The Andes Mountains, a 6,000-km long rampart of rugged land speckled with several peaks over 6 km high, rim the western edge of South America (Figure 13.1a). Along the range, powerful volcanoes occasionally spew clouds of ash skyward. Halfway around the world, Mt. Everest, the highest mountain in the Himalayan chain (and in the world), rises 8.5 km above sea level (Figure 13.1b). At its peak, air density is so low that climbers use bottled oxygen to stay alive. Why did the immense masses of rock comprising such mountains rise to such elevations? Before the 1960s, geologists really didn’t know. But plate tectonics theory provides a ready explanation—the Andes Mountains formed where the Pacific Ocean floor subducts beneath South America along a convergent plate boundary, while the Himalayan Mountains rose when India rammed into Asia, forming a collisional orogen.

van1701a_color

(a)

(b)

FIGURE 13.1. (a) Photo of the southern Andes Mountains of Chile. Rocks exposed on these peaks include relicts of old accretionary prisms, as well as granitic intrusions of a continental volcanic arc. (b) Himalayas, looking south from over the Tibetan Plateau, including the summits of Makalu (left; 8,462 m) and Everest (right; 8,850 m), which is the highest mountain on Earth. [NASA] [17.1]

Complex suites of structures (involving thrust faults, folds, and tectonic foliations) develop at convergent plate boundaries and collisional orogens. As a consequence, the crust shortens and thickens. In the process, metamorphism and, locally, igneous activity takes place. And, though it may seem surprising at first, gravity can cause the high regions of convergent and collision orogeny to collapse and spread laterally, yielding extensional faulting. In this chapter, we describe both the structural features and the rock assemblages that develop during convergent-margin tectonism and continental collision.

CONVERGENT PLATE MARGINS

When oceanic lithosphere first forms at a mid-ocean ridge, it is warm and relatively buoyant. But as lithosphere moves away from the ridge axis, it cools and the lithospheric mantle thickens, so that when lithosphere has aged more than 10 or 15 million years, it becomes negatively buoyant. In other words, old oceanic lithosphere is denser than underlying hot asthenosphere, and thus can subduct (or, sink) into the asthenosphere, like an anchor sinks through water. Such subduction occurs at a convergent plate boundary (which may also be called a subduction zone or a convergent margin). Here, oceanic lithosphere of the downgoing plate (or downgoing slab) bends and sinks into the mantle beneath the overriding plate (or overriding slab). An overriding plate can include either oceanic crust or continental crust, but a downgoing plate can include only oceanic crust, because continental crust is too buoyant to subduct.

Exactly how the subduction process begins along a given convergent plate boundary remains somewhat unresolved. Possibly it is a response to compression across a preexisting weakness such as may occur at a contact between continental and oceanic lithosphere along a passive continental margin, at a transform fault, or at an inactive mid-ocean ridge segment. Conceivably, compression causes thrusting of the overriding plate over the soon-to-be subducting plate. Once the subducting plate turns down and enters the asthenosphere, it begins to sink on its own because of its negative buoyancy. The subducting plate pulls the rest of the oceanic plate with it and gradually draws this plate into the subduction zone. In other words, because of its negative buoyancy, the subducted plate exerts a slab-pull force on the remaining plate and causes subduction to continue. When the subducting slab reaches a depth of about 150 km, it releases volatiles (H2O and CO2) into the overlying asthenosphere, triggering partial melting of the asthenosphere. The melt rises, some making it to the surface, where it erupts to form a chain of volcanoes called a volcanic arc.

FIGURE 13.2. Plate boundaries of modern Earth. Ocean ridges (heavy lines), trenches (lines with teeth on the overriding plate), and transform faults (thin lines) mark the plate boundaries. [14.14]

Convergent plate boundaries at the surface are marked by trenches (Figure 13.2). Presently they bound much of the Pacific Ocean and the volcanism along these boundaries led geographers to refer to the Pacific Rim as the “ring of fire.” Other present-day convergent plate boundaries define the east edge of the Caribbean Sea, the east edge of the Scotia Sea, the western and southern margin of southeast Asia, and portions of the northern margin of the Mediterranean Sea. In the past, the distribution of convergent plate boundaries on the surface of the Earth was much different. For example, during most of the Mesozoic, the west coast of North America and southern margins of Europe and Asia were convergent plate margins, but convergence at these localities ceased during the Cenozoic. Some convergent plate boundaries mark localities where oceanic lithosphere subducts under oceanic lithosphere (e.g., along the Mariana Islands and Aleutian Islands), and others mark localities where oceanic lithosphere subducts under continental lithosphere (e.g., along the Andes).

FIGURE 13.3. Generalized cross section of a convergent plate margin and related terminology. In this case, the margin occurs along the edge of a continent, bordered by a sliver of trapped oceanic crust. [17.4]

If you make a traverse from the abyssal plain of an ocean basin across a convergent plate boundary, you will find several distinctive tectonic features (Figure 13.3). A deep trough, the trench, marks the actual boundary between the downgoing and overriding plates. An accretionary wedge or accretionary prism, consisting of a package of intensely deformed sediment and oceanic basalt, forms along the edge of the overriding plate adjacent to the trench. Undeformed strata of the forearc basin buries the top of the accretionary prism and, in some cases, trapped seafloor or submerged parts of the volcanic arc. This basin lies between the exposed accretionary prism and the chain of volcanoes that comprises the volcanic arc. For purposes of directional reference, we refer to the portion of a convergent margin region on the trench side of a volcanic arc as the forearc region, while we refer to the portion behind the arc as the backarc region. These and related terms of divergence tectonics are summarized in Supporting Material.

To get an overview of what a convergent plate margin looks like, we take you on a brief tour from the ocean basin across a convergent plate margin. We start on the downgoing slab, cross the trench, climb the accretionary wedge, and trundle across the forearc basin and frontal arc into the volcanic arc itself. We conclude our journey by visiting the backarc region.

The Downgoing Slab

The first hint that oceanic lithosphere is approaching a subduction zone occurs about 250 km outboard of the trench (i.e., in the seaward direction, away from the trench). Here, the surface of the lithosphere rises to form a broad arch called the outer swell or peripheral bulge (Figure 13.4a). The elevation difference between the surface of an abyssal plain of normal depth and the crest of the swell itself, is about 500–800 m. Outer swells form because of the flexural rigidity of the lithosphere. Specifically, downward bending of the lithosphere at a convergent plate boundary levers up the lithosphere outboard of the trench and causes it to rise; you can illustrate this phenomenon by bending a sheet of plastic over the edge of a table (Figure 13.4b). Oceanic crust stretches to accommodate development of an outer swell and, as a result, an array of trench parallel normal faults develops along the crest of the outer swell (Figure 13.4c).

FIGURE 13.4. The peripheral bulge of the downgoing plate. (a) Example of the peripheral bulge just east of the Mariana Trench in the western Pacific. The solid line is a profile of the actual surface of the downgoing slab, while the dashed line is a calculated profile assuming the plate behaves like a sheet with flexural rigidity. (b) Table-top model of a peripheral bulge, with the inset showing lever concept. (c) Stretching along the outer swell produces horsts and grabens at the surface of the downgoing slab. [17.5]

The Trench

Trenches are linear or curvilinear troughs that mark the boundary, at the Earth’s surface, between the downgoing slab and the accretionary prism of the overriding plate (Figure 13.5). They exist because the subducted portion of the downgoing slab pulls the slab downwards to a depth greater than it would be if the lithospheric plate were isostatically compensated. Mass deficit from this depression at trenches produces a negative gravity anomaly, which is a geophysical signature of subduction zones.

