Eastern Wopmay orogen comprises an external thin-skinned thrust-fold belt and a contiguous metamorphic-plutonic internal zone produced by the closing of a short-lived (1.90–1.88 Ga) ensialic(?) back-arc(?) basin on the west margin of Slave craton, an Archean granite-greenstone terrane in the northwest Canadian shield. Sets of high-amplitude, basement-involved folds subsequently developed parallel (D2) and oblique (D3) to the strike of the thin-skinned (Dl) thrust-fold belt. The D3 folds permit the construction, by means of axial projections, of composite cross sections of the Dl and D2 structures through a depth interval of 30 km. The cross sections show that the sole thrust of the externides and the underlying autochthonous cover continue beneath the metamorphic-plutonic internal zone, which was thrust as a hot allochthon onto the relatively cold Archean craton. In the externides, thrusting produced a minimum of 40 percent east-west shortening and thickening of the allochthonous cover. This is manifested by about 25 mappable imbrications of a lower cratonic shelf sequence (stratigraphic thickness of 1.5 to 2.5 km), and upright chevron-type folds in overlying foredeep flysch and molasse (minimum, 5 km of stratigraphic thickness). Thrust propagation was eastward (toward the craton), and an active accretionary wedge of relatively low taper is indicated by high basal “step-up” angles, common back-thrusts, and lack of overall change in exposed stratigraphic level across the belt. The sole thrust is consistently positioned 100 to 300 m above the unfaulted basement surface, independent of footwall lithology, for at least 100 km west of the frontal thrusts. In the internides, a basal allochthonous assemblage of rift-facies clastics and bimodal volcanics is characterized by recumbent isoclinal folds, most of which are east-vergent. Allochthonous syn-rift deposits occur in both the eastern and western internides, but the overlying slope-rise facies of the cratonic margin sequence and succeeding synorogenic foredeep deposits are limited to the east. The slope-rise and foredeep strata are deformed by east-vergent thrusts and folds with steeply dipping axial surfaces. The medial and western internides are occupied by a salic-through-mafic plutonic suite (Hepburn Intrusive Suite), emplaced as the basin closed. North-trending folds and thrusts both predate and postdate prograde metamorphic isograds that envelope the composite Hepburn batholith of the medial internides. Eastward translation of the hot batholith on the sole thrust caused a minimum 250°C inversion of the contemporaneous geothermal gradient beneath the allochthon. This is evident from inverted metamorphic isograds that transect the sole thrust and the underlying, upward-facing, autochthonous cover. Following thin-skinned deformation, the allochthon and autochthon were congruently shortened to form large-scale north-trending folds (D2) having a ca. 35 km wave-length and an apparent structural relief of 13 and 6 km on the basement surface in the internides and externides, respectively. Thrusting of the basement-cover contact is insignificant, and the folds, which are completely exposed in oblique cross section, are clearly not the result of basement involvement in D1 thrusting. However, the existence of a blind thrust at depth is not ruled out. The cross folds (D3) are similar in style to the D2 folds but trend northeast and have first-order wavelengths of 80 to 140 km and amplitudes at the basement-cover interface of up to 15 km. Locally, D3 produced belts of well-developed higher order folds, typically having cuspate-lobate basement-cover profiles, and inhomogeneous but locally strong northeast-striking foliations. In common with D2 folds, the D3 folds have no significant associated thrusts, and are well developed in areas where the contemporary basement temperature exceeded a threshold of ca. 350°C. Zones of strong D3 deformation occur east of the frontal D1 thrusts, far beyond the limit of tectonically transported heat. The observed evolution in modes of crustal thickening—early thin-skinned thrusting followed by coaxial, and later transverse, thick-skinned folding—appears to have occurred in several other well-exposed, broadly coeval, orogenic belts in the Canadian shield. Detailed mapping of the steeply plunging segments of these orogens will allow their three-dimensional structural and metamorphic configurations to be elucidated by the use of axial projections, as successfully employed in Wopmay orogen.

