The evolution of rifted margins

The evolution of rifted continental margins is governed by many of the same forces and processes that affect the formation of intracontinental rifts (Section 7.6). Thermal and crustal buoyancy forces, lithospheric flexure, rheological contrasts, and magmatism all may affect margin behavior during continental break-up, although the relative magnitudes and interactions among these factors differ from those of the pre-break-up rifting stage. Two sets of processes that are especially important during the transition from rifting to sea floor spreading include: (i) post-rift subsidence and stretching; and (ii) detachment faulting, mantle exhumation, and ocean crust formation at nonvolcanic margins.

Post-rift subsidence and stretching

As continental rifting progresses to sea floor spreading, the margins of the rift isostatically subside below sea level and eventually become tectonically inactive. This subsidence is governed in part by the mechanical effects of lithospheric stretching (Section 7.6.2) and by a gradual relaxation of the thermal anomaly associated with rifting. Theoretical considerations that incorporate these two effects for the case of uniform stretching predict that subsidence initially will be rapid as the crust is tectonically thinned and eventually slow as the effects of cooling dominate (McKenzie, 1978). However, the amount of subsidence also is influenced by the flexural response of the lithosphere to loads generated by sedimentation and volcanism and by changes in density as magmas intrude and melts crystallize and cool (Section 7.6.7). Subsidence models that include the effects of magmatism and loading predict significant departures from the theoretical thermal subsidence curves.

The amount of subsidence that occurs at rifted margins is related to the magnitude of the stretching factor (P). There are several different ways of estimating the value of this parameter, depending on the scale of observation (Davis & Kusznir, 2004). For the brittle upper crust, the amount of extension typically is derived from summations of the offsets on faults imaged in seismic reflection profiles that are oriented parallel to fault dips. Estimates of the combined upper crustal extension and lower crustal stretching are obtained from variations in crustal thickness measured using wide-angle seismic surveys, gravity studies, and seismic reflection data. This latter approach relies on the assumption that the variations are a consequence of crustal extension and thinning. At the scale of the entire lithosphere, stretching factors are obtained through considerations of the flexural isostatic response to

Present-day seismically observed section 0 Ma

Syn-rift | | = erosion surfaces

100 Ma

Water

150 Ma

Reverse modeled to the base of post-rift sequences

200 Ma

Figure 7.35 Schematic diagram showing application of flexural backstripping and the modeling of post-rift subsidence to predict sequential restorations of stratigraphy and paleobathymetry. Restored sections are dependent on the p stretching factor used to define the magnitude of lithospheric extension and lithospheric flexural strength (after Kusznir et al., 2004, with permission from Blackwell Publishing).

loading (Section 7.6.4) and thermal subsidence. One of the most commonly used approaches to obtaining litho-spheric-scale stretching factors employs a technique known as flexural backstripping.

Flexural backstripping involves reconstructing changes in the depth to basement in an extensional sedimentary basin by taking into account the isostatic effects of loading. The concept behind the method is to exploit the stratigraphic profile of the basin to determine the depth at which basement rock would be in the absence of loads produced by both water and all the overlying layers. This is accomplished by progressively removing, or backstripping, the loads produced by each layer and restoring the basement to its depth at the time each layer was deposited (Fig. 7.35). These results combined with knowledge of water depth theoretically allow determination of the stretching factor (P). Nevertheless, as discussed further below, relationships between stretching factor and subsidence curves may be complicated by interactions between the lithosphere and the sublithospheric mantle. In practice, flex-ural backstripping is carried out by assigning each layer a specific density and elastic thickness (Te) (Section 7.6.4) and then summing the effects of each layer for successive time intervals. Corrections due to sediment compaction, fluctuations in sea level, and estimates of water depth using fossils or other sedimentary indicators are then applied. This approach generally involves using information derived from post-rift sediments rather than syn-rift units because the latter violate assumptions of a closed system during extension (Kusznir et al., 2004). The results usually show that the depth of rifted margins at successive time intervals depends upon both the magnitude of stretching factor (P) and the flexural strength of the lithosphere. Most applications indicate that the elastic thickness of the lithosphere increases as the thermal anomaly associated with rifting decays.

