Modes of shortening in foreland foldthrust belts

A common characteristic of fold and thrust belts is the presence of one or more décollement (or detachment) surfaces that underlie shortened sequences of sedimentary and volcanic rock (Section 9.7). The geometry of these surfaces tends to conform to the shape of the sedimentary and volcanic sections in which they form. In most foreland basins sedimentary sequences thin toward the foreland, resulting in décollements that dip toward the hinterland (Figs 10.5b, 10.7). In thin-skinned thrust belts (Section 10.2.4, Fig. 10.5b), the lowermost, or basal, décollement separates a laterally displaced sedimentary cover from an underlying basement that is still in its original position. In thick-skinned styles (Fig. 10.5d), the décollement surface cuts down through and involves the crystalline basement.

The development of thin- or thick-skinned styles of shortening commonly is controlled by the presence of inherited stratigraphic and structural heterogeneities in the crust. In the central Andean foreland, for example (Section 10.2.3), variations in the thickness and distribution of sedimentary sequences have been linked to different modes of Neogene shortening (Kley et al., 1999). Thin-skinned styles preferentially occur in regions that have accumulated >3 km of sediment, where the low mechanical strength of the sequences localizes deformation above crystalline basement (Allmendinger & Gubbels, 1996). Thick-skinned styles tend to occur in regions where Mesozoic extensional basins have inverted (Sections 9.10, 10.3.3). As these latter basins experience the shift from extension to contraction, old normal faults involving basement rock reactivate (Turner & Williams, 2004; Saintot et al., 2003; Mora et al., 2006).

In many fold and thrust belts, and especially in thick-skinned varieties, shortening results in some faults that dip in a direction opposite to that of the basal décollement, creating a doubly vergent thrust wedge composed of forward-breaking and back-breaking thrusts. These doubly vergent wedges may occur at any scale, ranging from relatively small basement massifs (Fig. 10.5d) to the scale of an entire collisional orogen (Fig. 8.23b,d). Their bivergent geometry reflects a condition where the material on the upper part of an advancing thrust sheet or plate encounters resistance to continued forward motion (Erickson et al., 2001; Ellis et al., 2004). The resistance may originate from friction along the décol-

lement surface as the wedge thickens or as material moves over a thrust ramp (Section 9.7). It also may result where an advancing thrust sheet encounters a buttress made of a strong material, such as the volcanic arc in an accretionary prism (Section 9.7) or the boundary between a rigid plate and a weaker plate (Section 8.6.3). Buttresses also may result from a change in lithol-ogy across an old normal or thrust fault, from a thickening sequence of sedimentary rock, or from any other mechanism.

The lateral (across-strike) growth of thrust wedges (Section 9.7) and the involvement of deep levels in the deformation are controlled by the temperature and relative strengths of the shallow and deep crust. If the upper crust is strong and the deep crust relatively hot and weak, then shortening may localize into narrow zones and thick-skinned styles of deformation result (Ellis et al., 2004; Babeyko & Sobolev, 2005). A weak middle and lower crust promotes ductile flow and inhibits the lateral growth of the thrust wedge. Deep crustal flow also tends to result in low critical tapers (Section 9.7) and a symmetric crustal structure that includes both forward- and back-breaking thrusts. During basin inversion, more normal faults tend to reactivate if the middle or lower crust is weak relative to the upper crust (Nem&k et al., 2005; Panien et al., 2005). By contrast, if the upper crust is weak and the deep crust is cool and strong, then shortening leads to a mechanical failure of upper crustal sequences and the orogen grows laterally by thin-skinned deformation. In scenarios involving a strong lower crust, thrust wedges tend to show high tapers, asymmetric styles (mostly forward-breaking thrusts), and rapid lateral growth.

