indicate that magma intrusion in the form of vertical dikes first becomes equally and then more important than faulting as rifting approaches sea floor spreading (Kendall et al., 2005). Repeated eruptions create thick piles that load the weakened plate causing older lava flows to bend down toward the rift axis. This process creates the seaward-dipping wedge of lavas (Section 7.7.1) that is typical of rifted volcanic margins (Section 7.7.1).

The rift segments in the Afar Depression illustrate that, as extension increases and the thickness of the lithosphere decreases, the asthenosphere rises and decompresses, and more melt is generated. Eventually all the border faults in the rift are abandoned as mag-matism accommodates the extension (Fig. 7.38c). At this stage the rift functions as a slow-spreading mid-ocean ridge that is bordered on both sides by thinned continental lithosphere (Wolfenden et al., 2005). As the melt supply increases and/or strain rate increases, new oceanic lithosphere forms in the magmatic segments and the crust and mantle lithosphere subside below sea level. This transition has occurred in the Gulf of Aden

fault

Crust 2 km

fault

Crust 2 km

Magmatic segments

Magmatic segments

Figure 7.38 Three-stage model for continental break-up leading to the formation of a volcanic passive margin (after Ebinger, 2005, with permission from Blackwell Publishing).

Figure 7.38 Three-stage model for continental break-up leading to the formation of a volcanic passive margin (after Ebinger, 2005, with permission from Blackwell Publishing).

(Fig. 7.2b) where conjugate rifted margins have formed recently. The margins on the western side of the Gulf are mostly buried by Oligocene-Miocene lavas from the Afar mantle plume. Those on the eastern side are starved of sediment and volcanic material and preserve 19-35 Ma structures that formed during oblique rifting and the transition to sea floor spreading (d'Acremont et al., 2005). Seismic reflection studies of these latter margins indicate that the southern rifted margin is about twice as wide as the northern one and displays thicker post-rift deposits and greater amounts of subsidence. As rifting gave way to sea floor spreading in this area, deformation localized in a 40-km-wide transition zone where magma intruded into very thin continental crust and, possibly, in the case of the northern side, exhumed mantle. The different widths and structure of the two margins indicate that the transition to sea floor spreading in the Gulf of Aden was an asymmetric process.

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