Investigations of lithospheric-scale stretching factors at both volcanic and nonvolcanic margins have revealed several characteristic relationships. Many margins show more subsidence after an initial tectonic phase due to stretching than is predicted by thermal subsidence curves for uniform stretching. Rifted margins off Norway (Roberts et al., 1997), near northwest Australia (Driscoll & Karner, 1998), and in the Goban Spur and Galicia Bank (Davis & Kusznir, 2004) show significantly more subsidence than is predicted by the magnitude of extension indicated by upper crustal faulting. In addition, many margins show that the magnitude of lithospheric stretching increases with depth within ~150 km of the ocean-continent boundary (Kusznir et al., 2004). Farther toward the continent, stretching and thinning estimates for the upper crust, whole crust, and lithosphere converge as the stretching factor (P) decreases. These observations provide important boundary conditions on the processes that control the transition from rifting to sea floor spreading. However, the causes of the extra subsidence and depth-dependent stretching are uncertain. One possibility is that the extra subsidence results from extra uplift during the initial stage of sea floor spreading, perhaps as a result of upwelling anomalously hot asthenosphere (Hopper et al., 2003; Buck, 2004). Alternatively, greater stretching in the mantle lithosphere than in the crust, or within a zone of mantle lithosphere that is narrow than in the crust, also may result in extra uplift. Once these initial effects decay the ensuing thermal subsidence during cooling would be greater than models of uniform stretching would predict. These hypotheses, although seemingly plausible, require further testing.

Observations of the southeast Greenland volcanic margin support the idea that the flow of low-density mantle during the transition to sea floor spreading strongly influences subsidence and stretching patterns. Hopper et al. (2003) found distinctive changes in the morphology of basaltic layers in the crust that indicate significant vertical motions of the ridge system. At the start of spreading, the system was close to sea level for at least 1 Myr when spreading was subaerial. Later subsidence dropped the ridge to shallow water and then deeper water ranging between 900 and 1500 m depth. This history appears to reflect the dynamic support of the ridge system by upwelling of hot mantle material during the initiation of spreading. Exhaustion of this thermal anomaly then led to loss of dynamic support and rapid subsidence of the ridge system over a 2 Ma period. In addition, nearly double the volume of dikes and volcanic material occurred on the Greenland side of the margin compared to the conjugate Hatton Bank margin located south of Iceland on the other side of the North Atlantic ocean. These observations indicate that interactions between hot asthenosphere and the lithosphere continue to influence the tectonic development of rifted margins during the final stages of continental breakup when sea floor spreading centers are established.

The flow of low-density melt-depleted astheno-sphere out from under a rift also may help explain the lack of magmatic activity observed at rifted non-volcanic margins. The absence of large volumes of magma could be linked to the effects of prior melting episodes, convective cooling of hot astheno-sphere, and/or the rate of mantle up welling (Buck, 2004). As sublithospheric mantle wells up beneath a rift it melts and cools. This process could result in shallow mantle convection due to the presence of cool, dense melt-depleted material overlying hotter, less dense mantle. Cooling also restricts further melting by bringing the mantle below its solidus temperature (Section 7.4.2). If some of this previously cooled, melt-depleted asthenosphere is pulled up under the active part of the rift during the transition to sea floor spreading, its presence would suppress further melting, especially if the rate of rifting or sea floor spreading is slow. The slow rates may not allow the deep, undepleted asthenosphere to reach the shallow depths that generate large amounts of melting.

Magma accretion, mantle exhumation, and detachment faulting

The transition from rifting to sea floor spreading at nonvolcanic margins is marked by the exhumation of large sections of upper mantle. Seismic reflection data collected from the Flemish Cap off the Newfoundland margin provide insight into the mechanisms that lead to this exhumation and how they relate to the formation of ocean crust.