FIGURE 13.5. Trenches fill with turbidites. Much of this sediment flows down submarine canyons and then accumulates in turbidite fans on the floor of the trench. [17.10]

The deepest locations in the oceans occur at trenches. In fact, the floor of the Mariana Trench in the western Pacific reaches a depth of over 11 km, deep enough to swallow Mt. Everest (nearly 9 km high) without a trace. But not all trenches are so deep. For example, the Juan de Fuca Trench in the Pacific, off the coast of Oregon and Washington (northwestern USA), is not much deeper than the adjacent abyssal plain of the ocean floor. Trench-floor depth reflects two factors: (1) the age of the downgoing slab (the floor of older oceanic lithosphere is deeper than the floor of younger oceanic lithosphere), and (2) the sediment supply into the trench (if a major river system from a continent spills into a trench, the trench fills with sediment). To see the effect of these parameters, let’s compare the geology of the Mariana Trench and that of the Oregon-Washington Trench. The great depth of the Mariana Trench is a result of its location far from a continental supply of sediment and the fact that the plate being subducted at the Mariana Trench is relatively old (Mesozoic). In contrast, the trench along the Pacific northwest margin of the United States has filled with sediments carried into the Pacific by the Columbia River, and the downgoing slab beneath the trench is quite young (Late Cenozoic).

Whereas the thickness of sediments in trenches is variable, all trenches contain some sediment, which is called trench fill. Typically, trench fill consists of flat-lying turbidites and debris flows that descended into the trench via submarine canyons (Figure 13.5). The sediment comes from the volcanic arc and its basement, from the forearc basin, and from older parts of the accretionary wedge. Eventually, the trench fill becomes incorporated into the accretionary prism, where it becomes deformed.

The Accretionary Wedge

During the process of subduction, the surface of the downgoing plate shears against the edge of the overriding plate. As we have already noted, the shear between the two plates produces an accretionary wedge. This is a tectonic element consisting of deformed pelagic sediment and oceanic basalt, which were scraped off the downgoing plate, and of deformed turbidite that had been deposited in the trench. There are two different accretionary prism geometries, which we illustrate for the case of a convergent margin near a continent. In Figure 13.6a, the prism forms seaward of a trapped sliver of oceanic crust, whereas in 13.6b, the edge of the continent comes directly in contact with the surface of the downgoing plate.

FIGURE 13.6. Two configurations of accretionary wedges (roughly to scale). (a) An accretionary wedge caught beneath the lip of trapped ocean lithosphere. (b) An accretionary wedge being scraped off the edge of a continent. The wedge, in this case, is bivergent. (c) The bulldozer analogy for the formation of an accretionary wedge as in (a). The blade acts as the backstop. (d) Sandbox model showing the formation of a bivergent wedge, as in (b). [17.11]

Traditionally, the process of forming an accretionary prism has been likened to the process of forming a sand pile in front of a bulldozer (Figure 13.6c). The blade of the bulldozer can be called a backstop, in the sense that it is a surface that blocks the movement of material that had been moving with the downgoing plate. Another way to visualize accretionary prism formation process comes from a simple sand-box model. In this model, a sand layer buries a sheet of mylar (thin plastic) that can be pulled through a slit in the base of the box; the slit represents the contact between the overriding and downgoing plates. As the mylar sheet moves down through the slit, the nonmoving sand on the “overriding plate” acts as a backstop, so the sand brought into the subduction zone piles up (Figure 17.11d). Note that in this configuration, a bivergent wedge forms. This means that the prism consists of a forewedge and retrowedge. In the forewedge, the portion of the accretionary prism closer to the trench, structures verge toward the trench (i.e., toward the downgoing plate), while in the retrowedge, the portion of the prism closer to the arc, structures verge toward the arc (i.e., toward the overriding plate). Note that the material of the wedge itself serves as the backstop.

Compressional deformation in the accretionary prism produces thrust faults, folds, and cleavage. But tectonic compressional stress is not the only cause of strain in a prism. Gravity sliding causes slumping of rock and sediment down the slope of the prism toward the trench. And once the prism has become very thick, it begins to undergo extensional collapse under its own weight, like soft cheese. This means that gravitational energy overcomes the strength of material at depth in the internal part of the prism, so this material spreads sideways, leading to horizontal stretching in the above prism. As a consequence of this stretching, the region near the surface of the prism undergoes normal faulting (Figure 13.7).

FIGURE 13.7. (a) Schematic of an accretionary wedge, showing different regimes of deformation, including reverse and normal faulting. [17.12]

The overall consequence of deformation and mass wasting processes in accretionary wedges results in long-term internal “circulation” of material within the wedge. During this process material first moves down to the base of the wedge and then moves back up toward the surface. Thus, sediment that accreted to the base of the wedge by subduction may later end up exposed at the surface of the wedge. This movement reflects both internal thrusting in the wedge that pushes material up, and normal faulting and slumping that strips away the overlying material of the wedge (Figure 13.8). Such net material flow within an accretionary wedge explains how blueschist, formed at the base of the wedge, can eventually be brought to the surface of the wedge. The combination of tectonic deformation, gravity-driven slumping, and extensional collapse that takes place in prisms makes them structurally complex and heterogeneous.

FIGURE 13.8. Bulk flow path of sediment in an evolving accretionary wedge. The large arrows trace the average movement of a grain of material. Note that during its movement, the grain may enter the region of blueschist metamorphism. [17.14]

In some places, accretionary prisms consist of coherent sequences of strata containing parallel arrays of folds and faults and an axial-planar cleavage. Elsewhere, prisms consist of broken formation in which beds can be traced for only a short distance before they terminate at another lithology, or mélange, a chaotic mixture of different rock types (Figure 13.9). In mélange, bedding cannot be traced very far at all, and rocks of radically different lithology and metamorphic grade are juxtaposed. On a traverse across an accretionary prism, you will find exposures of slate and lithic sandstone (derived from trench-fill turbidites), bedded chert (derived from pelagic silicic ooze), micrite (derived from pelagic carbonate ooze), greenstone (altered seafloor basalt), and blueschist (metamorphic rocks containing a blue amphibole called glaucophane).

FIGURE 13.9. Photograph of Paleozoic mélange exposed in an outcrop in north-central Newfoundland (Canada). Note how layers are disrupted, strongly cleaved, and complexly folded, and that a variety of different rock types are visible. [17.15]

The origin of blueschist puzzled geologists for years, because glaucophane can only form under unusual conditions of very high pressure (as occurs at depths >20 km) and relatively low temperature. Because of the geothermal gradient (rate of increase of temperature with depth) characteristic of continental crust, glaucophane does not form in normal continental crust. But once geologists began to understand the process by which accretionary prisms formed, the location of blueschist formation became clear. Blueschists form at the base of the deepest part of the prism, where pressures are great, due to the overburden of 20 km of sediment, but temperatures are relatively low because the underlying oceanic lithosphere is relatively cold.

The Forearc Basin and the Volcanic Arc

As we continue our tour up the slope of the accretionary prism and toward the volcanic arc, we find that the top of the prism is defined by an abrupt decrease in slope. This topographic ridge is the trench-slope break (Figure 13.3). In a few locations around the world (e.g., Barbados, east of the Lesser Antilles volcanic arc along the east edge of the Caribbean), the trench-slope break emerges above sea level. At many convergent margins, a broad shallow basin covers the region between the trench-slope break and the volcanic arc (Figure 13.3). This forearc basin contains flat-lying strata derived by erosion of the arc and the arc’s substrate. Typically, strata of the forearc basin overlie older, subsided, portions of the prism. But locally, these strata overlie ocean lithosphere that had been trapped between the arc axis and the trench when subduction initiated. The strata may also overlie older parts of the volcanic arc and its basement.

The volcanic arc is the chain of volcanoes that forms along the edge of the overriding plate, about 100–150 km above the surface of the subducted oceanic lithosphere. As noted earlier, most of the magma that rises to feed the arc forms by partial melting in the asthenosphere above the surface of the downgoing slab. This partial melting takes place primarily because of the addition of volatiles (H2O or CO2) released from the downgoing plate into the mantle as the downgoing plate heats up. Some researchers argue that small amounts of melt may be derived locally from the downgoing plate.

