A mountain belt is many tens to hundreds of kilometres wide and hundreds to thousands of kilometres long. It stands above the surrounding surface, which may be a coastal plain, as along the western Andes in northern Chile, or a high plateau, as within and along the Plateau of Tibet in southwest China. Mountain ranges or chains extend tens to hundreds of kilometres in length. Individual mountains are connected by ridges and separated by valleys. Within many mountain belts are plateaus, which stand high but contain little relief. Thus, for example, the Andes (Andes Mountains) constitute a mountain belt that borders the entire west coast of South America; within it are both individual ranges, such as the Cordillera Blanca in which lies Peru's highest peak, Huascarán, and the high plateau, the Altiplano, in southern Peru and western Bolivia.

Chains of active volcanoes, such as those occurring at island arcs, are commonly marked by individual high mountains separated by large expanses of low and gentle topography. In some chains, namely those associated with “hot spots” (see below), only the volcanoes at one end of the chain are active. Thus those volcanoes stand high, but with increasing distance away from them erosion has reduced the sizes of volcanic structures to an increasing degree.

The folding of layers of sedimentary rocks with thicknesses of hundreds of metres to a few kilometres often leaves long, parallel ridges and valleys termed fold belts, as, for example, in the Valley and Ridge province of Pennsylvania in the eastern United States. The more resistant rocks form ridges, and the valleys are underlain by weaker ones. These fold belts commonly include segments where layers of older rocks have been thrust or pushed up and over younger rocks. Such segments are known as fold and thrust belts. Typically their topography is not as regular as where folding is the most important process, but it is usually dominated by parallel ridges of resistant rock divided by valleys of weaker rock, as in the eastern flank of the Canadian Rocky Mountains or in the Jura Mountains of France and Switzerland.

Most fold and thrust belts are bounded on one side, or lie parallel to, a belt or terrain of crystalline rocks (crystalline rock). These are metamorphic and igneous rocks that in most cases solidified at depths of several kilometres or more and that are more resistant to erosion than the sedimentary rocks deposited on top of them. These crystalline terrains typically contain the highest peaks in any mountain belt and include the highest belt in the world, the Himalayas, which was formed by the thrusting of crystalline rocks up onto the surface of the Earth. The great heights exist because of the resistance of the rocks to erosion and because the rates of continuing uplift are the highest in these areas. The topography rarely is as regularly oriented as in fold and thrust belts.

In certain areas, blocks or isolated masses of rock have been elevated relative to adjacent areas to form block-fault mountains or ranges. In some places, block-fault ranges with an overall common orientation coalesce to define a mountain belt or chain, but in others the ranges may be isolated.

Block faulting can occur when blocks are thrust, or pushed, over neighbouring valleys, as has occurred in the Rocky Mountains of Colorado, Wyoming, and Utah in the western United States or as is now occurring in the Tien Shan, an east–west range in western China and Central Asia. Within individual ranges, which are usually a few hundred kilometres long and several tens of kilometres wide, crystalline rocks commonly crop out. On a large scale, there is a clear orientation of such ranges, but within them the landforms are controlled more by the variations in erosion than by tectonic processes.

Block faulting also occurs where blocks are pulled apart, causing a subsidence of the intervening valley between diverging blocks. In this case, alternating basins and ranges form. The basins eventually fill with sediment, and the ranges—typically tens of kilometres long and from a few to 20–30 kilometres wide—often tilt, with steep relief on one side and a gentle slope on the other. The uniformity of the gently tilted slope owes its existence to long periods of erosion and deposition before tilting, sometimes with a capping of resistant lava flows on this surface prior to tilting and faulting. Both the Tetons of Wyoming and the Sierra Nevada of California were formed by blocks being tilted up toward the east; major faults allowed the blocks on their east sides to drop steeply down several thousand metres and thereby created steep eastern slopes.

In some areas, a single block or a narrow zone of blocks has subsided between neighbouring blocks or plateaus that moved apart to form a rift valley between them. Mountains with steep inward slopes and gentle outward slopes often form on the margins of rift valleys. Less commonly, large areas that are pulled apart and subside leave between them an elevated block with steep slopes on both sides. An example of this kind of structure, called a horst (horst and graben), is the Ruwenzori in East Africa.

Finally, in certain areas, including those that once were plateaus or broad uplifted regions, erosion has left what are known as residual (residual landform) mountains. Many such mountains are isolated and not part of any discernible chain, as, for instance, Mount Katahdin in Maine in the northeastern United States. Some entire chains (e.g., the Appalachians in North America or the Urals in Russia), which were formed hundreds of millions of years ago, remain in spite of a long history of erosion. Most residual chains and individual mountains are characterized by low elevations; however, both gentle and precipitous relief can exist, depending on the degree of recent erosion.

Two properties of rocks contribute to the support of mountains, mountain belts, and plateaus, namely strength and density. If rocks had no strength, mountains would simply flow away. At a subtler level, the strength of the material beneath mountains can affect the scale of the topography.

In terms of strength, the lithosphere, the thickness of which varies over the face of the Earth from a few to more than 200 kilometres, is much stronger than the underlying layer, the asthenosphere (see plate tectonics). The strength of the lithosphere is derived from its temperature; thick lithosphere exists because the outer part of the Earth is relatively cold. Cold, thick, and therefore strong lithosphere can support higher mountain ranges than can thin lithosphere, just as thick ice on a lake or river is better able to support larger people than thin ice.

In terms of chemical composition, and therefore density, the Earth's crust is lighter than the underlying mantle. Beneath the oceans, the typical thickness of the crust is only six to seven kilometres. Beneath the continental regions, the average thickness is about 35 kilometres, but it can reach 60 or 70 kilometres beneath high mountain ranges and plateaus. Thus, most ranges and plateaus are buoyed up by thick crustal roots. To some extent the light crust floats on the heavier mantle, as icebergs float on the oceans.

It should be noted that the crust and lithosphere are defined by different properties and do not constitute the same layer. Moreover, variations in their thicknesses have different relationships to the overlying topography. Some mountain ranges and plateaus are buoyed up by a thick crust. The lithosphere beneath such areas, however, can be thin, and its strength does not play a significant role in supporting the range or plateau. Other ranges may overlie thick lithospheric plates, which are flexed down by the weight of the mountains. The crust beneath such ranges is likely to be thicker than normal but not as thick as it would be if the lithosphere were thin. Thus, the strength of the lithosphere supports these mountains and maintains the base of the crust at a higher level than would have been the case had the strong layer been absent. For instance, the Himalayas have been thrust onto the crust of the Indian shield, which is underlain by particularly cold, thick lithosphere that has been flexed down by the weight of the high range. The thickness of the crust is about 55 kilometres beneath the high peaks, which stand more than 8,000 metres high. The thickest crustal segment of 70 kilometres, however, lies farther north beneath the Plateau of Tibet (Tibet, Plateau of) (or Tibetan Plateau), whose altitude is about 4,500 to 5,000 metres but whose lithosphere is much thinner than that beneath the Himalayas. The strong Indian lithosphere helps to support the Himalayas, but the buoyancy of the thick Tibetan crust maintains the high elevation of the plateau.