A combination of these effects may explain why con-tractional deformation led to the rapid lateral growth of a foreland fold and thrust belt in the Central Andes and not in the Southern Andes (Section 10.2.3). Allmendinger & Gubbels (1996) recognized that deformation in these two regions involved two distinctive modes of shortening. In an older pure shear mode of shortening, deformation of the upper and lower crust occurred simultaneously in the same vertical column of rock. North of 23°S, this type of deformation was focused within the Altiplano. Later, during the Late Miocene the deformation migrated eastward, forming a thin-skinned foreland fold and thrust belt in the sub-Andean ranges while the middle and lower crust of the Altiplano continued to deform. This latter mode of shortening, where deformation in the upper crust and the deep crust are separated laterally, is known as simple shear.

The underthrusting of the Brazilian Shield beneath the Altiplano most likely drove the simple shear (Section 10.2.5). South of 23°S the pure shear mode of shortening lasted longer and was replaced by a thick-skinned thrust belt involving a mix of both pure and simple shear.

These differences in the style and mode of shortening along the strike of the Andes appear to be related to variations in the strength and temperature of the foreland lithosphere. Allmendinger & Gubbels (1996) postulated that the shallower basement and lack of a thick sedimentary cover in the Sierra Santa Bárbara ranges, south of latitude 23°S, precluded thin-skinned deformation and allowed deformation to remain in the thermally softened crust of the Puna for a long period of time. In addition, the mantle lithosphere beneath the Puna is significantly thinner than beneath the Altiplano, suggesting that the crust in the former is hotter and weaker.

To test this idea, Babeyko & Sobolev (2005) conducted a series of thermomechanical experiments where the cold, rigid lithosphere of the Brazilian Shield indented into the warm, soft lithosphere of the adjacent Altiplano-Puna. Figure 10.12a shows the model setup, which includes a thick plateau crust on the left and a three-layer crust on the right above mantle lithosphere. The three layer crust includes an 8-km-thick layer of Paleozoic sediments. The mechanical strength of this layer and the temperature of the foreland are the two main variables in the model. A Mohr-Coulomb elasto-plastic rheology simulates brittle deformation and a temperature- and strain rate-dependent viscoelastic rheology simulates ductile deformation. The whole system is driven by a constant shortening rate of 10 mm a-1 applied to the right side of the model.

After 50 km of shortening, the models show distinctive modes of shortening. In the case where the Paleozoic sediments are strong (or absent) and lie on top of a cold strong Brazilian Shield (Fig. 10.12b), the crust and mantle deform together homogeneously in pure shear mode (Fig. 10.12d). No deformation occurs in the indenting foreland where cold temperatures inhibit lateral growth of a thrust wedge. In the case where the Paleozoic sediments are weak and the foreland cold and

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Plateau

Brazilian shield

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Brazilian shield

Paleozoic sediments

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Felsic crust

Mafic crust

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Initial temperature. Cold foreland

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Initial temperature. Warm foreland

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Initial temperature. Cold foreland

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Initial temperature. Warm foreland

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Figure 10.12 (a-c) Initial setup and (d-f) results of numerical simulations of foreland deformation (after Babeyko & Sobolev, 2005, with permission from the Geological Society of America). (d-f) Accumulated finite strain after 50 km of shortening for three modes. White solid lines are boundaries of lithologic units. See text for explanation.

10 mm a

strong (Fig. 10.12e), the foreland displays a simple shear thin-skinned mode of deformation. Underthrusting of the shield is accompanied by the eastward propagation of the thin-skinned thrust belt above a shallow décollement at 8-14 km depth and drives deformation in the lower crust beneath the plateau. This style conforms well to observations east of the Altiplano and north of 23°S. It also simulates the conditions of the Himalayan fold-thrust belt south of the Tibetan Plateau (Section 10.4.4). In the case where the Paleozoic sediments are weak and the foreland warm and weak (Fig. 10.12c), deformation in the foreland is thick-skinned with a deep décollement at ~25 km depth (Fig. 10.12f). This latter style conforms well to observations east of the Puna and south of 23°S and results because the foreland is weak enough to deform by buckling.

These observations and experiments illustrate that variations in lithospheric strength and rheology play an important role in controlling the tectonic evolution of compressional basins and fold-thrust belts. These effects are prominent at scales ranging from individual thrust sheets to the entire lithosphere.

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