FIGURE 13.10. Satellite image of the Aleutian volcanic arc off the Alaskan Peninsula, showing the curvature of arcs. The Aleutians are a chain of hundreds of volcanic islands that extend more than 1500km westward from the Alaskan Peninsula. [NASA]

Island arcs form where one oceanic plate subducts beneath another (Figure 13.10), or where the volcanic arc grows on a sliver of continental crust that had rifted from a continent; continental arcs grow where an oceanic plate subducts beneath continental lithosphere. Volcanism at island arcs formed on oceanic crust tends to produce mostly mafic and intermediate igneous rocks, whereas volcanism at continental arcs also produces intermediate and silicic igneous rocks, including massive granitic batholiths. The large volumes of silicic magmas in continental arcs form when hot mafic magmas rising from the mantle transfer heat into the surrounding continental crust and cause melting of the crust. While partial melting of mantle peridotite (an ultramafic rock) yields basaltic (mafic) magma, partial melting of mafic or intermediate continental crust yields intermediate to silicic magma.

The arc-trench gap (Figure 13.11), meaning the distance between the arc axis and the trench axis, varies significantly among convergent margins. Two factors control the width of the arc-trench gap at a given convergent margin:

Dip of downgoing slab. Geometric principles dictate that if the downgoing slab dips very steeply, then the arc-trench gap must be narrow, but if the downgoing slab dips gently, then the arc-trench gap must be broad.
Width of accretionary prism. Where subduction has continued for a long time, or where a large river fills the trench with sediment, the accretionary prism grows to be very large. When this happens, the prism acts as a weight that flexurally depresses the downgoing slab, and as the prism builds seaward, the trench location migrates seaward.

The Backarc Region

The region on the opposite side of the volcanic arc from the forearc basin is called the backarc region. The structural character of backarc regions varies with tectonic setting (Figure 13.11). For the purpose of discussion, we recognize three types of backarc regions: (1) contractional, (2) extensional, and stable backarc regions.

(a) Andean Type - with contractional backarc region

(b) Mariana Type - with extensional backarc region

FIGURE 13.11. Different kinds of volcanic arcs and backarc regions. (a) An Andean-type

continental arc, with a compressional backarc region. Here, the volcanic arc grows on continental crust, and compression has generated a fold-thrust belt and basement-cored uplifts. Large granitic plutons develop. This situation develops where the velocity of the overriding plate (vOR) is in the same direction and exceeds the rollback velocity (vRB). (b) A Mariana-type island arc, with an extensional backarc. Here, the volcanic arc grows on oceanic crust, and a backarc basin develops in which there is seafloor spreading. A remnant arc, composed of a riftedoff fragment of the arc may occur in the basin. This situation develops when vOR < vRB, or is in the opposite direction to vRB. The island arc must move to keep up with the rollback. [17.16]

In a contractional backarc (Figure 13.11a), a backarc basin does not form. Rather, crustal shortening generates a fold-thrust belt and/or a belt of basement-cored uplifts. Both types of deformation developed on the east side of the Andes during the Cenozoic, so a contractional backarc is commonly called an Andean-type backarc. The style of deformation that develops in a contractional backarc depends on the angle of subduction and on the nature of the crustal section in the overriding plate. If subduction angles are moderate to steep and the backarc region contains thick strata of a former passive margin, a fold-thrust belt develops. If, alternatively, subduction angles are shallow, so the subducted plate shears along the base of the overriding plate, stress activates preexisting basement-penetrating faults further to the foreland, and reverse-sense movement on these faults uplifts basement blocks. Overlying strata drape over the block uplifts and form monoclinal folds.

In an extensional backarc, crustal stretching takes place (Figure 13.11b). This stretching produces a backarc basin. If the stretching produces only a continental rift, then continental crust underlies the backarc basin, but if seafloor spreading takes place, then oceanic crust underlies the backarc basin. A clear example of a backarc basin formed by seafloor spreading occurs behind the Mariana Island Arc in the western Pacific Ocean, so extensional backarcs are commonly called Mariana-type backarcs. Backarc basins appear to initiate by rifting along the length of the volcanic arc. When the rift evolves into a new mid-ocean ridge, it splits off a slice of the volcanic arc, and then seafloor spreading separates this now-inactive slice of arc crust from the still active volcanic arc. A slice of arc crust, separated from the active arc by a new segment of seafloor, is called a remnant arc. Large backarc basins like the Philippine Sea contain several remnant arcs. In the Philippine Sea these appear to have been produced by a succession of separate rifting episodes, each of which yielded a short-lived mid-ocean ridge.

Lastly, when no deformation occurs in the backarc region, we call this stable backarcs. Some stable backarcs may have once been contractional or extensional, but then later became stable when plate motions changed. Others are composed of oceanic lithosphere trapped behind the arc when a convergent margin developed far off the coast of the passive margin. The Bering Sea, north of the Aleutian Islands, for example, is underlain by Mesozoic-age trapped ocean floor that was isolated from the rest of the Pacific plate when the Aleutian volcanic arc formed.

Why do we observe such a wide range of kinematic behavior in backarc regions? The answer comes from examining the relative motion between the backarc region and the volcanic arc. As subduction progresses, the location of the bend in the downgoing plate rolls back, away from the backarc (Figure 13.11). The axis of the volcanic arc moves with the rollback. Thus, if the overriding plate is moving in the same direction but at a rate faster than rollback, a contractional backarc develops. If, however, the overriding plate is stationary or moves away from the trench, then rifting and a backarc basin develop. And if the overriding plate moves in the direction of rollback at the same rate as rollback, then the backarc region is stable. If there is a component of lateral motion between the overriding and downgoing plates, then some strike-slip faulting may also occur in the backarc. Backarc regions can evolve with time. For example, the Japan Sea started as an extensional backarc. Japan’s basement consists of continental crust that originally linked to eastern Asia before spreading in the backarc produced the Japan Sea. Presently, however, the character of earthquakes suggest that shortening, accompanied by strike-slip deformation, takes place in the Japan Sea. Because of this motion, we predict that the Japan Sea will close, suturing Japan back to mainland Asia.

Plate Coupling at Convergent Margins

Taking into account the description of convergent margins that we’ve provided earlier in this section, we see that not all convergent margins display the same suite of rocks and structures. Based on the contrasts among various convergent margins worldwide, geologists distinguish between two end-member types.

In a coupled convergent margin, the downgoing plate pushes tightly against the overriding plate, so the plate boundary overall is under compression. As a consequence, large shear stresses develop across the contact, causing efficient offscraping and tectonic underplating, and therefore buildup of a large accretionary wedge. This shear stress also triggers devastating earthquakes, and development of a contractional backarc region. Perhaps because compression squeezes crustal fractures closed, magma rises slowly, and therefore has time to fractionate and/or cause partial melting of adjacent continental crust before intruding at shallow depth or erupting at the surface. Since partial melting of continental crust produces intermediate to felsic magma, and since fractionation removes mafic minerals from a melt, intermediate to felsic igneous rocks predominate at coupled convergent margins.

In an uncoupled convergent margin, the downgoing plate does not push tightly against the overriding plate, so compression across the margin is not great. As a consequence, shear stresses across the plate boundary are relatively small, thrust earthquakes at the boundary have smaller magnitude, and relatively little offscraping and underplating occurs. In uncoupled systems we find extensional backarcs, and cracks in the overriding plate remain somewhat open, so mantle-derived magmas rise directly to the surface before significant fractionation or crustal contamination occurs. Thus, mafic igneous rocks are more common at such convergent margins.

