The East African Rift System: Geodynamics and Natural Resource Potentials

By Athanas Simon Macheyeki and Dalaly Peter Kafumu

The East African Rift System: Geodynamics and Natural Resource Potentials provides state-of-the-art knowledge and skills on how to explore, model, and extract the resources, using the East African Rift System (EARS) as a model. Each aspect to be discussed in the East African Rift System shall have its equivalent case study and readers interested in each rift of the world will find something connected or linked to his/her rift system of interest, be it a sub-chapter on earthquakes, geothermal energy models, etc.

The East African Rift System: Geodynamics and Natural Resource Potentials also describes rifting models of all other known rifts (especially continental rifts) of the world such as the Basin and Range Province, Rio Grande (USA); Rhine Graben (France and Germany); the Tibetan Rohai (Tibet); the Shaanxi Bohai (China); Lake Baikal (Russia); North Island (Australia); and the Aegean Sea Rift (Turkey). Key aspects to be presented shall be: rift type, rift age, rift physical dimensions, geothermal gradient models, natural resources, and models of exploration.

• Connects the science of rift systems to their economic potentials using the East African Rift System as the prime example
• Includes discussions and case studies from rift systems around the world
• Features chapters dedicated to natural resources, such as mineral deposit types (Au, He, REE, U) and the basic principles of their exploration?

• Language: English
• Publisher: Elsevier Science
• Release date: Mar 14, 2024
• ISBN: 9780323956437

Athanas Simon Macheyeki

Athanas S. Macheyeki is a consultant geologist, expert of structural geology and geochemistry, Senior Lecturer at the University of Dodoma and former Commissioner for the Mining (Minerals) Commission in Tanzania. He is also the former CEO of the Tanzania Extractive Industries Transparency Initiative, former Manager of Applied Geology - Geological Survey of Tanzania, former Principal of the Mineral Resources Institute – Dodoma and Founder of the Earth Sciences Institute of Shinyanga (www.esis.ac.tz). Prior to working for the government of Tanzania, he served as Mineral Exploration Geologist for Kabanga Nickel Company Ltd and for Anglo American Exploration Company (Tanzania). He is the holder of BSc. (UDSM), MSc. (UWC) and PhD (UGHENT) and developer of lithogechemical ratios useful for Ni-Cu sulphide exploration. He has authored / co-authored over 20 publications including articles, geological maps, book chapters and a book titled Applied Geochemistry: Advances in Mineral Exploration Techniques.

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Chapter 1

Rifts and rifting

Abstract

Rifts are linear structures on the crust that are involving or have involved the lithosphere in their dynamics. In other words, they are penetrative structures. They are the loci of crustal extension occurring in various tectonic settings ranging from cratonic crust to orogenic settings, meaning that although they are generally considered to be extensional features on the crust, they can also be associated with earlier, contemporaneous, or later compression settings. Rifts occur as narrow (50–100 km wide), elongated (up to 1000 km long) structures and are often segmented. Geomorphologically, rifts are defined as elongated depressions, bounded by normal faults. They are, in many parts, characterized by abnormally high heat flow, causing some parts of the rifts to raise in order to achieve isostatic equilibrium. The elevation of the thermal bulge varies with the activity of the rift system but is usually about 1.5–2 km above the regional average.

Keywords

Rifts; core complex rifts; lithospheric extension; magma; lithosphere

1.1 Rift systems and their origins

Rifts are linear structures on the crust that are involving or have involved the lithosphere in their dynamics (e.g., Braile et al., 2006). In other words, they are penetrative structures. They are the loci of crustal extension (Olsen & Morgan, 2006) occurring in various tectonic settings ranging from cratonic crust to orogenic settings (Roberts & Bally, 2012), meaning that although they are generally considered to be extensional features on the crust, they can as well be associated with earlier, contemporaneous, or later compression settings (e.g., Olsen & Morgan, 2006). Rifts occur as narrow (50–100 km wide), elongated (up to 1000 km long) structures and are often segmented (Leeder, 1995; Olsen & Morgan, 2006). Geomorphologically, rifts are defined as elongated depressions, bounded by normal faults (Allen & Allen, 2005; Olsen, 1995). They are, in many parts, characterized by abnormally high heat flow, causing some parts to raise in order to achieve isostatic equilibrium. The elevation of the thermal bulge varies with the activity of the rift system but is usually about 1.5–2 km above the regional average (Roberts, 2008).