As noted above, individual mountains, mountain ranges, mountain belts, and plateaus exist because tectonic processes have elevated terrains faster than erosion could destroy them. High elevations are created by three major processes: these are volcanism, horizontal crustal shortening as manifested by folding and by faulting, and the heating and thermal expansion of large terrains.

Most, but not all, volcanoes consist of material that is thought to have melted in the mantle (at depths of tens of kilometres), which rose through the overlying crust and was erupted onto the surface. To a large extent, the physical characteristics of the erupted material determines the shape and height of a volcano. Material of low density can produce taller mountains than can denser material. Lavas (lava) with low viscosity, such as in Hawaii, flow easily and produce gentle slopes, but more viscous lavas mixed with explosively erupted solid blocks of rocks can form steeper volcanic cones, such as Mount Fuji in Japan, Mount Rainier in the northwestern United States, or Mount Kilimanjaro in Africa.

Many volcanoes are built on elevated terrains that owe their existence to the intrusion into the crust of magmas (magma)—i.e., molten rock presumably derived from the mantle. The extent to which this process is a major one in mountain belts is controversial. Many belts, such as the Andes, seem to be underlain, at least in part, by solidified magmas, but the volume of the intruded material and its exact source (melting of either the crust or the mantle) remain poorly understood.

In most mountain belts, terrains have been elevated as a result of crustal shortening by the thrusting of one block or slice of crust over another and/or by the folding of layers of rock. The topography of mountain ranges and mountain belts depends in part on the amount of displacement on such faults, on the angles at which faults dip, on the degree to which crustal shortening occurs by faulting or by folding, and on the types of rocks that are deformed and exposed to erosion. Most of the differences among mountain belts can be ascribed to some combination of these factors.

Rocks, like most materials, expand when they are heated. Some mountain ranges and plateaus are high simply because the crust and upper mantle beneath them are unusually hot. Most broad variations in the topography of the ocean floor, the mid-ocean ridges and rises, are due to horizontal variations in temperature in the outer 100 kilometres of the Earth. Hot areas stand higher—or at shallower depths in the ocean—than cold areas. Many plateaus, such as the Massif Central in south central France or the Ethiopian Plateau, are elevated significantly because the material beneath them has been heated.

Besides erosion, which is the principal agent that destroys mountain belts, two tectonic processes help to reduce high elevations. Horizontal crustal extension and associated crustal thinning can reduce and eliminate crustal roots. When this happens, mountain belts widen and their mean elevation diminishes. Similarly, the cooling and associated thermal contraction of the outer part of the Earth leads to a reduction of the average height of a mountain belt.

Mountain belts differ from one another in various respects, but they also have a number of similarities that enable Earth scientists to group them into certain distinct categories. Each of these categories is characterized by the principal process that created a representative belt. Moreover, within individual belts different tectonic processes can prevail and can be associated with quite different landforms and topography. Thus, for any category there are exceptions and special cases, as well as subdivisions.

Volcanoes (volcano) typically form in any of three tectonic settings. At the axes of the mid-ocean ridge system where lithospheric plates diverge, volcanism is common; yet, high-standing volcanoes (above sea level) rarely develop. At subduction zones where one plate of oceanic lithosphere plunges beneath another plate, long linear or arcuate chains of volcanoes and mountain belts associated with them are the norm. Volcanoes and associated landforms, as well as linear volcanic chains and ridges (e.g., the Hawaiian chain) also can exist far from plate boundaries.

Where two lithospheric plates diverge, new material is intruded into the gap between the plates and accreted to each of them as they diverge. The vast majority of volcanic rocks ejected onto the surface of the Earth is erupted at the mid-ocean ridges and rises where this process occurs. Thus, such submarine landforms comprise very long, narrow volcanic centres. Although volcanoes do form as isolated seamounts along the axes of mid-ocean ridges, they constitute only a small fraction of the erupted material. Moreover, areas along the ridges and rises where volcanism is particularly abundant are considered unusual; the excess amount of volcanic activity is generally attributed to “hot spots” in the mantle (see below). Finally, most of the relief that defines the mid-ocean ridges and rises is not due to volcanism at all but rather to thermal expansion, as will be explained below.

Linear or arcuate belts of volcanoes are commonly associated with subduction zones. Volcanoes typically lie 150 to 200 kilometres landward of deep-sea trenches, such as those that border much of the Pacific Basin. The volcanoes overlie a zone of intense earthquake activity that begins at a shallow depth near such a trench and that dips beneath the volcanoes. They often form islands and define island arcs (island arc): these are arcuate chains of islands such as the Aleutians or the Lesser Antilles (see ocean: Deep-sea trenches (ocean)). Volcanoes usually are spaced a few to several tens of kilometres apart, and single volcanoes commonly define the width of such belts. Elsewhere, as in Japan, in the Cascade chain of the northwestern United States and southwestern Canada, or along much of the Andes, volcanoes have erupted on the margin of a continent. Nearly all features typical of an island arc, including the narrow belt of volcanoes, deep-sea trench, and intense earthquake activity, can be found at such continental margins.

The landscape of island chains of this kind is characteristically dominated by steep volcanic cones topped by small craters, and the relief between these volcanoes is low. A few such volcanoes have undergone massive eruptions and have expelled a large fraction of their interiors, as did Mount St. Helens in the northwestern United States in 1980. In the most intense eruptions of this sort, the remnants of the volcano collapse into the void at its centre, sometimes leaving a caldera (a very large crater with relatively low rims). Examples of such structures include those formed by Krakatoa in Indonesia in 1883 and by Thera (also called Santorin or Santoríni) in the Aegean Sea a few thousand years ago.

The lavas erupted at these volcanoes are thought to be derived from the mantle in the wedge of asthenosphere above the lithospheric plate plunging into it. water carried down in the interstices of the subducted rock and by hydrous minerals to which water is loosely bound chemically is expelled into the wedge of asthenosphere above the subduction zone. The introduction of water reduces the melting temperature of the rocks and allows material in the wedge to melt and rise to the surface.

