Km

Figure 7.3 (a) Major faults and segmentation pattern of the northern Main Ethiopian rift and (b) cross-section of Adama Rift Basin showing half graben morphology (images provided by C. Ebinger and modified from Wolfenden et al., 2004, with permission from Elsevier). MS, magmatic segment; BF, border fault. In (b) note the wedge-shaped geometry of the syn-rift Miocene and younger ignimbrite and volcanic units (vertical lined pattern and upper shaded layer). Pre-rift Oligocene flood basalts (lowest shaded layer) show uniform thickness.

(Fig. 7.4b). Inside the rift, earthquake clusters parallel faults and volcanic centers in a series of 20 km wide, right-stepping zones of magmatism (Fig. 7.4c). Up to 80% of the total extensional strain is localized within these magmatic segments (Bilham et al.,

1999; Ebinger & Casey, 2001). The largest earthquakes typically occur along or near major border faults, although the seismicity data indicate that the border faults are mostly aseismic. Earthquakes are concentrated around volcanoes and fissures at depths of

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Figure 7.4 (a) Seismicity and focal mechanisms of East Africa between I960 and 2005. Late Cenozoic volcanoes shown by triangles. (b) Seismicity of rift segments in the northern Main Ethiopian Rift (MER) between October 200I and January 2003. (c) Faults that cut <I.9Ma lavas and late Cenozoic eruptive centers comprising magmatic segments (MS). Miocene border faults in rift basins also shown. Size of focal mechanism solutions indicates relative magnitude of earthquakes. Black arrows show approximate range of plate velocity vectors derived from geodetic data. (d) Earthquake depth distribution across profile A-A' shown in (c). Cross-section B-B' shown in Fig. 7.5 (images provided by D. Keir and modified from Keir et al., 2006, by permission of the American Geophysical Union. Copyright © 2006 American Geophysical Union).

Figure 7.4 (a) Seismicity and focal mechanisms of East Africa between I960 and 2005. Late Cenozoic volcanoes shown by triangles. (b) Seismicity of rift segments in the northern Main Ethiopian Rift (MER) between October 200I and January 2003. (c) Faults that cut <I.9Ma lavas and late Cenozoic eruptive centers comprising magmatic segments (MS). Miocene border faults in rift basins also shown. Size of focal mechanism solutions indicates relative magnitude of earthquakes. Black arrows show approximate range of plate velocity vectors derived from geodetic data. (d) Earthquake depth distribution across profile A-A' shown in (c). Cross-section B-B' shown in Fig. 7.5 (images provided by D. Keir and modified from Keir et al., 2006, by permission of the American Geophysical Union. Copyright © 2006 American Geophysical Union).

less than 14 km (Fig. 7.4d), probably reflecting magma movement in dikes. In the rift flanks, seismic activity may reflect flexure of the crust (Section 7.6.4) as well as movement along faults. The orientation of the minimum compressive stress determined from earthquake focal mechanisms is approximately horizontal, parallel to an azimuth of 103°. This stress direction, like that in Afar, is consistent with determinations of extension directions derived from tension fractures in young <7000 year old lavas, geodetic measurements, and global plate kinematic data (Fig. 7.4c).

3 Local crustal thinning modified by magmatic activity. Geophysical data indicate that continental rifts are characterized by thinning of the crust beneath the rift axis. Crustal thicknesses, like the fault geometries in rift basins, are variable and may be asymmetric. Thick crust may occur beneath the rift flanks as a result of magmatic intrusions indicating that crustal thinning is mostly a local phenomenon (Mackenzie et al., 2005; Tiberi et al., 2005). Variations in crustal thickness may also reflect inherited (pre-rift) structural differences.

Mackenzie et al. (2005) used the results of controlled-source seismic refraction and seismic reflection studies to determine the crustal velocity structure beneath the Adama Rift Basin in the northern part of the Main Ethiopian Rift (Fig. 7.5a). Their velocity model shows an asymmetric crustal structure with maximum thinning occurring slightly west of the rift valley. A thin low velocity layer (3.3 km s-1) occurs within the rift valley and thickens eastward from 1 to 2.5 km. A 2-5-km-thick sequence of intermediate velocity (4.5-5.5 km s-1) sedimentary and volcanic rock lies below the low velocity layer and extends along the length of the profile. Normal crustal velocities (Pn = 6.0-6.8 km s-1) occur to depths of 30-35 km except in a narrow 20-30 km wide region in the upper crust beneath the center of the rift valley where Pn velocities are 5-10% higher (>6.5 km s-1) than those outside the rift (Fig. 7.5a). These differences probably reflect the presence of mafic intrusions associated with magmatic centers. A nearly continuous intracrustal reflector at 20-25 km depth and Moho depths of 30 km show crustal thinning beneath the rift axis. The western flank of the rift is underlain by a ~45 km thick crust and displays a ~15 km thick high velocity (7.4 km s-1) lower crustal layer. This layer is absent from the eastern side, where the crust is some 35 km thick. Mackenzie et al. (2005) interpreted the high velocity lower crustal layer beneath the western flank as underplated material associated with pre-rift Oligocene flood basalts and, possibly, more recent magmatic activity. Variations in intracrustal seismic reflectivity also suggest the presence of igneous intrusions directly below the rift valley (Fig. 7.5b).