Rifting is a divergent tectonic process that started earlier than 3 Ga (during the formation of Ur supercontinent) and continued to the Holocene (10BP–0BP) (Faure & Mensing, 2010; Li et al., 2008). Rifting is a process that is as old as the continents. Geological and geophysical evidence shows that in a series of repeated accretions and fragmentation of continents, rifting is a major force. For instance, the primordial Paleo-Mesoproterozoic Columbia supercontinent assembled along global-scale collisional orogenic belts around 2.1–1.8 Ga (Faure & Mensing, 2010; Zhao et al., 2004); it contained all of the Earth’s continental blocks (Zhao et al., 2004). The Columbia supercontinent broke apart by rifting along margins 1.6 Ga ago (Li et al., 2008). Other supercontinents include the Neoproterozoic Rodinia supercontinent that assembled 1.1–0.9 Ga ago (Fig. 1.1B) and began to rift apart in the late Neoproterozoic c. 750–633 Ma ago (Gray & Foster, 2005; Zhao et al., 2004), and during the Permian (225 Ma), Pangaea supercontinent was formed (Fig. 1.1A).

Figure 1.1 Relationship between Pangaea and Rodinia supercontinents.

(A) The Pangaea supercontinent accreted from the (B) collisional belt Rodinia Supercontinent. The breakup of Pangaea was initiated by mantle plumes creating rifting forces: Note the positions of mantle plumes (black dots). From Zhao, G., Sun, M., Wilde, S. M., & Li, S. (2004). A Paleo-Mesoproterozoic supercontinent: Assembly, growth and breakup. Earth-Science Reviews, 67(1–2), 91–123. https://doi.org/10.1016/j.earscirev.2004.02.

In the Triassic (200 Ma) to the Jurassic (180 Ma), Pangaea began to break up into two parts; the Northern part was made up of Laurasia containing Laurentia (North America) and the Eurasia (Europe and Asia) continents and the Southern part made up of Gondwana containing Africa, South America, Antarctica, Australia, and India (Roberts, 2012; Zhao et al., 2004). Rifting was followed by drifting. The drifting apart of the Pangaea continued through to Quaternary Period, the process that serves to illustrate here the role of rifting in the formation of continents and rift systems. During the Jurassic Period, several mantle plumes must have developed within the Pangaea continental interior as exemplified by a triple junction joining South America and Africa where two arms spread to form the Atlantic Ocean seafloor and the other arm failed and became the Benue Trough (Windley, 1978). Another example is the Tertiary triple junction that occurred in the Afar Triangle in Ethiopia where the two arms separate (ed) into the Red Sea and the Gulf of Aden, while the third arm failed to spread further creating the Ethiopian Rift (Bosworth & McClay, 2012; Windley, 1978).

As rifting has been a continuous live activity in the geological history it is, therefore, logical that major areas in the world today are currently undergoing continental rifting, and hence stretching (e.g., Buitera et al., 2023). Some of the continental rifts include the EARS, Baikal Rift (in Russia), Rio Grande Rift (in North America), Rhine Rift (between Switzerland, Germany, the Netherlands to the North Sea), West Antarctic Rift System, Basin and Range Province (in USA), the Woodlark Rift (in Papua New Guinea), the Aegean Sea Rift (located in the Aegean Sea separating Greece and Turkey), Bosphorus Rift (in Turkey), Tibetan–Himalayan Rift, and the Shaanxi Bohai Rift or Bohai Bay Basin (in China) (Fig. 1.2A).

Figure 1.2 Overview of past and present rifts and their evolution.

(A) Selected rifts and rifted margins. Currently active rifts cluster in East Africa, Western North America, and Central and Southern Europe. Aborted, so-called failed rift systems achieved significant amounts of crustal thinning but stopped extending before the lithosphere was broken. Rifted margins, by contrast, were formed by rifts that led to continental breakup and can host major sedimentary basins. Colors along rifts and rifted margins indicate the time-averaged rift obliquity from Brune et al. (2012). (B) Conceptual evolution of continental rifts. Inception of rifting involves localization of an array of normal faults that may be accompanied by magmatic intrusions and volcanism. Stretching of the crust results in isostatic subsidence and formation of sedimentary basins. Continental breakup takes place when the continental lithosphere is broken and seafloor spreading initiates. Following breakup, rifted continental margins experience cooling and further subsidence, which may be enhanced due to continuous sedimentation (Buitera et al., 2023). From Buitera, S.J.H., Bruneb, S., Keird, D., & Peron-Pinvidic, G. (2023). Rifting Continents. Chapter 19. Edited by João C. Duarte: Elsevier, Pages 459–481, ISBN 9780323857338. Dynamics of Plate Tectonics and Mantle Convection. https://doi.org/10.1016/B978-0-323-85733-8.00016-0.