COLLISIONAL TECTONICS

As subduction consumes an oceanic plate, a piece of buoyant lithosphere attached to the downgoing plate may eventually be brought into the convergent boundary. Examples of buoyant lithosphere include large continents, smaller continental fragments, island arcs, oceanic plateaus (broad regions of anomalously thick oceanic crust, formed by hot-spot volcanism), and spreading ridges. Examples of significant buoyant crustal blocks on today’s Earth is shown in Figure 13.12.

FIGURE 13.12. Global map showing blocks of buoyant crust that includes volcanic arcs, oceanic plateaus and continental fragments. Convergence between two major plates would cause terrane accretion and collisional tectonics. [17.19]

Regardless of type, buoyant lithosphere generally cannot be completely subducted, and when it merges with the overriding slab, the boundary becomes a collision zone. When the forces driving collision cease, the relative motion between the colliding blocks ceases, and when this happens, the once separate blocks of lithosphere have merged to become one. The shear surface that marks the boundary between these once-separate plates is a suture Slivers of ophiolites, oceanic crust and mantle thrust over continental crust, crop out locally along sutures. We define these terms, as well as others used in collision tectonics in Supporting Material.

(a)

FIGURE 13.13. (a) Zipper-like collision between two continents. Here, the ocean between the two continents is closing progressively from north to south. (b) Map showing the convergence of two irregularly-shaped continental margins. Promontories and recesses make the west coast of Continent B irregular. Because of the change in trend of the subduction zone along Continent A, frontal convergence (and, eventually, frontal collision) will occur to the north, while oblique convergence (and, eventually, oblique collision) will occur to the south. [17.20]

The types of rocks and structures formed during a particular collisional orogeny depend on numerous variables, including:

• Relative motion between colliding blocks. Frontal collisions yield thrust faults whose movement is perpendicular to the edge of the colliding blocks, while strain in oblique collisions may be partitioned between thrusting and strike-slip faulting. Where blocks collide obliquely, the timing of collision may be diachronous along strike (Figure 13.13a), and the blocks merge together like the two sides of a zipper.

• Shape of colliding pieces. The collision of broad, smooth continental margins yields fairly straight orogens, while the collision of irregular continental margins (with promontories, which are seaward protrusions of the continental margin, and recesses, which are indentations along the continental margin) yields sinuous orogens, in map view (Figure 13.13b). Promontories act as indenters that push into the opposing margin, creating a localized region of high strain. In some cases, slices of crust may move sideways (relative to the frontal collision direction) along strike-slip faults to get out of the way of the colliding blocks. This movement is a type of lateral escape (see below).

• Physical characteristics of colliding pieces. Physical characteristics, such as temperature, thickness, and composition, influence the way in which crustal blocks deform during collision. For example, warmer and, therefore, softer crust of a younger orogenic belt will develop greater strains during collision than will old, cold cratonic crust. During collision, a craton acts as a rigid indenter that pushes into the relatively soft, younger orogenic belt. The collision between India, an old craton, and southern Asia, a weak Phanerozoic orogen, illustrates such behavior. During this collision, much more deformation has happened in the weak southern margin of Asia than in strong India. In fact, a map of the collision zone (see below) shows that India has actually pushed into Asia, so that a transform fault now bounds each side of the Indian subcontinent.

Because so many variables govern the nature of a collisional orogeny, no two orogenies are exactly the same. Nevertheless, we can provide a basic image of the collision process by outlining the various stages in an idealized collision between two continents A and B (below). For reference, we call the portion of the orogen that is on the craton side of the collision the foreland and the internal part of the orogen the internal zone.

Stage 1: Pre-collision and Initial Interaction

FIGURE 13.14. Continent A has a passive-margin basin on its east coast, while Continent B has a convergent margin on its west coast. [17.21]

Let’s begin by setting the stage for the collision between two continents. In this scenario, Continent A moves toward Continent B as the oceanic lithosphere connected to Continent A subducts beneath the margin of Continent B (Figure 13.14). The margin of Continent A is a passive margin, along which a passive-margin sedimentary basin has developed; in contrast, the margin of Continent B is an Andean-type convergent margin along which a volcanic arc has developed. Continent A remains oblivious to the impending collision until the leading edge of the continent begins to bend, prior to being pulled into the subduction system by the downgoing plate. When this happens, flexure causes the surface of the continental margin to rise, so that the continental shelf rises above sea level and undergoes erosion. The margin of Continent A also undergoes stretching, and as a result, normal faults trending parallel to the edge of the margin start to slip.

Stage 2: Abortive Subduction and Suturing

FIGURE 13.15. During the initial stage of collision, the passive margin is uplifted, and an unconformity (locally, with karst) develops. Turbidites derived from Continent B soon bury this unconformity (see inset). Normal faults break up the strata of the passive-margin basin, due to stretching. But soon, thrusts begin to develop, transporting the deeper parts of the basin over the shallower parts. [17.21]

With continued convergence, the surface of Continent A’s continental shelf becomes the floor of the trench (Figure 13.15). When this happens, turbidites derived from the margin of Continent B and its volcanic arc bury the now-eroded surface of the shelf. Thus, a major unconformity defines the contact between the passive-margin basin sedimentary sequence and the turbidites.

Prior to collision, formation of the accretionary prism progressed along the margin of Continent B as new thrusts cut seaward into the strata on the downgoing oceanic plate. But during collision, the strata on the downgoing plate consist of thick, well-stratified sedimentary beds of the former passive-margin basin. Thus, thrusts propagate into these beds, producing a fold-thrust belt that, with time, grows toward the foreland (i.e., in the direction of the continental interior) of Continent A. The stack of thrust slices acts as a load that depresses the surface of Continent A, yielding a foreland sedimentary basin that spreads out over the edge of Continent A’s craton. Such basins are asymmetric; they are thickest along the margin of the orogen and become thinner toward the interior of the continent. Meanwhile the backarc fold-thrust belt on Continent B continues to be active.

Shortening during collision also reactivates the normal faults that bound basement slices at the base of the passive margin; because reactivation occurs in response to compression, these faults now move as reverse faults. The overall process of transforming the passive-margin basin into a thrust belt is called basin inversion. We use the term to emphasize that, during this process, a region that had previously undergone extension and subsidence during basin formation, now telescopes back together by a reversal of shear sense on preexisting faults, and undergoes uplift.

Eventually, a slice of the oceanic lithosphere that had once separated Continent A from Continent B may thrust over the inverted passive-margin of Continent A. This slice, which appears in the orogen as a band of highly sheared mafic and ultramafic rock, defines the suture; rock on one side of the suture was once part of Continent A, while rock on the other side was once part of Continent B. Meanwhile, in the internal part of the orogen the crust thickens considerably, and folding (creating large, tight to isoclinal folds), mylonitization and regional metamorphism (creating schists and gneisses) occur at depth. With progressive deformation, the plastically deformed metamorphic rocks move upwards and toward the foreland. In some cases large recumbent folds develop. In the European literature, large sheets of such transported rock, locally containing recumbent folds, are called nappes (Figure 13.16).

FIGURE 13.16. (a) Generalized cross section through the European Alps. Here, we do not show today’s erosional land surface, but add features not shown in the collision zones sequence. A mid-crustal weak zone serves as a basal detachment for faulting in the upper crust. Large recumbent folds (nappes) develop in the metamorphic interior of the orogen. (b) Photograph of a major thrust nappe (the Glarner Thrust) in the Swiss Alps that places Permian volcaniclastics over Tertiary shales. [17.23]

Metamorphic rock of the internal zone eventually becomes exposed in the peaks of the mountain range due to exhumation, the combination of processes that strips off rock at the surface of Earth to expose rock that had been deeper. Eventually, the downgoing oceanic lithosphere breaks off the edge of Continent A and sinks slowly into the mantle. Without a source of new magma, the convergent-margin volcanic arc of Continent B shuts off. On Continent B, deformation styles are the same, but the vergence of structures is opposite to those that form on the edge of Continent A; rocks on the Continent B side of the orogen thrust toward the interior of Continent B. Thus, taken as a whole, the orogen is bivergent, meaning that opposite sides of the orogen, overall, verge in opposite directions.