Gravity data provide additional evidence that the crustal structure of rift zones is permanently modified by magmatism that occurs both prior to and during rifting. In Ethiopia and Kenya, two long-wavelength (>1000 km) negative Bouguer gravity anomalies coincide with two major ~2 km high topographic uplifts: the Ethiopian Plateau and the Kenya Dome, which forms part of the East African Plateau (Figs 7.2, 7.6a). The highest parts of the Ethiopian Plateau are more than 3 km high. This great height results from the eruption of a large volume of continental flood basalts (Section 7.4) between 45 and 22 Ma, with the majority of volcanism coinciding with the opening of the Red Sea and Gulf of Aden at ~30 Ma (Wolfenden et al., 2005). The negative gravity anomalies reflect the presence of anomalously low density upper mantle and elevated geotherms (Tessema & Antoine, 2004). In each zone, the rift valleys display short-wavelength positive Bouguer gravity anomalies (Fig. 7.6b) that reflect the presence of cooled, dense mafic intrusions (Tiberi et al., 2005).

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Figure 7.5 (a) P-wave velocity model and (b) interpretation of the Main Ethiopian Rift (after Mackenzie et al., 2005, with permission from Blackwell Publishing). Location of profile (B-B') shown in Fig. 7.4c.

4 High heat flow and low velocity, low density upper mantle. Heat flow measurements averaging 70-90 mW m-2 and low seismic velocities in many rift basins suggest temperature gradients (50-100°C km-1) that are higher than those in the adjacent rift flanks and nearby cratons. Where the asthenosphere is anomalously hot, such as in East Africa, domal uplifts and pervasive volcanism result. Nevertheless, there is a large degree of variability in temperature and volcanic activity among rifts. The Baikal Rift, for example, is much cooler. This rift displays low regional heat flow of 40-60 mW m-2 (Lysack, 1992) and lacks volcanic activity.

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Figure 7.6 (a) Map showing long-wavelength Bouguer gravity anomalies after removal of the short wavelength (image provided by A. Tessema and modified from Tessema & Antoine, 2004, with permission from Elsevier). (b) Profiles (A-A') of topography, short-wavelength Bouguer gravity anomaly, and crustal thickness estimates of the central Main Ethiopian Rift (MER) (images provided by C. Tiberi and modified from Tiberi et al., 2005, with permission from Blackwell Publishing). Profile location shown in (a). Circles in crustal thickness profile indicate depths estimated from receiver function studies.

In East Africa, relatively slow Pn wave velocities of 7.7 km s-1 in the upper mantle beneath the Adama Rift Basin in Ethiopia (Fig. 7.5a) suggest elevated temperatures (Mackenzie et al., 2005). Elsewhere upper mantle Pn wave velocities are in the range 8.0-8.1 km s-1, which is expected for stable areas with normal heat flow. Tomographic inversion of P- and S-wave data (Fig. 7.7a-c) indicate that the low velocity zone below the rift is tabular, approximately 75 km wide, and extends to depths of 200-250 km (Bastow et al., 2005). The zone is segmented and offset away from the rift axis in the upper 100 km but becomes more central about the rift axis below this depth (Fig. 7.7c). In the more highly extended northern section of Ethiopian Rift (Fig. 7.7d), the low velocity anomaly broadens laterally below 100 km and may be connected to deeper low velocity structures beneath the Afar Depression (Section 7.4.3). This broadening of the low velocity zone is consistent with the propagation of the Main Ethiopian Rift, during Pliocene-Recent times, toward the older spreading centers of the Red Sea and Gulf of Aden.

In addition to the high temperatures, the low velocity zones beneath rifts may also reflect the

P-wave % velocity anomaly (b) 75 km

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