When rifts involve the continental lithosphere, they are called continental rifts, or else, if they involve the oceanic lithosphere, they are called oceanic rifts. Continental rifts are nascent plate boundaries where the lithosphere is thinned by tectonic activity (Brune et al., 2022, 2023). As explained earlier, in normal situations, an extending continental rift would later on become an oceanic rift (e.g., Nicolas, 2014), sometimes called a rift basin, and consequently result in continental breakup (e.g., Buitera et al., 2023). The extending region forms a rift valley where sedimentary basins may develop, and if stretching remains localized, the continent breaks and rifted margins form (Fig. 1.2B). Olsen and Morgan (2006) prefer the term oceanic rift valleys instead of oceanic rifts. Detailed classification of rifts by Merle (2011) as subduction-, plume-, mountain-, and transform-related rifts is based on linking rifts to their origin of formation, that is, plate boundaries (Fig. 1.2).

1.2 Types of rifts

1.2.1 Based on mode of extension

Rifts can be classified based on initial thermal conditions and the thickness of crust (e.g., Brune, 2016) involved in the faulting. When the crust is relatively thin and of a typical width of 100–150 km, if the initial lithosphere is cold (Buck, 2015; Hopper & Buck, 1996) and stable, and when the resistance of the brittle upper crust dominates the deformation process (Benes & Davy, 1996), narrow rifts are formed. Narrow rifts are characterized by intense normal faulting, large lateral gradients in crustal thickness and topography, and have a hot (possibly asthenospheric) upper mantle beneath the rifts (e.g., Prodehl et al., 1997). Typical narrow rifts include the East African Rift System, Southern Rio Grande Rift, Baikal Rift, Northern Red Sea (Buck, 1991, 2015), West Antarctic Rift, and the European Cenozoic Rift System (Corti et al., 2003; Rosendahl, 1987) (Fig. 1.3).

Figure 1.3 Narrow, wide, and core complex rifts. After Buck, W.R. (2015). The dynamics of continental breakup and extension. Treatise on Geophysics, 325–379. https://doi.org/10.1016/b978-0-444-53802-4.0.

Wide rifts are rifts that are typically expressed on the surface by a large number of separated basins in a relatively wider region typically over 1000 km wide—a width that is greater than the thickness of lithosphere (e.g., Hopper & Buck, 1996). Wide rifts will form if the lithosphere is initially warmer and unstable in terms of extension (Benes & Davy, 1996; Hopper & Buck, 1996) and when the brittle upper crust is thin and characterized by ductile flow in the lower crust allowing expansion of the deformation zone (e.g., Buck, 1991). Periodic instabilities with two wavelengths dominate wide rifting (Benes & Davy, 1996). Examples of wide rifts are the Basin and Range Province (Buck), (Figs. 1.2 and 1.3).

Core complex rifts result from a mode of extension that is transitional between unstable and stable lithospheric extension, that is, it does not correspond to a particular mode of extension (Benes & Davy, 1996; Brun, 1999) but rather occurs due to local anomalies within wide rifts (Brun, 1999). Core complex rifts begin as distributed zones of unstable extension, due to thin upper brittle crust but lower crustal flow induced by the gravity force, causing a re-localization of strain into narrow zones (e.g., Collettini et al., 2009; Huismans & Beaumont, 2003; Martinez & Taylor, 2002, 2003). According to Brun (1999), models suggest that core complex rifts probably develop in zones of the upper crust located above heterogeneities of the ductile lower crust that are weak enough to localize stretching. Consequently, local stronger thinning of the brittle upper crust is compensated by the uprise and exhumation of the ductile lower crust. In other words, core complex rifts consist of high-grade metamorphic rocks originating in the middle to lower crust and which are exposed at the surface, exhumed by low-angle normal faults, uplift, and erosion (e.g., Brune, 2014; Cheney, 1980). Examples of rifts that fall in this category include the Tibetan and the Aegean rift (Figs. 1.2–1.5).

Figure 1.4 The main features of a metamorphic core complex.

Simplified sketch showing the main features of a metamorphic core complex. From Tirel et al. (2009, Fig. 1.1).’– with the reference to ‘Tirel, C., Gautier, P., van Hinsbergen, D. J. J., & Wortel, M. J. R. (2009). Sequential development of interfering metamorphic core complexes: Numerical experiments and comparison with the Cyclades, Greece. Geological Society, London, Special Publications, 311, 257–292, https://doi.org/10.1144/SP311.10.