The tectonic evolution of the central Andes is depicted through analysis of the 30° to 33°S segment, which encompasses the highest part of the Andean Cordillera. A period of rifting, starting as early as 600 Ma, was a milestone in the evolution of the different tectonic regimes that are responsible for the present geological composition and structure of the Andes. The early Paleozoic was an important period of continental accretion when allochthonous terranes, such as Chilenia, were incorporated onto the western margin of Gondwanaland. The subduction zone was located about 300 km east of the present trench. An early Paleozoic magmatic arc was developed in the western Sierras Pampeanas. Sedimentary facies of that age record a continental margin between Chilenia and the magmatic arc, which is associated with a disrupted ophiolite sequence. The Famatinides orogeny produced the first deformation of the Andes and the uplift of the Protoprecordillera during middle to late Devonian times. The Gondwanides orogeny is characterized by subduction of an oceanic plate beneath a continental margin, with the accretion of minor exotic terranes in the southern part of the Andes. Magmatism, eastward migration of the volcanic front, sedimentation pattern, and deformation defined an evolving orogenic sequence. This is correlated to a varying convergence history linked to variations in the apparent polar wandering path of western Gondwanaland during late Paleozoic-early Mesozoic times. The Patagonides orogeny was also governed by changes in relative plate motions during middle to late Mesozoic times. The paleotectonic history suggests that two orogenic styles were produced: a stage with little compression and back-arc volcanism, and a stage with high compression without volcanism but with important deformation and emplacement of postorogenic granitoids. The Andean orogenic cycle is distinguished by conspicuous segments of the orogen that are controlled by the segmentation of the subducted oceanic Nazca plate related to the subduction of aseismic ridges. But the compressive phases and mountain building are more closely related to changes of plate motion, which affected thousands of kilometers of the continental margin, exceeding the length of any individual segment. The age of the subducted oceanic slab is an important factor that controls the magmatic activity and the presence or absence of retroarc magmatism. The relative influence of each proposed mechanisms varies substantially among the different segments.

Approximately 3,000 Ar, Sr, and Pb isotopic age determinations for Canadian Cordilleran rocks have been cataloged, categorized as to reliability and significance, and plotted on histograms, distribution maps for different time intervals, and space-time plots to show the magmatic evolution in this 2,300-km portion of the Circum-Pacific Mobile Belt. The history revealed is episodic, with stable distribution patterns within episodes and distinct lulls and changes in distribution between the episodes. From 230 to 214 Ma (during Late Triassic time), extensive mafic volcanism occurred in the Wrangell, Quesnel, and Stikine terranes. Volcanic-related ultramafic complexes are found scattered through the two latter terranes. Large calc-alkaline granitic plutons are known only in a belt crossing Stikinia in northern British Columbia. At the same time, blueschists formed in the Cache Creek accretion wedge. From 214 to 200 Ma (end of Triassic and part of Early Jurassic time), Early to Middle Jurassic arc magmatism began in Wrangellia and in the northern Quesnel, Stikine, and Yukon terranes. A distinct magmatic event is recognizable only in southern Quesnellia. Magmatism was absent on the North American craton. The Cache Creek and Quesnel terranes were definitely linked, Stikine and Cache Creek terranes were probably linked, and a regional metamorphic episode was completed in the Yukon Terrane by this time. From 200 to 155 Ma (late Early to early Late Jurassic time), magmatism was extensive in the Wrangell, Quesnel, Stikine, and Yukon terranes. Magmatism over-lapped into North America only east of southern Quesnellia after about 180 Ma. By the middle of this time interval, the southern Quesnel-Slide Mountain-North America linkage was complete, and major deformation and metamorphism had affected the Omineca Belt in British Columbia. Early to Middle Jurassic magmatism in southern Wrangellia (Vancouver Island) is distinctly older than the Middle to Late Jurassic magmatism that occurred in central Wrangellia (Queen Charlotte Islands). From 155 to 140 Ma (during Late Jurassic time), a few last-gasp plutons of the late Early to early Late Jurassic episode and other rocks with partially reset 155- to 145-Ma dates occur in the Wrangell, Quesnel, and Stikine terranes. Late Jurassic magmatism (160 to 140 Ma) occurred in the Alexander Terrane (Saint Elias region). From 145 to 138 Ma (latest Jurassic and beginning of Early Cretaceous time), plutonism