Some volcanic phenomena occur at large distances from plate boundaries (for example, on the Hawaiian Islands or at Yellowstone National Park in the western continental United States). Also, as noted above, volcanism is especially intense at some parts of the mid-ocean ridge system (as in Iceland or the Galápagos Islands in the eastern Pacific). Magmas erupted in these settings originate in the asthenosphere, perhaps at depths of several hundred kilometres or more at what are called hot spots in the mantle. Such sources of melting may be due to chemical differences rather than to heat (see volcano: Intraplate volcanism (volcano)). Active volcanoes are usually localized in a region with dimensions of 100 to 200 kilometres or less.

A chain of extinct volcanoes or volcanic islands (and seamounts), like the Hawaiian chain, or a volcanic ridge, like Walvis Ridge between the islands of Tristan da Cunha and the east coast of Africa, can form where a lithospheric plate moves over a hot spot. The active volcanoes all lie at one end of the chain or ridge, and the ages of the islands or the ridge increase with their distance from those sites of volcanic activity. Older volcanoes are more eroded than younger ones and are often marked only by coral reefs that grow on the eroded and subsiding volcanic island.

Volcanic chains of this kind are not common in continental regions, in part because most continental masses move slowly over hot spots. Volcanic activity, however, can be particularly abundant when a plate moves so slowly with respect to a hot spot. Moreover, a long duration of volcanism often results in a warming of the lithosphere. This warming causes a localized thermal expansion and consequently a localized upwarping or doming (dome) of the Earth's surface, as in the case of the Yellowstone area or the Massif Central in France. The resulting domes cover areas a few to several hundred kilometres in extent, and the mean elevations are rarely as much as 1,000 metres higher than the surrounding regions. Thus, except for the isolated volcanoes that lie on the upwarps, relief is gentle and due largely to erosion.

Some hot spots are associated with massive eruptions of lava and ash, primarily of basaltic (basalt) composition, which cover vast areas as extensive as tens or hundreds of square kilometres. Such flood basalts, or traps, buried the Snake River Plain west of Yellowstone a few million years ago, the Columbia River Valley some 20,000,000 years ago, and central India (the Deccan traps) 60,000,000 to 65,000,000 years ago. Flood basalts create a remarkably flat surface that is later dissected into a network of sharply incised valleys (see plateau).

Most volcanoes that cannot be ascribed either to a subduction zone or to seafloor spreading at mid-ocean ridges are attributed to hot spots. There are, however, some volcanoes, volcanic fields, and flood basalts that cannot yet be ascribed to hot spots with any certainty. Nevertheless, the landforms associated with such volcanic phenomena resemble those in other settings for which a simple cause can be offered.

Most mountain belts of the world and nearly all of those in Europe, Asia, and North America have been built by horizontal crustal shortening and associated crustal thickening. The landforms associated with such belts depend on the rates, amounts, and types of crustal deformation that occur and on the types of rocks that are exposed to erosion. To some extent the deformation can be related to different tectonic settings. Large thrusted crystalline terrains and parallel fold and thrust belts are commonly associated with continental collisions in which two separate continents have approached each other and one has been thrust onto the other. Continental collisions are responsible for Alpine-, or Himalayan-, type mountain belts. Fold and thrust belts can also be associated with active continental margins or Andean-type margins, where oceanic lithosphere is subducted into the asthenosphere but where crustal shortening occurs landward of the volcanic arc on the overriding continental plate. Block-faulted ranges commonly form as intracontinental mountain ranges or belts, far from collision zones and subduction zones.

These belts are thought to have been created by the movement of one continent beneath another. In general, a thick layer of light, buoyant continental crust cannot be carried deep into the asthenosphere. Instead, the leading edge of the descending continent is scraped off, and the rest of the continent then plunges beneath the off-scraped slice. Eventually the convergence between the two plates carrying the continents comes to a halt, but usually not before several slices of continental material have been removed from the underthrusting continent and stacked on top of it.

The sedimentary rocks deposited on the continental crust and its margin long before the collision often constitute one or part of one of the off-scraped slices. They commonly are deformed into a fold and thrust belt as the basement under them continues to plunge beneath the overriding plate at the subduction zone. Layers of strong sedimentary rock detach from the underlying basement at weak layers that commonly consist of evaporites (salt, gypsum, or anhydrite) or of shale by a process called décollement (from the French word meaning “ungluing”). The stronger layers of sedimentary rock are then folded into linear, regularly spaced folds—alternating anticlines and synclines—and thrust on top of one another. The Valley and Ridge province of Pennsylvania, which was formed during the collision of Africa and North America near the end of Paleozoic time (about 240,000,000 years ago), is a classic example.

Convergence between two lithospheric plates can be rapid in such settings—10 to 100 millimetres per year—and the amount of displacement on the major thrust faults also can be large—tens to more than 100 kilometres. Thus, when a slice of crystalline rock from deep in the crust is scraped off the remainder of the continent and is underthrust by it, much of the slice is uplifted and pushed onto the relatively flat, ancient surface of the intact portion of the continent. Erosion generally removes the sedimentary cover of such slices and leaves expanses of crystalline rocks, as can be seen on Himalayan or Alpine peaks.

Faults along which a slice of continental crust is torn from the rest of the continent and thrust onto it are called ramp overthrusts. When the fault first forms, it dips at 10° to 30° (or more). Slip on this fault (i.e., the movement of one face of the fault relative to the other) brings the leading edge of the off-scraped slice of crust to the surface of the Earth, where it then slides along the surface. The intact continent is flexed down by the weight of the material thrust on top of it. As a consequence, its initially flat surface dips at a very gentle angle of only a few degrees. Accordingly, a ramp overthrust consists of two segments. The first segment, the ramp, dips relatively steeply; slip on it causes uplift of the overriding slice and of the crystalline rocks from deep in the crust to create high relief and the high range. The other segment, which was once the top surface of the continent, has been flexed down and dips at a gentle angle. Slip on it allows the overthrust slice to advance over the rest of the continent, where it plows the sedimentary layers in front of it into folds and smaller overthrusts.

Collision zones are thus commonly identified by narrow belts of elevated crystalline terrain and parallel fold and thrust belts. The crystalline terrain has been thrust upward and toward the fold and thrust belt. Deformation is generally confined to shallow depths of only a few kilometres at such belts but penetrates deeply into the Earth beneath the crystalline terrains. The rapid uplift of these resistant rocks creates a high range. A crystalline terrain often exhibits large folds in which the rocks appear to have flowed instead of having been bent. Folds of this sort have formed at depths where the rocks were hot and soft before they reached the relatively cold surface of the Earth. The overthrusting of crystalline terrains onto intact continental crust can occur at rates of tens of millimetres per year, which is rapid for rates of slip on faults, and the crystalline rocks can be uplifted 10 to 20 kilometres by slip on ramp overthrusts.