Stage 3: Crustal Thickening and Extensional Collapse

FIGURE 13.17. In a mature collision orogen, the subducting slab has broken off, a suture has formed, and metamorphic rocks are uplifted and exhumed in the interior of the orogen. [17.21]

The last stage in collision tectonics represents final convergence between buoyant continental blocks and deformation of collision zone elements and the leading edges of the colliding blocks (Figure 13.17). So far, we mostly focused on the horizontal shortening that takes place in the crust during collision tectonics. But keep in mind that, as crust shortens horizontally, it also thickens. In fact, the crust beneath collisional orogens may attain a thickness of 60–80 km, almost twice the thickness of normal crust. Shortening during collision also causes the lithospheric mantle to thicken substantially (Figure 13.18a).

FIGURE 13.18. Lithospheric thickening and orogenic collapse. (a) During collision, the crust thickens by thrusting. (b) Later, as collapse occurs, extensional faults develop in the upper crust, while plastic flow occurs at depth. This process may contribute to development of a broad plateau. (c) The soft-cheese analogy for extensional collapse. A block of cold cheese can maintain its thickness. Once the cheese warms up, it weakens and spreads laterally. The rind of the cheese ruptures, and small faults develop. [17.24]

Thickening of the crust cannot continue indefinitely because, as the crust thickens, rock at depth becomes warm and, therefore, weaker. As a consequence, the differential stress developed in the orogen due to the weight of overlying rock (the “overburden”) exceeds the yield strength of the rock at depth, and the rock begins to flow and develop horizontal extensional strain (Figure 13.18b). In other words, because of the force of gravity, very thick orogens collapse under their own weight. As we pointed out earlier, you can picture this process by imagining a block of cheese that is heated in the sun (Figure 13.18c). Eventually, the cheese gets so soft that it spreads out, and the thickness of the block diminishes. This process is called extensional collapse (or, orogenic collapse). Extensional collapse at depth in an orogen causes stretching of the upper crust, where rock is cooler and still brittle. Therefore, during collapse, rock of the upper crust ruptures and normal faults form. Because collapse decreases the thickness of the uppermost crust, it causes decompression of the lower crust. This decompression may trigger partial melting of the deep crust, or even the underlying asthenosphere, producing magmas.

FIGURE 13.19. Post-orogenic plutonism and lithospheric delamination. (a) Thickening forms a keel-shaped mass of cool lithosphere to protrude down into the asthenosphere. (b) The keel drops off by mantle delamination or "mantle drips", and is replaced by warm asthenosphere, causing partial melting and formation of anorogenic (or, post-orogenic) plutons. The surface of the crust may rise as a consequence. [17.25]

Melting may also be caused by lithosphere delamination, which means that a deep keel of lithosphere that develops during thickening drops off (Figure 13.19). Warm asthenosphere rises to take its place and heats the remaining lithosphere. Some researchers argue that a broad portion of the mid-crust may actually become partially molten at this stage. What is certain is that magma does form and intrude the upper crust in many collisional orogens after deformation has ceased. Because it intrudes after deformation, the granite has no tectonite fabric, so geologists refer to the granite as post-orogenic granite. The process of extensional collapse can occur while shortening and thrusting continue along the margins of collisional orogen, and it may continue after shortening has ceased. Extensional collapse, together with erosion, keeps mountain ranges from exceeding elevations of about 8 to 9 km, and contributes to the development of broad, high plateaus like the Tibetan Plateau of Asia.

COLLISIONAL OROGENY

Throughout we emphasized the central role of contraction in convergence and collision. We introduced several geometries and structures, including accretionary wedges, nappe structures and fold-thrust belts. In this section we will look at some examples of fold-thrust belts and the kinematics associated with contractional structures at convergent plate boundaries.

FIGURE 13.20. Folds in Cretaceous shales below a thrust that moved Paleozoic carbonates westward in the frontal segment of the foreland fold-thrust belt of the Canadian Rockies (near Jasper, Alberta).

Fold-Thrust Belts

Picture yourself hiking among the massive cliffs of the Canadian Rockies (Figure 13.20). That alone is a good reason to be a geologist. Where did the beds of sedimentary rock forming these mountains come from, and how did they end up exposed on cliffs 2 km above sea level? The sediment composing the beds originally accumulated on the floor of a sea, tens of kilometers to the west of their present location. Eventually, the sediment was deeply buried until it lay several kilometers below the Earth’s surface. Thus, these strata had to move large distances both horizontally and vertically to get to their present location in the Canadian Rockies. After uplift, erosion by rivers and glaciers carved the rugged cliffs on which the strata crop out today. Formation of the Canadian Rockies involved displacement on multiple thrust faults. The forces driving thrust sheet movement was generated by convergence and/or microplate collision along North America’s western border during Mesozoic and Early Cenozoic orogenies. As a consequence of thrusting, once-horizontal beds of sediment became tilted and folded. Geologic domains, such as the Canadian Rockies, in which regional horizontal tectonic shortening of the upper-crust yields a distinctive suite of thrust faults, folds, and associated mesoscopic structures, are called fold-thrust belts.

Settings of Fold-Thrust Belts

Fold-thrust belts occur worldwide in a variety of tectonic settings—basically, anywhere that a layer of the upper crust undergoes significant horizontal shortening under low-grade or sub-metamorphic conditions. To describe these settings, we first need to introduce a few terms. When specifying relative locations in fold-thrust belts that have formed on continental crust, we use the undeformed region of a continent outside of the fold-thrust belt as a point of reference. The foreland direction is toward the undeformed continental interior, whereas the internal zone is the orogen’s more intensely deformed and metamorphosed region (Figure 13.21). Similarly, when referring to accretionary prisms formed on oceanic lithosphere, geologists use the term foreland to refer to the less deformed side closer to the trench. Portions of some fold-thrust belts on continents involve strata that were deposited in a foreland basin. In accordance with our definition of foreland, a foreland basin is a wedge of sediment deposited on the surface of the continent in the foreland of the orogen. Foreland basins form because the stack of thrust sheets in a fold-thrust belt acts as a weight that bends down the surface of the continent and creates a depression that collects sediment eroded from the orogen.

FIGURE 13.21. (a) Schematic cross section illustrating the location of a fold-thrust belt in an orogen. Note that the belt occurs between the foreland basin and the internal metamorphic zone of the orogen. Thrusts eventually cut across strata of the foreland basin and incorporate the basin material into the fold-thrust belt. [18.2]

Once subduction consumed the oceanic lithosphere between two continents, they collide. Strata that had originally accumulated in a passive-margin basin2 on the downgoing plate get caught in a vise between the two continents and undergo tectonic shortening (Figure 13.22). As a consequence, a fold-thrust belt evolves in which strata of the former passive margin undergo thrusting toward the foreland. During this process, strata of the deeper-water portion of the passive-margin basin may be placed on top of strata of the shallower water part of the basin. In the internal portions of such fold-thrust belts, normal faults form during the rifting that originally created the passive margin reactivate as thrust faults (i.e., inversion), so that basement thrust slices move up and over sedimentary strata. Erosion attacks the rising collisional orogen, providing sediment that collects in a foreland basin. Foreland-basin sediment eventually becomes incorporated in the fold-thrust belt as well. The Valley and Ridge Province of the Appalachians, the Jura Mountains of Switzerland, and the Himalayan Mountains of Asia are examples of fold-thrust belts in collisional orogens.