occurred in the Endako area of central British Columbia (Francois Lake suite) but is virtually unknown elsewhere. From 135 to 125 Ma (during Early Cretaceous time), there was a magmatic lull of major significance present throughout western North America. From 110 to 90 Ma (middle Cretaceous time), widespread plutonism occurred across all terranes. Dual culminations are evident: the Coast Plutonic and Ominica belts. Before this time all sutures except those outboard of Wrangellia had been closed. From 80 to 70 Ma (during Late Cretaceous time), a narrow, sinuous belt of magmatism persisted, mostly in the southeastern Coast Plutonic Belt, southwestern Yukon Territory, and scattered across the Skeena and Stikine arches. From 70 to 60 Ma (latest Cretaceous to Paleocene time), a distinct lull in magmatism occurred. Rare plutons of this time interval are known in the Coast Plutonic Belt, on the Skeena Arch, and in the southern Intermontane Belt. From 55 to 45 Ma (latest Paleocene to Middle Eocene time), widespread and voluminous magmatism occurred in all terranes. The early Cenozoic volcanic front crossed the Coast Plutonic Complex from its east side in the south to its west side in the north. Associated thermal and tectonic effects were strong even into the Omineca Belt, producing large reset metamorphic areas in the Coast and Omineca belts. This was a short-lived event, synchronous from southern British Columbia through the Yukon Territory. West of the volcanic front, offshore of Wrangellia, Metchosin volcano growth was underway at this time. Late in this time interval, the 50?–45–36-Ma Catface–Leech River event(s) of southern Wrangellia occurred. There is also time overlap with a diffuse Massett magmatic event in the Queen Charlotte Islands, and with Baranoff Island and Yakutat–Saint Elias region magmatism. Initial 87 Sr/ 86 Sr ratios and petrographic characteristics of Canadian Cordilleran igneous rocks are reviewed in the time frame just described. These reflect the nature of underlying crust, contemporaneous lithosphere thickness, and distance from the subduction zone. Comparisons with other parts of the Circum-Pacific Magmatic Belt shows both out-of-phase magmatism (Japan and southwestern Alaska) and perfect matching of some episodes (Sierra Nevada). Major magmatic episodes correspond to times of increased westward motion of North America with respect to hot spots or to times of increased convergence between western North America and the Farallon Plate.

The west-east-directed part of the Penninic ocean basins, with their associated intraoceanic rises of the Alps and the western Carpathians, is considered as a complex Jurassic-Early Cretaceous transform zone that links the northern end of the Piemont-Ligurian ocean with the northern end of the Vardar ocean. The Maghreb transform belt may furnish an analogous example. Several traverses through the Alps are briefly described. Correlation between the Penninic zones of the Western and Eastern Alps is still controversial.

We use the paleogeographic reconstructions of the Tethyan belt made during a recent French-Soviet cooperative program (Dercourt and others, 1985, 1986) to consider the relationship between plate kinematics and plate tectonics. These reconstructions are based on kinematic and paleomagnetic syntheses. They enable us to obtain a quantitative estimate of the total amount of shortening along the Tethyan belt, both in subduction and in collision. We propose that, during Mesozoic time, the basic system was a three-plate system that involved the northward migration of a mid-ocean ridge across the ocean, its subduction beneath Eurasia, and the subsequent formation of a new Rift along the African margin. This might explain consecutive northward migrations of microcontinents across the Tethys. Our quantitative evaluation of continental collision indicates that about half the total amount of continental crust involved must have disappeared into the mantle. We suggest that this is because only the upper brittle part of the crust is involved in the stacking of nappes, whereas the lower ductile crust disappears into the mantle. As a result, the Tethyan continental collision resulted in a 75-m lowering of sea level. A discussion of the paleo-stress trajectories over the West European platform since the Eocene shows that Oligocene time was dominated by east-west extension related to relative motion between western Europe and central Eurasia. This relative motion is confirmed by the difference between the plate kinematics deduced from Arctic magnetic lineations and those deduced from North Atlantic magnetic lineations. The paleo-stress trajectories are remarkably parallel over most of the West European platform at all stages. However, since latest Miocene time, there is an outward fanning of the compressive stress trajectories in a zone 300 to 150 km wide, along the forefront of the Alps. The significance of this zone of fanning stress trajectories is not clear.