At some continental margins, oceanic lithosphere is subducted. At some of these sites, the landscape is dominated by volcanoes, such as along the Cascades of western North America or in Japan, but at others, such as along much of the Andes of South America, volcanoes constitute only a small or even negligible part of the relief. At Andean-type margins, the crust is typically thicker than normal, and high mountains can exist even in the absence of volcanoes. Some of the thickened crust is due to the intrusion of magma from the mantle, and some to crustal shortening.

Oceanic lithosphere is commonly subducted at active continental margins at rates of tens to more than 100 millimetres per year, but crustal shortening within the overriding plate typically occurs at rates of only a few millimetres annually. As at continent-continent collision zones, the crustal shortening occurs both by overthrusting of crystalline terrain onto intact continental crust, which in this case lies landward of the volcanic belt, and by the formation of a fold and thrust belt within sedimentary rock lying on the intact continent. The thrusting of crystalline terrain is probably facilitated by a heating and consequent weakening of the rocks near the volcanoes. The presence or absence of a parallel fold and thrust belt depends in part on the presence or absence of thick sedimentary rocks within which detachment of separate layers can take place.

Notwithstanding large variations in topography and in the style of deformation among Andean belts in general, the scales of deformation and uplift are less than those at collision zones. Overthrust crystalline terrains are smaller, and the crystalline rocks themselves have not been thrust up from depths as great as those at collision zones. Much of the Andes, for instance, consists of sedimentary rock that never was buried deeper than a few kilometres and therefore has not been metamorphosed (heated to high temperature or put under high pressure) or at most only has been mildly metamorphosed. Topography in the high parts of the Andes is typically much gentler than in the Himalayas. The most impressive relief is on the eastern flank of the Andes where rivers responding to a wet climate have cut deep canyons.

Fold and thrust belts can be very well developed at Andean margins. The eastern Cordillera of the Bolivian Andes is an extremely wide fold and thrust belt, but only along the eastern third of the cordillera do simple parallel folds control the topography. Farther west, both the greater role of thrust faulting in the evolution of the cordillera and the longer duration of erosion have diminished the role of folding. Except where rivers have cut deep canyons, relief is not exceptionally great. Similarly while oceanic lithosphere was underthrust beneath the west coast of Canada during the Mesozoic Era (248,000,000–65,000,000 years ago), the Canadian shield was underthrust more than 200 kilometres beneath the Canadian Rocky Mountains, with crustal shortening occurring by décollement and by folding and thrust faulting within the sedimentary cover.

Thus Andean-type belts have a narrow belt of volcanoes and often a fold and thrust belt on their landward margin. The volcanoes of some belts are built on a high range that is more of a long, narrow plateau than a mountain range, for relief on it is not necessarily great.

In some regions, mountain belts have been formed by crustal shortening within a continental mass, rather than where two continents have collided. Some 40,000,000 to 80,000,000 years ago, the Rocky Mountains of Colorado, Utah, and Wyoming formed in this way, and today both the Tien Shan and the Atlas Mountains of northwestern Africa are actively forming within a continent. In general, intracontinental mountain belts are characterized by block faulting. Blocks, tens of kilometres wide and hundreds of kilometres long, are uplifted along faults that dip beneath them at angles of 25° to 45°. Because of the displacement on steep faults, crystalline rocks commonly crop out in the mountains. The edges of the ranges can be sharply defined. Fold and thrust belts are not common and are usually narrow where present.

At the edges of such ranges, sedimentary rocks are commonly tilted up, and, where resistant, they can form narrow, sharp-crested ridges called hogbacks that are parallel to the front of the ranges. A particularly prominent hogback lies along the east edge of the Front Range in eastern Colorado.

Intracontinental belts generally consist of elongated block-faulted ranges, which in some cases overlap but are not necessarily parallel to one another. Thus, in parts of the Tien Shan, two or three nearly parallel, sharply bounded ranges are separated from one another by parallel basins that are 10 to 30 kilometres wide. The ranges of this great mountain system are being overthrust onto the basins, and one such basin, the Turfan Depression, has dropped below sea level (see tectonic basins and rift valleys). In contrast with the parallel ranges in the Tien Shan, the northwest-trending Wind River Range in Wyoming, the east–west trending Uinta Mountains in Utah, and the north–south trending Front Range in Colorado are all part of the same intracontinental belt, the Rocky Mountains.

The southern part of the Andes in southern Chile and Argentina consists of a narrow range only 100 to 200 kilometres wide. A chain of volcanoes follows the axis of the range, but crustal thickening due to crustal shortening is a principal cause of the high range, and many of the volcanoes are built on folded and faulted sedimentary rock.

From northern Argentina to northern Peru and Ecuador, the Andes are much wider, with the widest segment across southern Bolivia. There, the mountain belt consists of two parallel ranges, the Cordillera Occidental (or Western Cordillera) and the Cordillera Oriental (or Eastern Cordillera (Real, Cordillera)), which surround the high plateau, the Altiplano.

The volcanic chain has been constructed on thick crust and forms the Cordillera Occidental. The Brazilian shield has been underthrust beneath the Cordillera Oriental, which comprises the western edge of a wide fold and thrust belt. This fold and thrust belt is marked by north–south trending folds and north–south trending ridges and valleys in northern Argentina and southeastern Bolivia. North of the latitude where the west coast of South America bends, the trend of the Andes, including that of both cordilleras, is northwesterly parallel to the coast of Peru. The fold and thrust belt east of the Cordillera Oriental is narrower than that farther south but is well defined by a few northwesterly trending ridges and valleys.

Lying between the two cordilleras in northern Argentina, western Bolivia, and southern Peru, the Altiplano stands at an average height of about 3,800 metres. Within it lies Lake Titicaca, the highest navigable lake in the world. The Altiplano is a high arid basin that captures sediment eroded from the eastern and western cordilleras bounding it (see plateau). Older rocks that crop out within it have been folded; thus crustal shortening probably has been an important factor in creating the high elevations and the thick crust that underlies this plateau.

In Colombia, the Andean chain diverges into three separate chains, each about 100 kilometres wide. Volcanoes occur in the westernmost chain, but all three have undergone crustal shortening. For example, the easternmost of the three, which continues into Venezuela as the “Venezuelan Andes,” is being underthrust from the northwest by the Maracaibo Basin and from the southeast by the Guiana Shield underlying southeastern Venezuela. Thus the Venezuelan Andes are an intracontinental mountain belt.

The divergence of the Andes into three chains in Colombia extends northward. The western chain continues into Panama and through Central America. The central chain continues toward the Caribbean. The Venezuelan Andes intersect an east–west trending chain along the north coast of South America in Venezuela.