Figure 1.5 Predicted modes of extension as a function of crustal thickness and surface heat flow compared to data for regions showing those modes of extension, or in the case of the Altiplano and Tibet having the conditions that might lead to core complex formation. After Buck, W.R. (2015). The dynamics of continental breakup and extension. Treatise on Geophysics, 325–379. https://doi.org/10.1016/b978-0-444-53802-4.0.

Buck (1991, 2015) provide ranges of heat flow in mWm−2 versus thickness in km conditions for predicted modes of extension as a function of crustal thickness and surface heat flow (Fig. 1.5).

Understanding the lithospheric extension requires one to understand well the processes of uplift, faulting, and magmatism (Baldridge et al., 2006) and the timing of the same. The state of the lithosphere (cold or warm) affects its response to rifting. The Baikal rift zone for example appears to have formed in a cold, stable lithosphere, while the Rio Grande rift formed in a region that had experienced almost continuous deformation and magmatism for the past 50 Ma. Differences in preexisting conditions such as these may explain some of the differences and similarities among rifts or rift zones (e.g., Fig. 1.5).

1.2.2 Based on geometry and kinematics

Rifts can also be classified as orthogonal or oblique based on their structure, geometry, and kinematics, that is, according to the environment of the overall displacement and strain field in which they form (e.g., Busby & Ingersoll, 1995; McClay & White, 1995). The length of fault segments in rifts is inversely proportional to the degree of obliquity of the rift (e.g., McClay et al., 2002), and the lateral propagation of rift segments results in the merging of the different rift basins in one single rift, in which the initial en échelon geometry remains visible (Mart & Dauteuil, 2000; McClay & White, 1995).

When long rift border faults and relatively straight with shorter intrarift faults that are perpendicular to the direction of extension (i.e., α, the angle between the rift axis and the extension direction=90 degrees), they are referred to as orthogonal rifts (e.g., McClay & White, 1995; Morley et al., 2004; Zwaan et al., 2016).

During orthogonal rifting, magma is passively squeezed away from the rift axis and forced to go on the sides of the rift and extrude there—such a phenomenon results in off-axis magmatism (Fig. 1.6A) – a typical case observed in the Ethiopian rift, the Limagne Graben, and the Red Sea (Corti et al., 2003).

Figure 1.6 Horizontally sliced models of three-dimensional rift experiments.

Interpretations of horizontally sliced models of three-dimensional rift experiments; (A) orthogonal extension, α=90 degrees and (B) oblique extension, α=60 degrees. In the orthogonal rift, long linear faults develop that are oriented perpendicular to the direction of extension, whereas, in the oblique model, different smaller faults form in en échelon geometry. Gray shadings in the figure represent local depocentres. Arrows point to the extension direction. From McClay, K.R., & White, M.J. (1995). Analog modeling of orthogonal and oblique rifting. Marine and Petroleum Geology, 12, 137–15.

On the other hand, when the rift border faults with the direction of extension are not perpendicular to each other, that is, α ≠ 90 degrees, such rifts are called oblique rifts (e.g., Corti et al., 2003). The rift border faults are generally highly segmented forming en échelon arrays parallel to the zone of rifting and the intrarift faults are either perpendicular to the rift border faults or could be at higher angles (Duclaux et al., 2019; Fig. 1.6B).

Unlike in the orthogonal rifting where magma are passively squeezed off the axis, during oblique rifting, magma is emplaced within the main rift axis (depression) portraying elongated magma intrusions with an oblique, and in some cases, an en echelon pattern (Corti et al., 2003; Hus, 2004). During oblique rifting, the magma at depth influences the fault pattern localizing strain in the overlying crust—a typical phenomenon of the Main Ethiopian rift (Figs. 1.7 and 1.8). Increased extension in the oblique and moderately oblique rift systems (α=45 and 60 degrees) causes the intrarift faults to rotate toward parallelism with the rift border faults (Morley, 1999). Multiphase phenomena can also occur when orthogonal rifting precedes oblique rifting (McClay & White, 1995; Morley, 1999).

Figure 1.7 Digital elevation model of oblique rifts that developed within oblique weak zones.

(A) Main Ethiopian rift (Agostini et al., 2011; Philippon et al., 2014). (B) Western and eastern branches of East African rift system (Delvaux & Barth, 2010). (C) Baikal rift (Petit et al., 1996). White arrows indicate regional direction of extension (GPS motion) and red arrows indicate local minimum horizontal stress deduced from either fault slip data or earthquake focal mechanisms. From Philippon, M., Willingshofer, E., Sokoutis, D., Corti, G., Sani, F., Bonini, M., & Cloetingh, S. (2015). Slip re-orientation in oblique rifts. Geology, 43, 147–150. https://doi.org/10.1130/G36208.1: Courtesy to the Geological Society of America.

