Until the mid-1960s, although geologists had assembled much detailed information on geological processes both on continents and in the ocean, there was no comprehensive theory to explain the interrelationship between these processes. It is no exaggeration to say that with the development of plate tectonics, the subject of geology was revolutionized. It is now possible to understand and explain such apparently widely divergent subjects as the location of volcanoes, where and why earthquakes occur and why mountain belts are located only in certain confined areas. We now understand the structure of the ocean including why the deepest part of the ocean is near the continents and not, as was once thought, in the centre of the ocean and why the crust under the ocean is much younger than most of the rocks on the continents and are made almost exclusively of basalt. Many of the patterns seen in the evolution of fossils can now be explained and it has even been possible to explain why the mammals of Australasia are dominated by marsupials; the Americas have a few marsupials and Africa and Asia have none at all. The purpose of this chapter is to explain our present knowledge of the structure of the Earth, including both its interior and its outer layers, which we now know are divided into plates. The history of the development of global plate tectonic theory will be explained and the chapter concludes with a section explaining the location and development of many of the major features of the continents as we see them today. An understanding of plate tectonics is fundamental to understanding a wide range of other subjects in physical geography including biogeography, oceans and global landform development.
The Earth’s structure
The Earth is roughly spherical, as even some of the ancient Greeks realized. It is now known that it is an oblate spheroid, somewhat flattened at the poles with a radius of 6357 km at the poles and 6378 km at the equator. The flattening of the Earth is very small but the shape of the Earth is very close to that expected of a fluid rotating with the same velocity that the Earth travels through space. This suggests that, over a long time span, the Earth behaves in some ways like a fluid.
The interior of the Earth
If the Earth was cut in half and the interior structure exposed, At the centre of the planet is a core that has a radius of 3470 km. The inner core is solid and very dense (∼13 g cm3 ) with a thickness of 1390 km. It is magnetized and has a temperature probe in the region of 3000°C. The solid inner core is surrounded by a transition zone about 700 km thick, which in turn is surrounded by a 1380 km thick layer of liquid material that together form the outer core. It is somewhat less hot than the inner core and also less dense (∼11.5 g cm3 ). The core is thought to be made up of iron and nickel with a small amount of lighter elements such as silicon. The pressure at the core of the Earth is up to 3 million times the Earth’s atmospheric pressure. The next layer is called the mantle and contains the largest mass of any of the layers of the Earth, approximately 70% of the total mass. The outer part (180 km) of the mantle along with the overlying crust is called the lithosphere and is rigid; this floats on the more mobile asthenosphere. Beneath the asthenosphere is a transition zone (350–700 km deep) to the lower mantle, which is located from 700 km to 2900 km below the surface. The mantle is composed principally of magnesium–iron silicates and has a density of 3.35 g cm – 3. When the rocks in the mantle are subjected to differential pressure and heat, they flow slowly, rather like ice flows in a glacier.
The outer layers of the Earth
The Earth’s outermost layer is the cold, rigid, yet thin layer that we are directly familiar with, known as the crust. The large continental land masses are formed primarily of granite-type rocks, which have a high content of aluminium and silica. Quartz and sodium-rich feldspar are the two principal minerals present. The crust is comparatively thick (∼35970 km) and has a density of 2.892.9 g cm – 3. By contrast, the rocks underlying the oceans are mainly basalt, which is lower in silica and higher in iron and magnesium.
The principal minerals in basalt are olivine, pyroxenes and feldspars rich in calcium. The oceanic crust is much thinner than the continental crust, being only 6–10 km thick, but has a higher density (∼3 g cm – 3 ). The boundary between the crust and the mantle is called the Mohorovičić discontinuity, named after its Czech discoverer, and usually called simply the Moho. At one time it was thought that the Moho was the layer where the Earth’s rigid crust moved relative to the mantle. However, it is now thought that the crust and the upper mantle together form the rigid upper layer known as the lithosphere (60–100 km thick) which floats on the asthenosphere and moves as the Earth’s plates.
The difference between the average elevation of the continents and the oceans is determined principally by differences in the thickness and density of continental and oceanic crust. This is the principle of isostasy. A common analogy used to describe isostasy is that of an iceberg. The top of an iceberg is above sea level and is supported by the buoyancy of the displaced water below the surface. The deeper an iceberg extends below the surface, the higher the same iceberg reaches above the surface. Isostasy is, however, a dynamic situation.
For example, when the major ice sheets, which until 10 000 years ago covered Europe and North America, melted a significant weight was removed from the continents. These continents are still rebounding upwards as a result of slow flow processes in the upper mantle. An extreme example of this is in areas of west–central Sweden that are still moving upwards at a rate of 2 cm yr-1. Areas can remain elevated where tectonic plates are colliding, as the crust thickens where two continental plates come together. This has occurred in the Tibetan Plateau, for example. It can also occur in areas of volcanic activity such as Yellowstone National Park, USA.