Cordilleran detachment faults, as defined here, are extensional faults of low initial dip, probably less than 30°, and subregional to regional scale. Some detachment faults have large translational displacements, i.e., in excess of several tens of kilometers. First interpreted as Tertiary extensional structures in the eastern Great Basin by Armstrong (1972), they are now known to be widespread throughout those Cordilleran regions that have undergone greatest Cenozoic extension. Detachment faults are commonly, but not necessarily, associated with lower-plate mylonitic gneisses that compose the so-called “metamorphic core complexes.” Probably nowhere in the U.S. Cordillera are detachment faults more widely and spectacularly developed than in the region that borders the lower Colorado River in southernmost Nevada, southeastern California, and southwestern Arizona. We believe that our studies and those of numerous other workers in this region, the Colorado River extensional corridor of Howard and John (1987), provide a number of new perspectives on the origin, geometry, and evolution of Cordilleran detachment faults. Detachment faults are best explained as evolving shallow-dipping shear zones that have accommodated Tertiary crustal extension (Wernicke, 1981). The fault zones are believed to root at midcrustal or lower upper crustal depths into broad zones of intra-crustal flow, the tectonic regime in which mylonitic gneisses form. At their upper ends, major detachment faults either reach the surface directly or terminate at shallow depth into pull-apart complexes of closely spaced normal faults. Along these evolving shear zones, lower-plate mylonitic gneisses are drawn upward and out from beneath upper-plate rocks. As footwall gneisses rise structurally upward, they are retrograded, sheared, and shattered at progressively colder and shallower crustal levels to form the chloritic breccias and microbreccias characteristic of many major detachment faults. At advanced stages of detachment fault evolution, lower-plate mylonitic gneisses formed at depths >12 km are tectonically juxtaposed beneath unmetamorphosed supracrustal rocks and exposed at the surface through combinations of crustal upwarping, tectonic denudation, and erosion. Contrary to popular belief, the master detachment faults exposed today are probably not in their entirety those faults that formed at the start of extensional deformation, but rather are only the youngest in a succession of major detachment faults. Detachment faults undergo warping at high angles to the direction of crustal extension, probably in large part related to isostatically induced distortions of originally more planar faults. Such warping leads to the development of younger, more planar fault splays that either cut upward into former upper-plate rocks (excisement) or downward into former lower-plate rocks (incisement). Recognition of such geometric complexities offers fresh insights into deciphering the evolving strain patterns within major detachment terranes. Studies in the Whipple Mountains region of southeastern California indicate that: (1) detachment faults have formed by both excisement and incisement tectonics; (2) northeast-southwest-trending “folds” of major detachment faults, oriented parallel to the direction of extension, are in reality primary corrugations or flutes in the fault surface (a conclusion previously reached by other workers in nearby areas); (3) most normal faults in the upper plates of major detachments originally had listric geometries before losing their flattened lower segments as the consequence of excisement tectonics; and (4) detachment faults can transect upper crustal rocks as primary, low-dipping shear zones without pre-existing, shallow-dipping structural controls (e.g., thrust faults) on their localization; the northeast-southwest-trending curviplanar geometry of the Whipple fault does, however, seem to mimic preexisting fold structure in lower-plate mylonitic gneisses crossed by the fault. Finally, the rate of translation along master faults of some evolving detachment systems apparently can be very rapid (>1 cm/yr), much faster than rates that we (and perhaps other workers) once deemed reasonable. A very good case can be made on the basis of geochronologic and field studies that footwall mylonitic gneisses were transported upward along the Whipple detachment system from lower upper crustal depths to near surface levels in less than 2 m.y. (between 18 and 20 Ma).

