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1.1. Introduction
The Earth’s interior is composed of multiple layers of differing composition and thickness. The exterior layers are the oceanic and continental crust, followed by the upper mantle – together constituting the lithosphere (Figure 1.1). The Earth’s lithosphere is divided into tectonic plates that are in constant motion relative to each other, following a repeated evolution of opening, widening and closing oceanic basins. This plate tectonic cycle became known as the “Wilson Cycle” in honor of J. Tuzo Wilson (1908–1993), who actively participated in the important evolution of knowledge and understanding achieved in geoscience in the 1960s (Wilson 1966). By this point, breakthroughs in geophysical techniques had produced observations of magnetic anomalies, seafloor spreading ridges and hot spots that had established plate tectonics and plate motions as fact (e.g. Atwater 1970; Molnar Atwater 1973). This, in turn, allowed J.T. Wilson to identify the systematic separation and reassembly of continental masses.
Figure 1.1. Schematic representation of the main layers constituting the Earth
Within the Wilson Cycle, continental rifting corresponds to the extensional stage that leads to thinning and rupture of the continental lithosphere (Figure 1.2). The primary product of rifting is the rift, which may subsequently become a rifted margin if the extensional processes lead to ridge emplacement, genesis of oceanic crust and the opening of an ocean basin.
Rifts were identified as unique geological occurrences before plate tectonics were established (e.g. Suess 1891; Gregory 1896; de Lapparent 1898; Willis 1928; Bullard 1936). Initially, they were described as topographic depressions flanked by arrays of subvertical faults (e.g. Gregory 1923). It was not until the last century that rifts were recognized as extensional features rather than as results of compressional deformation (e.g. Vening Meinesz 1950).
Figure 1.2. Schematic representation of the Wilson Cycle in eight steps illustrating the various stages of the successive opening and closing of ocean basins (source: modified after Wilson et al. 2019)
Rifts can be summarized as geographical regions consisting of extensional sedimentary basins of various sizes, with various tectonic and sedimentary geometries that are linked in various structural contexts. The repeated use of the adjective “various” in this basic definition leads to a description that may sound rather vague and unprecise. This is deliberate, since the extension applied to a continental domain can lead to the formation of multiple structures with different geometries. If rifts are defined as topographic depressions that are created by normal faulting and associated subsidence, then the dimensions, physical characteristics and linkage characteristics between the different extended subdomains can span a wide variety of options. Because the range of observed architectural possibilities in rifts is broad, numerous classification systems for rifts have been proposed in the literature: passive versus active, narrow versus wide, intracontinental, diffuse, synorogenic and so on. However, even though each rift type may appear unique, a series of geologic features are commonly observed in most rift settings. These include specific basin types (grabens, half-grabens, pull-apart/strike-slip basins), extensional structures (normal faults, detachment faults, core complexes), sedimentary geometries (growth structures, unconformities) and salt- and magma-related features (e.g. Neumann and Ramberg 1978; Keen 1985; Ziegler 1992; Ruppel 1995; Corti et al. 2003; Merle 2011).
This chapter provides a brief summary of these features, a list of the main types of rifts and case examples. Our aim is to clarify the basic notions required to study rifted margins – this is done to avoid any misunderstanding of the use of some terms in the following chapters. Regularly, a Further reading section will be provided as an illustration of the many contributions published on the various topics. The reader interested in finding more in-depth explanations is referred to these contributions for further detailed information.
1.2. Rift classifications
This section begins with the definition of the active and passive rifting categories, as these are regularly mentioned in the literature and are the primary designations of extensional rift settings. However, because these two categories are very loosely defined and present-day data coverage is limited, it is difficult to fit all rifts in this classification scheme. For this reason, we prefer a classification that is based on the general tectonic setting of the rifts, as proposed by Ruppel (1995), which will be described in section 2.2.
1.2.1. Rift mechanism classification: active versus passive rifting
Originally, rifts were classified by the process of lithospheric thinning: in active rifts, deformation was thought of as driven by active asthenospheric upwelling, whereas in passive rifts, the asthenospheric upwelling was itself viewed as a passive response to lithospheric extension (Neumann and Ramberg 1978; Şengör and Burke 1978) (Figure 1.3).
Active rifting processes are governed by the upwelling of warm asthenospheric mantle beneath the base of the continental lithosphere, causing uplift and thinning of the continental plate by heating, basal shear and melting (Burke and Dewey 1973; Turcotte and Emerman 1983; Buck 1991). The idea is that in the earliest stages of this process, the lithospheric mantle is thinned by thermal erosion, while the continental crust is uplifted and affected by volcanic products, but preserves its initial thickness. The crust is thinned by extension in secondary stages (Merle 2011). This model was developed to explain various observations in rifts such as uplifted plateaus and abundant volcanism (e.g. flood basalts). Depending on the setting, the cause of asthenospheric rise is often related to a mantle plume, although the observations and constraints remain debated. For instance, the Rio Grande Rift (US–Mexico) is often listed as a type example of an active rift, with a narrow geometrical structural setting that is formed by laterally distributed pure shear extension (Wilson et al. 2005), with no observations suggesting the presence of a plume.
On the other end of the spectrum, passive rifting is driven by distant stresses that cause crustal and lithospheric mantle extension and thinning, resulting in a passive asthenospheric mantle rise (Turcotte Oxburgh 1973; Keen 1985; Ruppel 1995). The driving forces are generated by plate boundary motions (e.g. subduction pull and ridge push) or by convective motions at the base of the lithosphere. The main contrast between passive and active rifting is that in passive rifting, the rise of the asthenosphere occurs only in response to the thinning of the overlying lithosphere – as opposed to active upwelling, driving the extension for active rifting. This passive upwelling can then generate various secondary processes affecting the rift-like decompression melting and consequent mantle-derived magmas which are added to the crust (as underplating, intrusions and extrusions) and specific lateral thermal gradients (Ruppel 1995; Huismans et al. 2001).
Based on these definitions, it must theoretically be possible to distinguish the two rift types based on volcanism