Figure 1.8 Interpretative schematic 3D block diagram of the Main Ethiopian Rift evolution based on the modeling results, relating (1) off-axis volcanism to lateral flow of the lower crust and magma during the first orthogonal rifting phase (denoted by 1st) and (2) oblique faulting and magmatism to the second oblique rifting phase (denoted by 2nd). Note that this interpretation includes both the effect of isostatic compensation provided by the asthenosphere and the thermal thinning of the crust, which were not considered in the current models. UC, upper crust; LC, lower crust; m, magma; rf, rift fill; WFB, Wonji fault belt. From Corti, G., Bonini, M., Conticelli, S., Innocenti, F., Manetti, P., & Sokoutis, D. (2003). Analogue modelling of continental extension: A review focused on the relations between the patterns of deformation and the presence of magma. Earth-Science Reviews, 63, 169–247. https://doi.org/10.1016/S0012-8252(03)00035-7.

1.3 Rifting and their driving forces

Two main models can be considered for the rifting processes, that is, the far-field model and the mantle plume (Sengor & Burke, 1978). In either model, the lithosphere is stretched and thinned (Gao et al., 2013; Huismans et al., 2001; Koptev et al., 2016). In the far-field model, passive rifts are formed, whereas in the mantle plume model, active rifts are formed (Koptev et al., 2016). The mantle upwelling triggers partial melting in the upper mantle causing volcanic activities in the rift (Nicolas, 2014), and in situations where copious magmatic activities are happening in the rift, active rifting is accelerated even further because twice more melt is generated by a 100°C overheated mantle plume compared to normal rising asthenosphere (Nicolas, 2014; White & McKenzie, 1989). The upwelling overheated mantle plume creates a dome that may further develop into a triple junction in which case (if rifting proceeds further) two of the arms can form a seafloor spreading rift as two continents separated by an ocean and the other arm fails and becomes a graben or a trough (DiPietro, 2018) (Fig. 1.9). According to Olsen and Morgan (2006), the lower crust and/or the lower lithospheric mantle is more likely to deform in a ductile manner and that rupturing of the lithospheric may not occur until dike intrusion progresses to a stage of seafloor spreading, that is, until all crustal extension is accommodated by the intrusion of new, mantle-derived basaltic crust. There is a close relationship between lithospheric strength and the presence of magmatic bodies in or beneath a lithosphere (e.g., Buck, 2015). When a large volume of melt is emplaced in the continental lithosphere, the thermal and mechanical properties of the lithosphere are modified; this weakens the lithosphere and enhances deformation and strain localization (e.g., Morley, 1999). Without considering the rate of strain rate of the lithosphere and its corresponding yield stress, Buck (2015) shows how the strength of the lithosphere is reduced (weakened) due to advective lithospheric thinning (Fig. 1.10).

Figure 1.9 Rifting in map view and in cross-section.

(A) The map shows the initial development of a triple junction. The cross-section below the map shows bulging and stretching of the crust and development of normal faults and volcanism. (B) The map shows development of a divergent plate boundary between two continental fragments as well as a failed rift. The cross-section shows development of a passive continental margin on both sides of the divergent plate boundary based on Van der Pluim and Marshak (2004). From DiPietro, J.A., 2018. Forcing Agent. Geology and Landscape Evolution, Elsevier, pp. 59–77. https://doi.org/10.1016/b978-0-12-811191-8.0.

Figure 1.10 Plots on the left show schematic illustrations of processes that may lead to localization of strain during continental lithospheric extension.

Plots to the right show the approximate distribution of yield stress with depth both at the center of a rift (A) and at an area unaffected by the rifting (B). The difference in the two curves is marked with vertical hatchers and that area is proportional to the change, here reduction, in the tectonic force, needed for continued rifting. The scaling of these force changes is given within ovals. See the text for further explanation (Buck, 2015). ε, shear strain ratio; Hb, thickness of the brittle lithosphere; ΔFN, change in lithospheric strength due to extension; ΔFM, change in the tectonic force for an extension due to magma; ΔFC, change in the tectonic force for an extension due to crustal thinning or due to loss of cohesion; Δxd, lithospheric thickness; T*m, an effective temperature of magma; CHb, the yield stress at the base of the lithosphere; and ΔFl=change in local crustal buoyancy force; S, loss of cohesive strength. From Buck, W.R., 2015. The dynamics of continental breakup and extension. Treatise on Geophysics, 325–379.