Active fold-and-thrust belts and accretionary wedges along compressive plate boundaries are analogous to the wedges of soil or snow that form in front of moving bulldozers. Such wedges deform until they attain a critical taper, which corresponds to an internal state of stress that is everywhere on the verge of Coulomb failure. The critical taper depends on the wedge cohesion S 0 as well as the internal and basal coefficients of friction μ and μ b and the internal and basal Hubbert-Rubey pore-fluid pressure ratios λ and λ b ; an exact relation can be obtained if the cohesion increases linearly with depth, i.e., S 0 = Kz . Typical laboratory rock-strength parameters in the range μ = μ b = 0.6 to 0.85 and S 0 = 5 to 18 MPa are consistent with the known taper, thrust-fault dips, and pore-fluid pressures in the western Taiwan fold-and-thrust belt. Much lower rock strengths in the range μ = μ b ≈ 0.2 and S 0 ≈ 0, however, can also satisfy the Taiwan data. As fresh material enters the toe of a fold-and-thrust belt, accretionary wedge, or bulldozer wedge, the wedge grows self-similarly, maintaining its critical taper. A submarine or other noneroding wedge widens with time like t 1/2 , whereas an eroding wedge attains a steady-state width when the accretionary influx is balanced by erosion. The rate of erosion exerts a significant control on the deformation and metamorphic histories of rocks incorporated into mountain belts. To illustrate this, we develop a simple kinematical model of self-similar wedge growth, and use it to infer the mean trajectories and residence times of rocks in the rapidly eroding fold-and-thrust belt in central Taiwan, which develops a steady-state width of about 90 km. A typical rock resides in the Taiwan wedge for 2 to 3 m.y. before being uplifted and eroded, and it experiences strain rates in the range ∊̇ = 10 −13 to 10 −14 s −1 and finite strains in the range S = 1 to 10. Qualitatively, the model predicts that rocks transported farther into the wedge have been buried deeper and thus have been subjected to greater pressures and temperatures, in general agreement with the observed gradient of metamorphism in Taiwan and other mountain belts. In contrast with rapidly eroding mountain belts like Taiwan, more slowly eroding belts can grow to widths in excess of 300 to 500 km and have mean residence times greater than 50 m.y., which may be too long to reach true steady state before conditions change.

We review separately aspects of two types of forces that resist mountain building and therefore that affect the support, deep structure, and evolution of mountain ranges, using observations from the large-scale tectonics of Asia and the Andes to illustrate them. The first such force might be termed mechanical strength. In its simplest description, the lithosphère is flexed as an elastic plate under the weight of a mountain range thrust on top of the lithosphère. The gross shapes of foredeep basins and gravity anomalies over them show that the simple analyses, in terms of elastic plates over inviscid fluids, are reasonable first approximations. At the same time, the weights of some mountain ranges are inadequate to depress lithospheric plates to the depths of neighboring foredeep basins, and the weights of others would create deeper foredeep basins than observed if additional forces were not present. Thus, the strength of the lithosphère alone does not support all ranges. The strength of the lithosphère, however, does affect the geometry of major, deeply rooted thrust faults, which seem to behave as crustal-scale ramp-overthrust faults. Seismic activity seems to be concentrated on the steeper, deeper sections, whereas slip on the flatter planes parallel to the underthrusting basement seems to occur aseismically. The second force that resists mountain building is gravity; the forces that drive two plates together and that cause crustal thickening must do work against gravity. More gravitational potential energy is stored in a column of mass that includes a high range and thick crustal root than is in a column that is in isostatic equilibrium with the mountain belt but with thinner crust. Because of the increasing amount of work that must be done against gravity acting on an increasingly higher range, the range should reach a maximum mean elevation related to the force at which the plates are pushed together. In this sense, the mean elevations of high plateaus serve as crude pressure gauges for the average compressive stresses pushing on the margins of the plateaus. When the maximum elevations are reached, crustal shortening need not cease; convergence can continue as the range builds outward, growing laterally into a high plateau. Moreover, changes, possibly small ones, in the driving forces can lead to situations in which the crests of high ranges or high plateaus can undergo crustal extension while crustal shortening continues on the flanks of the range. From the simple analogy with the pressure gauge, we can use the simultaneous occurrence of crustal extension at high altitudes with crustal shortening on the flanks of the range to place crude limits on the average strength of the lithosphère at mountain ranges. Finally, we discuss aspects of the tectonic evolution of western North America that appear to be analogous with aspects of the active tectonics of the Himalaya, the Tibetan plateau, and the Tien Shan in Asia and the Andes.