Lithospheric stretching

During horizontal extension, lithospheric stretching results in a vertical thinning of the crust and an increase in the geothermal gradient within the zone of thinning (McKenzie, 1978). These two changes in the physical properties of the extending zone affect lithospheric strength in contrasting ways. Crustal thinning or necking tends to strengthen the lithosphere because weak crustal material is replaced by strong mantle lithosphere as the latter moves upward in order to conserve mass. The upward movement of the mantle also may result in increased heat flow within the rift. This process, called heat advection, results in higher heat flow in the rift because the geotherms become compressed rather than through any addition of heat. The compressed geotherms tend to result in a net weakening of the lithosphere, whose integrated strength is highly sensitive to temperature (Section 2.10). However, the weakening effect of advection is opposed by the diffusion of heat away from the zone of thinning as hot material comes into contact with cooler material. If the rate of heat advection is faster than the rate of thermal diffusion and cooling then isotherms at the base of the crust are compressed, the geotherm beneath the rift valley increases, and the integrated strength of the lithosphere decreases. If thermal diffusion is faster, isotherms and crustal temperatures move toward their pre-rift configuration and lithospheric weakening is inhibited.

England (1983) and Kusznir & Park (1987) showed that the integrated strength of the lithosphere in rifts, and competition between cooling and heat advection mechanisms, is strongly influenced by the rate of extension. Fast strain rates (10-13 s-1 or 10-14 s-1) result in larger increases in geothermal gradients than slow rates (10-16 s-1) for the same amount of stretching. This effect suggests that high strain rates tend to localize strain because inefficient cooling keeps the thinning zone weak, allowing deformation to focus into a narrow zone. By contrast, low strain rates tend to delocalize strain because efficient cooling strengthens the lithosphere and causes the deformation to migrate away from the center of the rift into areas that are more easily deformable. The amount of net lithospheric weakening or strengthening that results from any given amount of stretching also depends on the initial strength of the lithosphere and on the total amount of extension. The total amount of thinning during extension usually is described by the stretching factor (P), which is the ratio of the initial and final thickness of the crust (McKenzie, 1978).

The thermal and mechanical effects of lithospheric stretching at different strain rates are illustrated in Fig. 7.22, which shows the results of two numerical experiments conducted by van Wijk & Cloetingh (2002). In these models, the lithosphere is divided into an upper crust, a lower crust, and a mantle lithosphere that have been assigned different rheological properties (Fig. 7.22a). Figures 7.22b-d show the thermal evolution of the lithosphere for uniform extension at a rate of 16 mm a-1. At this relatively fast rate, heating by thermal advection outpaces thermal diffusion, resulting in increased temperatures below the rift and strain localization in the zone of thinning. As the crust thins, narrow rift basins form and deepen. Changes in stretching factors for the crust (P) and mantle (8) are shown in Fig. 7.22e,f. The total strength of the lithosphere (Fig. 7.22g), obtained by integrating the stress field over the thickness of the lithosphere, gradually decreases with time due to stretching and the strong temperature dependence of the chosen rheologies. Eventually, at very large strains, the thermal anomaly associated with rifting is expected to dissipate. These and many other models of rift evolution that are based on the principles of lithospheric stretching approximate the subsidence patterns measured in some rifts and at some rifted continental margins (van Wijk & Cloetingh, 2002; Kusznir et al., 2004) (Section 7.7.3).

The experiment shown in Fig. 7.22h-j shows the evolution of rift parameters during lithospheric stretching at the relatively slow rate of 6 mm a-1. During the first 30 Ma, deformation localizes in the center of the rift where the lithosphere is initially weakened as isotherms and mantle material move upward. However, in contrast with the model shown in Fig. 7.22b-d, temperatures begin to decrease with time due to the efficiency of conductive cooling at slow strain rates. Mantle upwelling in the zone of initial thinning ceases and the lithosphere cools as temperatures on both sides of the central rift increase. At the same time, the locus of thinning shifts to both sides of the first rift basin, which does not thin further as stretching continues. The mantle thinning factor (Fig. 7.22l) illustrates this behavior. During the first 45 Ma, upwelling mantle causes 8 to be larger in the central rift than its surroundings. After this time, 8 decreases in the central rift as new upwelling zones develop on its sides. The total strength of the lithosphere (Fig. 7.22m) for this low strain rate model shows that the central rift is weakest until about 55 Ma.

Differential stress

Differential stress

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(e) Crustal thinning 25-i

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(e) Crustal thinning 25-i

(f) Mantle lithosphere thinning (8)

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(k) Crustal thinning

(l) Mantle lithosphere thinning (8)

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7.2X10 6.8X10 6.4X10 6.0X10 5.6X10 5.2X10 4.8X10 4.4X10

Figure 7.22 (a) Three-layer lithospheric model where the base of the lithosphere is defined by the 1300°C isotherm at 120km. Differential stress curves show a strong upper crust and upper mantle and a lower crust that weakens with depth. Thermal evolution of the lithosphere (b-d) during stretching for a horizontal extensional velocity of 16 mm a~'. Evolution of lithospheric strength (g) and of thinning factors for the crust (e) and mantle (f) for a velocity of 16 mm a~'. Thermal evolution of the lithosphere (h-j) during stretching for a velocity of 6 mm a~'. Evolution of lithospheric strength (m) and of thinning factors for the crust (k) and mantle (l) for a velocity of 6 mm a- (image provided by J. van Wijk and modified from van Wijk & Cloetingh, 2002, with permission from Elsevier).

7.2X10 6.8X10 6.4X10 6.0X10 5.6X10 5.2X10 4.8X10 4.4X10

T=0"C

Figure 7.22 (a) Three-layer lithospheric model where the base of the lithosphere is defined by the 1300°C isotherm at 120km. Differential stress curves show a strong upper crust and upper mantle and a lower crust that weakens with depth. Thermal evolution of the lithosphere (b-d) during stretching for a horizontal extensional velocity of 16 mm a~'. Evolution of lithospheric strength (g) and of thinning factors for the crust (e) and mantle (f) for a velocity of 16 mm a~'. Thermal evolution of the lithosphere (h-j) during stretching for a velocity of 6 mm a~'. Evolution of lithospheric strength (m) and of thinning factors for the crust (k) and mantle (l) for a velocity of 6 mm a- (image provided by J. van Wijk and modified from van Wijk & Cloetingh, 2002, with permission from Elsevier).

After this time the weakest areas are found on both sides of the central rift basin. This model shows how the strong dependence of lithospheric strength on temperature causes strain delocalization and the formation of wide rifts composed of multiple rift basins at slow strain rates. The model predicts that continental break-up will not occur for sufficiently slow rift velocities.

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