The Flemish Cap is an approximately circular shaped block of 30-km-thick continental crust that formed during Mesozoic rifting between Newfoundland and the Galicia Bank margin near Iberia (Fig. 7.36a). The two conjugate margins show a pronounced break-up asymmetry. Seismic images from the Galicia Bank show a transition zone composed of mechanically unroofed continental mantle (Fig. 7.36b) and a strong regional west-dipping S-type reflection (Fig. 7.36b, stages 1 & 2) (Section 7.7.2). The transition zone is several tens of kilometers wide off the Galicia Bank and widens to 130 km to the south off southern Iberia. The S-reflec-tion is interpreted to represent a detachment fault between the lower crust and mantle that underlies a series of fault-bounded blocks. By contrast, the Newfoundland margin lacks a transition zone and shows no evidence of any S-type reflections or detachment faults (Hopper et al., 2004). Instead, this latter margin shows an abrupt boundary between very thin continental crust and a zone of anomalously thin (3 to 4 km thick), highly tectonized oceanic crust (Fig. 7.36b, stages 3, 4, and 5). Seaward of this boundary the oceanic crust thins even further to <1.3 km and exhibits unusual very reflective layering (I).

The five stage model of Hopper et al. (2004) explains these structural differences and the evolution of the conjugate margins. In Fig. 7.36b, the top panel shows a reconstruction of the two margins emphasizing their asymmetry at final break-up when the continental crust was thinned to a thickness of only a few kilometers (stage 1). During break-up, displacement within an extensional detachment fault (labeled S in Fig. 7.36b) unroofed a peridotite ridge (PR) above a zone of weak serpentinized upper mantle. Break-up west of the ridge isolated it on the Galicia Bank margin when, during stage 2, mantle melts reached the surface and sea floor spreading was established. Limited magma-tism produced the thinner than normal (3-4 km), highly tectonized ocean crust. During stage 3, a

Stage 1: Continental breakup and mantle exhumation

Incipient spreading center East ^ —Galicja" / Margin

S" detachment^

Stage 1: Continental breakup and mantle exhumation

S" detachment^

Stage 4: Magmatic event (submarine flood volcanism)

Stage S: Late extension, normal seafloor spreading

Stage S: Late extension, normal seafloor spreading

Figure 7.36 (a) Location of seismic surveys of the Flemish Cap and (b) five stage model of nonvolcanic margins (after Hopper et al., 2004, with permission from the Geological Society of America). MO in (a) is magnetic anomaly. Random-dash pattern, continental crust; v pattern, oceanic crust; light gray shading, serpentinized upper mantle; dark gray shading, unaltered upper mantle; thick lines, strong reflections; dashed lines, inferred crust-mantle boundary; dotted lines, oceanic layers; PR, peridotite ridge; S, reflections interpreted to represent a detachment fault; I, unusual very reflective oceanic crust;Z, reflections interpreted to represent a detachment fault buried by deep marine flood basalts.

Figure 7.36 (a) Location of seismic surveys of the Flemish Cap and (b) five stage model of nonvolcanic margins (after Hopper et al., 2004, with permission from the Geological Society of America). MO in (a) is magnetic anomaly. Random-dash pattern, continental crust; v pattern, oceanic crust; light gray shading, serpentinized upper mantle; dark gray shading, unaltered upper mantle; thick lines, strong reflections; dashed lines, inferred crust-mantle boundary; dotted lines, oceanic layers; PR, peridotite ridge; S, reflections interpreted to represent a detachment fault; I, unusual very reflective oceanic crust;Z, reflections interpreted to represent a detachment fault buried by deep marine flood basalts.

reduction in magma supply led to about 20 km of extension that was accommodated mostly by detachment faulting. The detachment faulting led to the exhumation of the mantle and formed an oceanic core complex that is similar to those found in slow-spreading environments at ridge-transform intersections (Section 6.7). Voluminous but localized magmatism during stage 4 resulted in a 1.5-km-thick layer of deep marine flood basalts that buried the detachment surface (reflection

Z). The intrusion of gabbroic material may have accompanied this volcanism. This magmatic activity marked the beginning of sea floor spreading that formed normal (6 km) thickness ocean crust (stage 5).

This example shows that, to a first order, the transition from rift to oceanic crust at nonvolcanic margins is fundamentally asymmetric and involves a period of magmatic starvation that leads to the exhumation of the mantle. This type of margin may typify slow-spreading systems (Section 6.6) where large fluctuations in melt supply occur in transient magma chambers during the early stages of sea floor spreading.

7.8 CASE STUDIES: THE TRANSITION FROM RIFT TO RIFTED MARGIN

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