FIGURE 13.22. Generalized cross section showing fold-thrust belts in a continent-continent collision setting. A fold-thrust belt forms in the foreland of the orogen on both sides of the orogen. Slivers of obducted ocean crust may separate lower-plate rocks from the metamorphic internal zone of the orogen and define the suture between the two plates. [18.3]

Generalized Cross-section of a Fold-Thrust Belt

We’ve considered the regional setting in which fold-thrust belts occur, now let’s focus on the individual structures, and arrays of structures, that occur within them. To begin our discussion, we examine a cross section of the Pine Mountain Thrust in the southern Appalachians (Figure 13.23). This fault formed in a sequence of Paleozoic strata during the Alleghanian Orogeny, when Africa collided with North America, forming Pangaea.

FIGURE 13.23. Pine Mountain Thrust of the Appalachians. (a) Location of the range in the eastern United States. (b) Simplified map showing the trace of the Pine Mountain Thrust. Note the transform (or, tear) faults at its terminations. Location of red box in (a). (c) Cross section of Pine Mountain thrust sheet, southern Appalachians (simplified to illustrate terminology for describing hanging-wall and footwall structures). Location of the cross section shown by the XX’ section line (dashed red) on map in (b). Barbs point toward hanging wall of the thrust faults. [18.7]

As shown in Figure 13.23c, thrust faults like the Pine Mountain Thrust cut up section in the direction that the hanging-wall moves. Displacement on the fault puts older strata on top of younger strata. In the case of the Pine Mountain Thrust, we see that the fault lies at the base of the Cambrian strata in the southeast, cuts up section to the northwest, and eventually flattens out in Siluro-Devonian strata. Movement on the Pine Mountain Fault placed Cambrian strata over Siluro- Devonian strata. Note that thrusting can duplicate a stratigraphic succession, so that a vertical hole drilled through the hanging wall, across the fault, and into the footwall could encounter the same stratigraphic units twice. Further, thrusting raises strata above its pre-faulting elevation. Strata in the hanging wall of the Pine Mountain Thrust lie approximately 2.5 km above their original pre-faulting elevation! Our example of the Pine Mountain Thrust also illustrates that some thrust faults resemble a flight of stairs, in that they consist of flats that lie approximately in the plane of bedding and ramps that cut across bedding. The key to determining whether a fault segment is a ramp or a flat is to look for cutoffs. A cutoff is the intersection between a bedding plane and a fault surface along a ramp. Flats commonly exceed ramps in cross-sectional length, and typically lie within incompetent (weak) strata like shale and evaporite. Ramps tend to develop in competent (strong) rocks like sandstone, dolomite, and limestone. Note that their placement depends on mechanical stratigraphy.

Kinematics of Thrust Systems

A thrust system refers to the family of related thrust faults that ramp up from a single detachment fault or décollement. The name “detachment” emphasizes that rock above the fault has detached or separated from rock below during movement. Detachments tend to develop in weak rock types, such as shale or evaporite. There are two end-member geometric types of thrust systems that we’ll discuss:

Imbricate fans. They are characteristic of accretionary wedges that form in the subduction complex (Figure 13.24). Displacement on each thrust is modest.
Duplexes. They are characteristically found in fold-thrust belts, especially in the foreland areas of final collision, hence their common name as forland fold-thrust belts (Figure 13.25). Thrusts may accumulate considerable displacement.

FIGURE 13.24. An idealized imbricate fan that develops by progressive break-forward thrusting. Note that successively younger thrusts cut into the footwall, and older faults and folds become deformed by younger structures. The dashed lines are the traces of fold axial surfaces. In the cross section showing

“Time 3,” the sequence of thrusts is labeled. Fault 1 is the oldest and Fault 3 is the youngest. On this cross section, tip lines and branch lines are points; in three dimensions, they go into and out of the page. Active fault segments in red. [18.12]

Individual thrusts that make up an imbricate fan branch upsection from a common detachment and terminate updip without merging into an upper detachment (Figure 13.24). The line (in three dimensions) along which the ramp connects to an underlying detachment is called a branch line, and the line at which the fault terminates and displacement decreases to zero is called the tip line.

FIGURE 13.25. Idealized flat-roofed duplex that develops by progressive break-forward faulting. Note that the roof thrust undergoes a sequence of folding and unfolding, and that formation of the duplex results in significant shortening. Active fault segments in red. [Animation] [18.13]

In a duplex, a series of thrusts branches upwards from a lower detachment and merges with a higher detachment (Figure 13.25). Sometimes geologists refer to the lower detachment of a duplex as the floor thrust, and the upper one as the roof thrust (Figure 18.13a and b). Note that adjacent thrust fault surfaces in a duplex completely surround bodies of rock; these fault-bounded bodies are called, somewhat oddly, horses.

In both imbricate fans and duplexes, the faults comprising a thrust system do not all initiate at the same time. Generally they initiate in a break-forward sequence (or “piggy-back”). This means that the faults of the system form one after the other, with each new fault forming to the foreland side of the previous one. Thus, the youngest fault in the system occurs on the foreland end of the system, whereas the oldest fault occurs at the internal end. There are exceptions to the break-forward sequence scenario. Slip at any given time may be partitioned among the youngest thrust and thrusts immediately behind it. Also, in some cases, existing thrusts of the system reactivate and/or new faults initiate to the internal part of the preexisting faults. These out-of-sequence faults can be recognized where they cross cut structures in the foreland. Within a given thrust system, most thrust sheets move in the same overall direction, called the regional transport direction. In the case of convergent-margin fold-thrust belts, the regional transport direction carries rocks from the internal zone of an orogen toward the foreland.

Insights from Modeling

In recent years, researchers are increasingly exploring the processes of convergence and collisional tectonics by means of laboratory analog models, such as sandbox experiment, and by computer models (Figure 13.26). This work allows researchers to simulate the evolution of orogens in front of their eyes and in real-time, and to determine how key variables affect this evolution. Variables that can be tested by modeling include detachment strength, subduction angle, rock strength, surface erosion rates, convergence geometry, among others. The effect of these variables on the evolution of an orogen can be understood in the context of critical-taper theory that we discussed in Chapter 5, as applied to bivergent orogens. With this in mind, we briefly consider the effect of changing some variables in an evolvingorogenic system.

FIGURE 13.26. Results from a computer modeling study showing the flow path of material points in an orogen under different climatic conditions. (a) Wind comes from east, so exhumed metamorphic rocks occur in a narrow belt on the east side. (b) If the wind comes from the west, a broad metamorphic belt occurs on the west. [17.30]

If there is an increase in the erosion rate (due to a change in climate), material transfers from the internal zone to the foreland, which decreasing the taper. Recall that this causes the orogen to grow wider and causes exhumation (erosion and removal of crust at the Earth’s surface, thereby allowing the uplift of crust that had been at depth) in the interior of the orogen. The geometry of erosion depends on the prevailing wind direction, for this determines whether erosion affects the retrowedge or the forewedge (Figure 13.26). The location of erosion, in turn, determines where exhumation takes place and thus where metamorphic rocks formed at depth eventually become exposed. As another example, if the foreland region contains a particularly weak detachment horizon, then the critical taper angle decreases, and orogen grows wider without becoming as thick. Again, this will affect metamorphic patterns in the interior of the orogen. If the geothermal gradient increases, crust at depth becomes weaker, which in turn decreases the taper angle. This decrease in crustal strength allows extensional collapse to take place so that the wedge achieves the lower taper angle. The modern literature in structural geology presents ever more sophisticated modeling studies that focus on coupling between parameters, such as climate and geometry, and exhumation and geometry.

RELATED PROCESSES OF COLLISIONAL TECTONICS

In our description of the stages in an idealized continent–continent collision, we focused only on tectonic phenomena occurring adjacent to the colliding margins, and mostly on movements that can be illustrated by a two-dimensional, cross-sectional plane. Here, we describe additional tectonic processes that can occur during collision, and we consider movements in the third dimension.