At the eastern end of the Caribbean Sea, the Lesser Antilles—volcanic islands that form a typical island arc—mark a zone where a part of the floor of the North Atlantic Ocean underthrusts that of the Caribbean Sea—namely, the Caribbean Plate. This plate has moved east relative to both North and South America at a rate of 10 to 20 millimetres per year for tens of millions of years. This displacement and the consequent overthrusting of the seafloor to the east are responsible for the volcanic arc that constitutes the Lesser Antilles as well as for the strike-slip displacement occurring in Venezuela.

Most of the major islands that define the northern margin of the Caribbean—Puerto Rico, Hispaniola, Cuba, and Jamaica—are mountainous, and these mountainous terrains, like that in northern Venezuela, are remnants of the period of convergence between North and South America and also of complicated deformation along the ancient margins of the Caribbean Plate. At present, crustal shortening occurs at only a very slow rate, if at all, on these islands.

At the western margin of the Caribbean Plate another small plate, the Cocos Plate, is being underthrust beneath Mexico and Central America. A belt of volcanoes extends from northern Panama to western Mexico, and virtually all of the highest mountains in this belt are volcanic. These volcanoes are built on thickened crust, and crustal shortening has occurred within the Central American Cordillera, but the principal tectonic process that has affected the landscape is volcanism.

Subduction of oceanic lithosphere presently occurs only beneath two segments of the coast of western North America. The subduction of a small plate, the Juan de Fuca Plate, beneath northern California, Oregon, Washington, and southern British Columbia is responsible for the Cascade chain of volcanoes, which includes Mount St. Helens. The very large Pacific Plate, which underlies most of the northern and western Pacific Basin, moves north-northwest with respect to North America nearly parallel to the coast, and is subducted beneath southern Alaska and the Aleutians. The volcanic chain that forms the Aleutian Islands and continues into the Alaskan peninsula to the Wrangell Mountains is a consequence of this convergence and subduction.

Most of the North American Cordillera was built during Mesozoic and Early Cenozoic times (between about 170,000,000 and 40,000,000 years ago) when lithospheric plates converged with North America at rapid rates of many tens to more than 100 millimetres per year. The Juan de Fuca Plate is the last remnant of one of these plates. The others have been subducted beneath western North America and have completely disappeared. Thus, in Mesozoic and Early Cenozoic times, an Andean margin similar to that which presently bounds the west coast of South America bounded western North America.

The Coast Ranges of central and northern California, Oregon, and Washington consist of folded and faulted slices of oceanic crust and its overlying sedimentary rocks. Much of the rock that constitutes these mountains was scraped off the oceanic lithosphere at the trench just west of the continent. The Olympic Mountains in northwestern Washington, for instance, consist largely of off-scraped seamounts. The rock of such coastal mountains was intensely deformed and metamorphosed before being elevated to produce the present range. Specifically, the hard basalt that makes up much of the oceanic crust has been metamorphosed into the easily deformed rock serpentinite, which contains the weak, fibrous mineral serpentine. The gentle relief of the Coast Ranges is due in part to the weakness of serpentinite, a characteristic that gives rise to frequent landslides and rapid erosion.

A belt of granite lies inland and forms a mountainous zone from the axis of Baja California (in Mexico), through southern California, along the Sierra Nevada in the states of California and Nevada, northwestward into Idaho, and then north-northwestward along the western margin of the Canadian Rocky Mountains to Alaska. This granite belt underlay the volcanoes that marked the subduction zone in Mesozoic and Early Cenozoic times. The intrusion of this granite was most intense between 170,000,000 and 70,000,000 years ago during the Mesozoic Era. The Sierra Nevada of California, which contains Mount Whitney, the highest peak in the contiguous United States, is composed almost entirely of this granite.

While subduction of oceanic lithosphere occurred beneath western North America, a major fold and thrust belt developed east of the granitic belt. During Mesozoic time, the Precambrian basement of Canada and North America was underthrust westward at least 200 kilometres beneath the Andean margin, and the sedimentary rocks covering it were folded and thrust onto one another. Although present in the western United States, this fold and thrust belt is most clearly revealed in the Canadian Rockies along the border between the provinces of Alberta and British Columbia, particularly in Banff and Jasper national parks.

In sum, throughout the latter half of the Mesozoic (Mesozoic Era) from about 170,000,000 to 65,000,000 years ago, the topography of western North America probably resembled that of western South America: a trench lay offshore; a belt of volcanoes underlain by granitic intrusions marked the western edge of a high range of mountains; and a fold and thrust belt lay east of the range. The tectonic history of western North America is more complicated, however, because during this period fragments of both continents and sub-oceanic plateaus were carried to the subduction zone and collided with North America. Most of the rock now found in westernmost Canada and Alaska consists of separate terrains of rock that were independently accreted to North America and that were subsequently deformed when the next such terrain collided with it. Moreover, tectonic processes occurring during the Cenozoic (since 65,000,000 years ago) have been different from those that occurred earlier and have severely modified the landscape.

Beginning about 70,000,000 to 80,000,000 years ago, the locus of crustal shortening in the United States shifted from the fold and thrust belt, whose remnants now lie along the borders of western Utah and eastern Nevada and of western Wyoming and eastern Idaho, to eastern Utah, Colorado, and central Wyoming. Between about 70,000,000 and 40,000,000 years ago, thrust faulting on the margins of the Front Range in Colorado, the Laramie Mountains and the Wind River Range in Wyoming, and the Uinta Mountains in Utah, among others, allowed the uplift of blocks of Precambrian rock that are now exposed in the cores of these ranges. Together, these intracontinental ranges of block-faulted mountains form most of the Rocky Mountains of the United States.

During roughly the same period, volcanic rocks were erupted and deposited in parts of the Rockies, such as in southwestern Colorado in what are now the San Juan Mountains. The area that now forms the Colorado Plateau, in southern Utah and northern Arizona, underwent only very mild deformation in the form of small faults and folds and apparently lay at relatively low elevation. Sediment derived from the fold and thrust belt to its west and from the Rockies to its north and east was deposited on this relatively stable area. Thus, some 40,000,000 years ago, a high range of mountains lay along the western margin of North America. This range consisted of a volcanic chain along most of its western edge and an eroded fold and thrust belt on its eastern edge. At the latitude of Wyoming, Colorado, and Utah, another belt of mountains, the present-day Rocky Mountains, lay farther east.

The topography of the western United States has been modified extensively by tectonic processes during the last 20,000,000 years. Much of the mountainous terrain of Utah, Nevada, and California underwent large-scale crustal extension, beginning more than 40,000,000 years ago but accelerating about 15,000,000 years ago. The crustal extension approximately doubled the surface area of the region between central Utah and the Sierra Nevada, presumably with a reduction in the mean elevation of the mountains.

The present topography of the Basin and Range Province of North America is a direct manifestation of this crustal extension (see tectonic basins and rift valleys). The most prominent basins, such as Death Valley and Owens Valley in California, are small rift valleys that were formed during the last few million years. This phase of the crustal extension continues even today, with such basins becoming deeper and the surrounding ranges increasing in height. This condition is readily discernible in the case of Owens Valley and the Sierra Nevada. The occurrence of a major fault on the east side of the Sierras has allowed the valley to drop with respect to the mountain range, which has been tilted up toward the east.

Concurrent with this extension, the uppermost mantle beneath parts of the western United States has become hotter. The considerable height of the Colorado Plateau, for instance, appears to be the result of the warming of the underlying mantle during roughly the past 10,000,000 years. Such mantle heating also seems to have been responsible, at least in part, for the present elevation of much of the North American Cordillera.

The one area where rapid subduction of oceanic lithosphere (more than 50 millimetres per year) has continued is southern Alaska, where the Pacific Plate is being underthrust beneath the coast. The St. Elias Mountains (Saint Elias Mountains), the tallest in southeastern Alaska and the Yukon, appear to be the direct consequences of this convergence and rapid underthrusting. Deformation of the southern Alaskan crust extends northward several hundred kilometres to the Alaska Range, where the highest mountain in North America, Mount McKinley (McKinley, Mount), is found.

North–south crustal shortening in southern Alaska occurs both by thrust faulting and by strike-slip faulting on nearly vertical, northwesterly trending planes. Mount McKinley lies adjacent to one such major strike-slip fault, the Denali Fault. The rocks that make up Mount McKinley have been displaced several tens of kilometres northwestward relative to the rocks north of the Denali Fault and a few kilometres upward. This small vertical component, compared with the large horizontal component, has created the high peak.

In the central part of the Japanese island of Honshu, the Circum-Pacific System diverges into two chains. One continues southward along the Izu, Bonin, and Mariana islands. These volcanic islands form island arcs (island arc) where the Pacific Plate is subducted beneath the floor of the Philippine Sea to the west. Southwest of Honshu, the Ryukyu Islands are another island arc where the Philippine Sea floor is subducted beneath the Yellow Sea.

The Ryukyu island arc ends abruptly at the island of Taiwan, which is not part of the Ryukyu arc. Taiwan is a small mountainous island consisting of folded and thrusted sedimentary rocks on the southeastern margin of the Asian continent. The sedimentary rocks of Taiwan were deposited on that margin under tranquil conditions, much as sedimentary rocks have been deposited on the margins of the Atlantic Ocean. Then, in the last few million years a segment of the Asian continental margin encountered a subduction zone that dipped east-southeast. As that short segment of the margin began to be underthrust, the sedimentary rocks were scraped off its leading edge and thrust back on top of it. Thus, not only the mountains of Taiwan but also virtually the entire island consists of folded and thrust sedimentary rocks that have rapidly piled up on what had been a submerged continental shelf.

A couple of volcanic islands south of Taiwan mark the southward continuation of this subduction zone to Luzon, the large northern island of the Philippines. The mountainous landscape of the Philippine Islands is a consequence both of subduction of the South China Sea floor eastward beneath Luzon and of subduction of the Philippine Sea floor westward beneath the southern Philippine islands. Volcanism and, in Luzon, crustal shortening have built the major mountains.

A major system of island arcs extends across the Indonesian (Indonesia) islands of Sumatra and Java and eastward almost to the island of New Guinea and then again eastward along the New Britain, Solomon, and New Hebrides (Vanuatu) chains. Virtually all of the high mountains of the Sunda (Sunda Islands), or Indonesian, arc are volcanoes, some of which are associated with particularly noteworthy eruptions. In 1883 the massive eruption of the volcano on the island of Krakatoa, in the straits between Java and Sumatra, was followed by a collapse of its caldera, which caused a huge sea wave that was recorded all around the world. The eruption in 1815 of the Tambora (Tambora, Mount) Volcano on Sumbawa was perhaps the greatest in recorded history. Debris from this eruption darkened the skies for several months and caused a temporary global cooling that made 1816 “the year without a summer.” The Sumbawa volcanic arc is associated with the northward subduction of the Indian Ocean floor beneath Indonesia. Similarly, the volcanic arcs of New Britain, the Solomon (Solomon Islands) and New Hebrides (Vanuatu) islands, are associated with the northward subduction of the floor of the Solomon Sea and that of the Coral Sea beneath these island arcs.

A high range of mountains forms the backbone of the island of New Guinea between the Sunda and New Britain arcs. Whereas seafloor continues to be subducted beneath these arcs, the northern margin of the Australian continent has encountered the segment of the subduction zone between these arcs. The mountains of New Guinea consist of folded and faulted volcanic and sedimentary rocks. The volcanic rocks include both ancient seafloor and old island arcs that were thrust up and onto the northern margin of Australia. The sedimentary rock includes a full complement of Paleozoic, Mesozoic, and Cenozoic rock deposited in the tranquil conditions of an ancient continental shelf. Thrust faulting has elevated metamorphic rock to the crest of the high range where glaciers persist even at the Equator, while the sedimentary rock is being deformed in a fold and thrust belt along the southern margin of the range.

East of the New Hebrides Islands, the Circum-Pacific System is defined by the Tonga and Kermadec Islands, volcanic islands associated with the westward subduction of the Pacific Plate. The subduction zone continues southward to the North Island of New Zealand, where volcanism is the principal tectonic process that has created mountains and relief. The mountains of the South Island of New Zealand, however, have been produced by different tectonic processes. Whereas the convergence between the Pacific Plate and the seafloor beneath the Tasman Sea manifests itself as subduction of the Pacific Plate at the Tonga-Kermadec-North Island zone, it results in crustal shortening across the South Island. The Southern Alps of New Zealand have resulted from this crustal shortening, which occurs by folding, by thrust faulting, and by vertical components of slip on predominantly strike-slip faults that trend southwest across the northern and western parts of the island. Rapid uplift, possibly as much as 10 millimetres per year, keeps pace with the rapid erosion of the easily eroded schists of the Southern Alps.

The Circum-Pacific System continues southwest of New Zealand along a submarine ridge, the Macquarie Ridge. In short, the Circum-Pacific System consists of a variety of mountain types and ranges where different tectonic processes occurring at different geologic times in the past have shaped the landscape. The grouping of these different belts into this single system is thus only a crude simplification.