Extrusion Tectonics

Collisional orogeny along a continental margin may lead to the development of regional-scale strike-slip faults that propagate far into the interior of the overriding plate. As an example, several large strike-slip faults start at the Himalayas and cut eastward, following curved trajectories across Asia to the Pacific margin (Figure 13.27). It has been proposed that movement on these faults accommodates lateral translation of large wedges of Asia relatively eastward, in response to the collision of India with Asia.

FIGURE 13.27. Extrusion tectonics. (a) Sketch map of major structures in southeastern Asia. Note the major faults that slice across China. The large arrows indicate the motion of large crustal blocks. (b) Map-view sketch of a laboratory experiment to simulate such lateral escape tectonics. A wooden block (representing the Indian craton) is pushed northwards into a clay cake. The cake is restrained on the west, but not on the east. As the block indents, strike-slip faults develop in the clay, and large slices are squeezed eastwards. As we see in the chapter on wrench tectonics, these faults are classified as transfer faults.

[17.26]

To simulate the development of these faults, researchers constructed models in which they push a wooden block into a cake of clay (Figure 13.27). The wooden block represents the rigid craton of India, while the clay represents the relatively soft crust of Asia. The crust of Asia is not a craton; it is relatively soft because it formed during protracted Phanerozoic orogeny. In this model, a rigid wall, representing the mass of western Asia, constrains the left side of the clay cake. The right side, representing the Pacific margin, remains unconstrained. As the wooden block moves into the clay, wedges of clay, bounded by strike-slip faults, move laterally to the right. This scenario is, therefore, also known as lateral escape.

Substantial debate continues about the significance and age of faulting in southern Asia, and the underlying mechanics, but it does illustrate how blocks of the overriding plate in a collisional orogen can be squeezed sideways out of the path of an indenter, much like a watermelon seed squirts sideways when you squeeze it between your fingers. When extrusion tectonics occurs, strain resulting from collision cannot be depicted in a single cross section, because of movement into or out of the plane of a 2D section.

Plateau Uplift

Continental collisions cause the uplift of a linear belt of mountains, a collisional range, but in some cases also leads to the development of a broad region of uplifted land, known as a plateau. As a case in point, we return to the India-Asia collisions and the development of the Tibetan Plateau in southern Asia (Figure 13.27). Determining the time of uplift of this plateau remains a challenging issue, but available evidence suggests that it rose up in concert with India’s collision with Asia.

Researchers continue to debate the reason for the uplift of the Tibetan Plateau. Some consider it to be a consequence of thickening of the crust, either by plastic flow in the lower crust, or because the crust of India has been emplaced under the crust of Asia. Others suggest that it reflects heating that occurred when the lower part of the lithospheric mantle detached (delaminated) from the base of the lithosphere and sank down into the mantle. Such lithosphere delamination would juxtapose hot asthenosphere at the base of the remaining lithosphere and cause the lithosphere to heat up. Thus, to maintain isostatic equilibrium, the surface of the lithosphere would have to rise, resulting in a plateau.

Accretionary Tectonics

If we look at the present-day Pacific Ocean region, we find many pieces of crust that differ in thickness and/or composition from the typical oceanic crust produced at a mid-ocean ridge. These pieces include:

• Small fragments of continental crust, such as Japan or Borneo, which rifted off larger continents in the past.

• Volcanic island arcs, such as the Mariana Arc, which formed along convergent plate boundaries.

• Seamount chains and oceanic island chains that formed above hot spots, such as the Hawaii-Emperor chain.

• Oceanic plateaus, broad regions of anomalously thick crust, probably composed of basalt extruded at particularly productive hot spots, such as Iceland.

All of these pieces are buoyant, relative to the asthenosphere, and thus cannot be subducted. Therefore, if subduction continues along the eastern margin of Asia for many more millions of years, the pieces would eventually collide with and suture to Asia. After each such docking event, a new convergent margin may form on the outboard (oceanic) side of each sutured piece. As a consequence, the continent grows. This overall process is called crustal accretion or accretionary tectonics. The small crustal pieces that have been attached to a larger continental block by accretion are called accreted (or exotic) terranes. In some cases, the area of a continent grows by the addition of a broad belt of accreted terranes called a tectonic collage (Figure 13.28).

FIGURE 13.28. Accreted terranes and tectonic collage. (a) Schematic cross section of an orogen, based loosely on the southern Appalachians, showing how the orogen includes an accreted volcanic arc, an accreted microcontinent, and crust that had been part of a large colliding continent. [17.29]

How do we identify accreted terranes in an orogen? The first hint that a block of crust may be an accreted terrane comes from studying the geologic history preserved in a block. Do the rocks and structures of the block correlate with those of adjacent crust? If not, then the block is probably accreted. Geologists can test this proposal further by using paleomagnetism, which may demonstrate that the accreted block and the continent to which it is now attached did not have the same apparent polar-wander paths prior to the time at which accretion occurred. Paleontology may also indicate that a block originated at a different latitude from the host continent.

Based on mapping, paleomagnetic and paleontologic study, geologists have demonstrated that, during the Mesozoic and Cenozoic, the North American Cordillera grew westward by crustal accretion. In fact, much of the vast tract of land that now comprises most of California, Oregon, Washington and Alaska in the United States, and British Columbia and the Northwest Territories in Canada originated as crustal fragments outboard of North America that were accreted to its continental margin. Much of this accretion occurred during oblique convergence, so strike-slip faults developed in the orogen and movement on them transported whole terranes, or slivers of them, along strike to the north. The Appalachian Mountains of eastern North America preserve a similar, but older story of accretion. The eastern edge of what was North America in the Precambrian lies well inboard of the continent’s present coastline.

When two major continents collide along a margin that previously was the locus of terrane accretion, you can imagine that the resulting collisional orogen will be very complex. It will contain several sutures separating different blocks, and each block has its own unique geologic history. Most major collisional orogens involve accretion of exotic terranes prior to collision of the larger continents and final closure of the intervening ocean, so such complexity is the norm rather than the exception. Therefore, you should not assume that the history in one particular region represents that of the whole orogen, nor should you be surprised to find radically different geologic histories preserved in adjacent areas of crust in the orogen.

OROGENIC ARCHITECTURE

Orogenic architecture describes the broad geometry of a mountain belt (Figure 13.29). Whereas the details of each individual mountain belt differ, they have many features in common.

FIGURE 13.29. Idealized section through a collisional orogen, showing basement blocks, foreland and fold-thrust belts, metamorphic internal zone (with nappes), inverted passive margin, strike-slip plate boundary, accreted volcanic arc, accreted microcontinent, and sutures (S). Most if not all mountain belts contain several of the features shown in this ideal section, but none probably contains all of them. The diagram is based on observation in Phanerozoic mountain belts; most Precambrian belts preserve only the deeper basement levels. [20.3]

Generally we find a deformed, originally wedge-shaped sedimentary sequence that was deposited at the stable continental margin. This sequence may contain marine carbonates if the area was located in the equatorial realm. Slivers of ophiolite, a rock assemblage containing ultramafic (mantle) rocks, gabbros, dikes, and pillow basalts, are remnants of ancient ocean floor that are also preserved in an orogen. It is noteworthy that, based on geochemical evidence, most ophiolites in ancient mountain belts are backarc basin oceanic lithosphere rather than representing main ocean basin rocks. Ophiolites are key evidence for the activity of modern-day plate tectonics in ancient mountain belts. Granites, associated with volcanic arc formation or the melting of overthickened crust, are variably present. As the orogen evolves, marine clastics (also called flysch, after such deposits in the European Alps) that are derived from the eroding mountain belt are deposited in foreland basins and at the waning stages of orogenic activity coarse continental clastics (also called molasse, also named after European Alps geology) are laid down. In young orogenic belts we find that isolated slivers of underlying basement rocks (or, crystalline basement) have become exposed by faulting. In some cases, mantle rocks are similarly exposed. In ancient orogens this basement component significantly increases, and the sedimentary sequence is mainly preserved in metamorphosed and highly deformed rocks. The oldest mountain belts consist nearly entirely of deformed mid-crustal to lower-crustal rocks of magmatic origin. In a way, these ancient orogens expose the roots of deeply-eroded mountain belts and, in combination with modern belts, they provide a fairly complete section through orogenic crust.