As in the case for the Circum-Pacific System, the grouping of these different mountain ranges into a single system is an oversimplification. The various ranges (and plateaus) of the Alpine-Himalayan System formed at different times, at different rates, and between different lithospheric plates, and consist of different types of rocks.

Physiographically, this chain can be subdivided into three parallel belts: the Lesser Himalayas, the Great Himalayas, and the Tethys Himalayas. (Some authorities prefer a subdivision into four belts, the additional one designated the Outer, or Sub-Himalayas.) The Great Himalayas are defined by an arcuate chain of the highest peaks. To the south lie the Lesser Himalayas, a belt about 100 kilometres wide with an average elevation of 1,000 to 2,000 metres that is dissected by the rivers emanating from the Great Himalayas and north of it. To the north, the Tethys Himalayas form the southern edge of the Tibetan Plateau.

The rocks of the Lesser Himalayas consist primarily of mildly metamorphosed sedimentary rock largely of Precambrian age. At present, the remainder of the Indian subcontinent underthrusts the Lesser Himalayas on a very gently dipping thrust fault, so that the rocks forming this belt are sliding over the ancient top surface of India. As a result, the uplift of the Lesser Himalayas seems to be relatively slow.

The rate of uplift in the Himalayas seems to be rapid in two parallel zones: (1) at the very front of the range where the ancient metamorphic and sedimentary rocks of the Lesser Himalayas have been thrust up and onto the young sediments, and (2) beneath the Great Himalayas. The thrust fault that carries the Himalayas onto the intact part of India is a ramp overthrust, with the steep part of the ramp dipping north beneath the Great Himalayas. Slip on this steep part allows the rapid uplift of the Great Himalayas, which in turn creates the high peaks and carries rock from deep in the crust to the Earth's surface.

Most of the constituent rocks of the Great Himalayas are metamorphic; they once constituted the middle and lower crust of India's ancient northern margin but were subsequently scraped off and thrust up onto the surface. The very tops of many of the peaks, however, consist of Paleozoic sedimentary rocks, which dip northward. North of the Great Himalayas, in the Tethys Himalayas, these Paleozoic rocks and the Mesozoic sedimentary rocks deposited on them along the southern edge of the Tethys Ocean have been folded and faulted into east–west ridges.

Geologically, the northern margin of the Himalayas follows the Indus River in the west and the Brahmaputra River (also called Tsang-po or Yarlung Zangbo Jiang) in the east. The last remnants of the Tethys Ocean floor can be found in what some refer to as the Indus-Tsang-po Suture Zone, where a jumble of volcanic and sedimentary rocks have been folded and thrust over one another in a narrow zone parallel to these rivers. North of this suture, a belt of granites forms the backbone of the Trans-Himalayan range. These granites were intruded into the crust of the southern margin of Asia between 120,000,000 and 50,000,000 years ago, when the Tethys Ocean floor was being subducted beneath southern Asia and before India collided with it.

Since India collided with Eurasia, it has penetrated 2,000 kilometres or more into the ancient Eurasian continent. The northern edge of India may have been subducted a few hundred kilometres beneath southern Tibet, but most of its penetration has been absorbed by crustal shortening north of the collision zone. The crust of the Tibetan Plateau (Tibet, Plateau of) appears to have been severely shortened; the thickness of its crust has approximately doubled. Although much of Asia underwent extensive deformation during phases of mountain building in Late Precambrian, Paleozoic, or Mesozoic times, the high altitudes of all of the mountain ranges surrounding the Tibetan Plateau, including the Pamir, Karakoram, Kunlun, Nan, Ch'i-lien (Qilian), and Lung-men (Longmen) mountains, have formed since India collided with Eurasia.

In some areas, blocks of crust undeformed since Precambrian time, such as beneath the Takla Makan (or Tarim Basin), have remained in that state, but have been displaced in response to India's penetration. The northward displacement of the Tarim Basin has caused intracontinental crustal shortening in the Tien Shan, the aforementioned east–west trending mountain range with peaks exceeding 7,000 metres in height that lies 1,000 to 2,000 kilometres north of India's northern edge. The Tien Shan was the site of Late-Paleozoic mountain building, but by the time India collided with Eurasia erosion had planed down the ancient Tien Shan to a featureless terrain buried in its own sediment. The present elevation therefore seems to be a consequence of India's penetration into Asia notwithstanding its great distance from India itself.

The penetration of India into Eurasia not only has caused crustal thickening in front of itself, but it also is squeezing parts of Asia eastward out of its northward path. One manifestation of this extrusion of material out of its path is the crustal shortening on the eastern margin of the Tibetan Plateau, where crustal thickening is actively occurring. The eastward displacement of crustal blocks along major strike-slip faults also seems to have caused rift systems to open in a northwest–southeast direction. The Baikal Rift Zone in Siberia and the Shansi Graben in northern China seem to have resulted from the east-southeastward extrusion of material out of India's path. Moreover, crustal thickening in the Tibetan Plateau has ceased, and now east–west extension of the plateau contributes to the eastward extrusion. The plateau is laced with northerly trending rift zones that are bounded by northerly trending tilted, block-faulted mountains. Thus, virtually all of the tectonic landforms of Asia seem to be attributable, directly or indirectly, to India's collision with Eurasia and its subsequent penetration into the continent.

The eastward extrusion has been facilitated by the lack of any major obstacle to the eastward displacement of South China. Minor westward displacement of material also has occurred in Afghanistan, but this process has been blocked by the collision of Arabia with southern Iran and Turkey, where to some extent the same processes have occurred as in eastern Asia.

While these overlying sedimentary rocks have become detached and folded, the penetration of the Arabian shield into Iran and Turkey has built plateaus in front of it and mountain ranges on the north sides of the plateaus: these include the Kopet-Dag (Kopet-Dag Range) and Elburz (Elburz Mountains) ranges north of the central Iranian plateau and the Caucasus north of the Anatolian plateau. North–south crustal shortening is the principal process by which these ranges were built, but volcanism has contributed in some cases. Many of the high mountains in this area are volcanoes, including Mount Demavand (Demavend, Mount), which towers over the city of Tehrān, Mount Ararat (Ararat, Mount) on the border of Turkey and Armenia where Noah reputedly landed, and Mount Elbrus (Elbrus, Mount), the highest peak in the Caucasus. The penetration of the Arabian Peninsula into eastern Turkey also has induced a westward extrusion of Anatolia, the high central part of Turkey. Thus, the same processes active in eastern Asia have affected the landscape of the western portion but only on a smaller scale.