Orogenic deformation is usually polyphase (denoted by “D”) and each phase can contain several fold generations (denoted by “F”), as we learned earlier. Within a single orogenic phase, the deformation sequence may look something like the following: Early structures are thrusts that repeat stratigraphy, or large recumbent folds that repeat and locally invert stratigraphy (called nappes). These thrusts often root in a detachment zone (or décollement) at depth. In metamorphic regions this stage has produced widespread transposition. These early structures are overprinted by upright folds that may contain an axial plane foliation, and later fold generations are commonly present as kinks and crenulations. These contractional features locally overprint evidence of an initial rifting stage (normal faulting) that formed at the passive plate margin. In addition to early rifting, extensional structures often form during the later stages of mountain building (called syn-orogenic extension) and during unroofing (called post-orogenic extension.

Curved Orogens (Oroclines)

Looking at a global map of mountain belts around the world, they are commonly curved in map view. Modern examples include the Alps and Cantabrian chains of Europe, the Rockies in North America and Mexico, and the Tasmanides in Australia. Curved mountain belts are often called oroclines, and can be the result of multiple scenarios (Figure 13.30):

the belt interacts with basement highs in the foreland, for the basement highs retard propagation of thrusts;
the strength of the detachment horizon beneath changes along strike, because thrusts propagate further to the foreland where the detachment is weaker;
there are lateral variations in the thickness of the stratigraphic sequence being deformed, for thrust belts are wider where they involve thicker stratigraphic successions;
preexisting thrusts are bent by interaction with a strike-slip fault that cuts across the belt;
a second phase of thrusting overprints a preexisting belt at a high angle to the original belt;
a fold-thrust belt impinges on the corner of a rigid craton.

FIGURE 13.30. Map-view sketches illustrating processes in the formation of curved orogenic belts. (a) Interaction with basement highs in the foreland. New thrusts forming between the two highs originate with curved traces. (b) Lateral pinch-out of a stratigraphic glide horizon. The thrusts propagate further to the foreland over a weak salt horizon and thrusts originate with curved traces. (c) Lateral variations in stratigraphic thickness. Thrusts propagate into the foreland where pre-depositional strata are thicker, so they originate with curved traces. (d) Interaction with a strike-slip fault. Motion on the strike-slip fault bends the thrust traces. (e) Overprinting of two non-parallel belts. Development of the younger belt bends the traces of the older belt. (f) Impingement against an irregular cratonic margin. Thrust sheets bend as they wrap around the cratonic corner. [18.27]

In some cases, the map-view curvature of fold-thrust belts develops when the thrusts initiate, so that right from the start the belt has a curved trace. But in other cases, a preexisting straight belt undergoes bending in map view during a subsequent phase of deformation. Map-view curve that forms by the bending of a preexisting straight belt, such that the arms of the curve rotate around a vertical axis, is called an orocline.

Speculation on Orogenic Styles

Most geoscientists accept the notion that earlier in Earth’s history the mantle was hotter overall than it is today, because the young Earth held more of its primordial heat and had a greater concentration of radioactive elements than it does today. For example, decay of radioactive elements produced three times as much heat at the beginning of the Archean and about 1.8 times as much heat at the beginning of the Proterozoic as it does today. Thus, heat flow and mantle convection in the younger Earth were probably more vigorous than today, but whether a hotter, more vigorously convecting mantle caused young continents to be warmer, and thus weaker, than those of today remains a point of debate. If excess heat of the Earth’s interior was lost, in part, by conduction through the continents, then the continents must have been warmer. However, if the additional heat of the young Earth was lost through convection involving oceanic lithosphere (because spreading rates were faster, or spreading occurred at a greater number of ridges, or there were a greater number of hot spots), then the continental crust may not have been substantially hotter.

Researchers who argue in favor of the idea that the crust was not substantially hotter in Precambrian times, point out that young continents lay above a thick lithospheric root and thus would have been insulated from the convecting asthenosphere. Uncertainty over the stability of the mantle beneath continents complicates interpretation of the thermal conditions of continents. The deep root of thickened, cooler mantle that formed beneath collisional orogens may delaminate, at which time hot asthenosphere flows against the base of the continent, causing an increase in heat flow into the continent. Delamination would be more likely in the Archean if mantle convection was more vigorous, which may explain the prevalence of high-temperature metamorphism in Archean terranes.

The height of a mountain range on Earth depends largely on the strength of the crust, because it collapses and spreads laterally under its own weight if the gravitational load of the elevated region exceeds the strength of rock at depth, and exceeds the magnitude of the horizontal tectonic forces that hold up the range. Thus, a decrease in crustal strength would mean that a mountain range could not grow as high during contractional orogeny, because a weak crust would allow the range to collapse and undergo lateral spreading before it built up as high mountains. Since the strength of the crust decreases as temperature increases due to the temperature dependence of deformation efficiency, crust with a higher geotherm will be weaker than crust with a lower geotherm. This contrast would imply that the width and height of Archean and Early Proterozoic orogenic belts would be less than those of Phanerozoic orogenic belts for a given amount of horizontal convergence (Figure 13.31). Thus, the cross-strike geometry of mountain ranges might have been different in the past; for example, Archean and Early Proterozoic orogens may have contained wider belts of plastically deformed rock.

FIGURE 13.31. Schematic cross-sections that contrast collisional orogens of Archean time with those of Phanerozoic time. The medium shaded layer represents supracrustal rocks, the light layer represents basement, and the dark lens represents a mid-crustal weak zone. (a) Phanerozoic collisional orogen with adjacent foreland basin. (b) Archean collisional orogen. The thin horizontal line above the orogen defines the comparative height of the Phanerozoic orogen. [20.4]

In addition to geometric observations, we increasingly recognize interconnections between the atmosphere, climate, erosion, and tectonics. Can changes in environmental conditions affect regional geology over time? If atmospheric circulation and climate were significantly different in the past, then ancient orogenic belts may have been different, both morphologically and structurally, from modern orogens. Earth’s early atmosphere may have been more corrosive than the modern atmosphere, because of the greater concentration of volcanic gases. If so, rainfall might have caused chemical weathering at faster rates than today. If the Archean and early Proterozoic atmosphere was richer in CO2 than the Phanerozoic atmosphere, the Earth was probably warmer most of the time, so atmospheric circulation and oceanic evaporation might have been faster, leading to greater rainfall. Because continents were smaller in Earth’s early history, storms would not be calmed by movement over broad areas of land. Thus, weathering and erosion may have been faster during Earth’s earlier history than they are today. So, exhumation rates would be faster and isotherms in the crust would rise significantly. To replace the mass deficit resulting from erosion, rocks metamorphosed at great depth would be rapidly brought to the surface, producing wide metamorphic belts as relicts of orogens. When tectonism eventually ceased, the next succession of supracrustal rocks would be deposited directly on high-grade gneiss. Further, foreland fold-thrust belts would be smaller, basement structures would be reactivated in the foreland (because uplift of isotherms would bring hot rocks to the surface in the foreland), and deep foreland basins would not develop. Although far from proven, these speculations offer a rich framework for comparison of modern and ancient orogens and a further understanding of plate tectonic processes.

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v1.2 Last update: 15-Sep-2023