The evolution of the western segment of the Tethyan System is the most complicated, involving more than just a collision of the African (Africa) continent with parts of Europe. In Early Jurassic time (about 180,000,000 years ago), Africa, which then lay close to Europe, moved southeastward away from it. In doing so, it caused new ocean floor (Tethys) and new continental margins to form. Much of the rock in the Alps, for instance, was deposited on this newly formed margin of southern Europe. Later, during the Cretaceous (about 100,000,000 years ago), the divergence of Africa and Europe ceased, and convergence between them began. Mountain ranges through northern Greece (the Pindus (Pindus Mountains)), the Yugoslav region (the Dinaric Alps), Romania, Hungary, the Czech Republic, and Slovakia (the Carpathians (Carpathian Mountains)), and Austria, Switzerland, France, and Italy (the Alps) all formed as the Italian Peninsula—a promontory on the African continent—moved first north-northeast toward Europe at 20 to 30 millimetres per year and later northwest at a slower rate of about 10 millimetres per year. The change in the direction of motion and the irregular shape of this promontory are two reasons why the tectonic evolutions of the different ranges of Europe are very different from each other. This is unlike the situation in the Himalayas, where the history of the belt is similar throughout the 2,500-kilometre-long range.

The best-studied of these ranges is the western Alps in Switzerland and France. The western end of the Tethys Ocean floor was subducted beneath northern Italy until about 45,000,000 to 35,000,000 years ago. At that time, southern Europe and northern Italy collided. As the southern margin of Europe began to be subducted beneath northern Italy, the sedimentary cover deposited on the European margin of the Tethys Ocean was detached and scraped off the margin. Thick layers of relatively strong sedimentary rock (e.g., limestone and sandstone) that had been deposited on weak layers of salt (and in some cases shale) became detached and folded into huge nappes—enormous, flat layers of rock that seem to have been folded and sometimes dragged over one another like sheets of cloth pushed over a table or bed.

As northern Italy continued to override the coast of southern Europe, it not only pushed the sedimentary cover farther onto the European landmass, but it also scraped up bits and pieces of the deeper metamorphic rocks of Europe's basement. Moreover, as the crust thickened, the increase in pressure and temperature metamorphosed the deeply buried rocks. Although there are exceptions, the northern and western parts of the Alps thus are dominated by folded, unmetamorphosed sedimentary rock, and the southern part consists largely of metamorphic rock.

As Europe was flexed down under the weight of the Alps thrust onto it, a foreland basin (see below) formed just north of the Alps: this is the Molasse Basin of northern Switzerland and southern Germany. Continental convergence in the past 10,000,000 years has caused folding and thrusting in the Jura Mountains of northwest Switzerland and France, and displacement on ramp overthrusts beneath the front of the Alps has elevated several crystalline massifs, including the Belledonne and Mont Blanc (Blanc, Mont) massifs in France and the Aare (or Aar) and Gotthard massifs in Switzerland. Moreover, with the elevation of the Alps above the Po plain of northern Italy, a southward overthrusting has carried the southern part of the Alps back onto the basin there as the Italian promontory has continued to penetrate into the rest of Europe.

The Apennines (Apennine Range), which form the backbone of the Italian peninsula, were built by the folding and faulting of sedimentary rock deposited on the peninsula. The deformation in a direction nearly perpendicular to that of the Alps was due in part to a phase of the northeastward movement of Italy toward the Adriatic coast of the Balkan Peninsula and also to the rotation of Corsica and Sardinia away from southern France and toward Italy. Thus, while the crust of the Alps was being shortened in its north–south or northwest–southeast dimension, that of the Apennines was being shortened in its northeast–southwest dimension.

While the Alps, the Apennines, and the ranges of eastern Europe were being built, different processes created mountain ranges in parts of western Europe and destroyed others in eastern Europe. For instance, while the last remnant of the Tethys Ocean, the eastern Mediterranean Sea, continues to be subducted beneath Greece and Turkey, north–south crustal extension and associated crustal thinning occurs in the Aegean area and western Turkey. This crustal thinning has already lowered the surface of what may have been a high range or plateau to below the level of the Aegean Sea and is reducing the average elevation of western Turkey.

In contrast, the western Mediterranean Sea—between Italy, Spain, and North Africa—was formed during the past 30,000,000 years and is not a remnant of the Tethys Ocean. Since that time and concurrently with the subduction of the Tethys lithosphere beneath southern Italy, Greece, and Turkey, fragments of crust have separated from southern Europe. As these fragments drifted across the ancient westernmost end of the Tethys Ocean, they opened the new western Mediterranean basin behind them.

The collisions of these fragments with parts of Italy and Africa have contributed to the building of mountain ranges in these areas. Corsica and Sardinia swung out from southern France, and the eastern margin of Corsica, which lies below sea level, collided with Italy. The Calabrian (Calabria) peninsula of southern Italy once lay against Sardinia, but its southward drift opened the Tyrrhenian Sea. The volcanoes of Italy, including Mount Vesuvius (Vesuvius) near Naples and Mount Etna (Etna, Mount) on Sicily, were formed as a result of the subduction of the ancient oceanic lithosphere of the Tethys beneath the Calabrian arc, which only recently collided with the rest of Sicily and the southern part of the Italian peninsula. Small fragments farther west collided with North Africa, causing crustal shortening and mountain building across northern Tunisia, Algeria, and Morocco.

The convergence of another small fragment with Europe built the Pyrenees. The Iberian Peninsula lay against the western margin of France until about 90,000,000 or 100,000,000 years ago, when it began to rotate into its present position and opened the Bay of Biscay (Biscay, Bay of) behind it. As the peninsula moved toward southern France, a combination of crustal shortening and strike-slip deformation along the Pyrenees built the narrow range that separates Spain and France.

Some high areas follow old mountain belts, but the present elevations are the result of recent uplift due to the heating of the lithosphere and to its thermal expansion. Strictly speaking, these belts are not residual mountain ranges. The mountainous topography of Norway and northern Sweden (Scandinavia), for instance, follows an Early Paleozoic belt that marks the zone where Europe and North America (including Greenland) collided more than 400,000,000 years ago, long before the present Atlantic Ocean formed. The present-day topography, however, probably exists because this area was heated when Greenland was rifted away from Europe some 55,000,000 years ago when the North Atlantic Ocean began to form. Similarly, the mountains of eastern Australia, including the Snowy Mountains that contain the continent's highest peak, follow a Paleozoic belt; yet, the present topography seems to be the result of the warming of the lithosphere both when New Zealand separated from the east coast of Australia some 80,000,000 to 90,000,000 years ago and again when Australia drifted over a hot spot in the asthenosphere tens of millions of years ago.

Most of the highlands of continents that are not characterized by chains of mountains have resulted from a heating of the lithosphere. The majority of them, however, are better described as plateaus than as mountain ranges.