Earth structure The interior of the Earth is inaccessible: all that we know about it has had to be deduced from our knowledge of the rocks accessible to us and from the shape and physical properties of the Earth as a whole, and from the results of geophysical and geochemical experiment.

Our desire to understand the Earth goes far back into history, to the Hebrews, the Ancient Greeks, and the Chinese, but our understanding of the internal structure of the Earth really starts with the Victorian physicists (e.g. Rayleigh and Rutherford) and then Sir Harold Jeffreys with his classic book The Earth, first published in 1924, who laid the foundations for modern geoscience. With the very rapid advances in equipment, and particularly in computer technology, our knowledge of the details of the fine structure and workings of the Earth's interior has improved greatly since the 1980s.

The Earth is, in the broadest sense, a series of concentric spherical shells, each shell having distinct physical or chemical properties (Fig. 1, Table 1). The outermost, and thinnest, shell is the crust. Then, descending into the interior, the next shell is the mantle, which extends to a depth of 2891 km. This is subdivided into two: the upper mantle and the lower mantle. Finally at the centre of the Earth is the core. This also is subdivided into two: the outer core and, the innermost sphere, the inner core.

Table 1 Volume, mass, and density of the Earth

	Depth	Volume	Volume	Mass	Mass	Density
(km)	(1018m3)	(% of total)	(1021kg)	(% of total)	(103kg m−3)
Crust	0–Moho	10	0.9	28	0.5	2.60–2.90
Upper mantle	Moho–670	297	27.4	1064	17.8	3.35–3.99
Lower mantle	670–2891	600	55.4	2940	49.2	4.38–5.56
Outer core	2891–5150	169	15.6	1841	30.8	9.90–12.16
Inner core	5150–6371	8	0.7	102	1.7	12.76–13.08
Whole Earth	–	1083	100	5975	100

The crust

The rocks exposed at the surface of the Earth are part of the crust. The crust is a thin layer of silica-rich rocks which have been derived from the underlying mantle by melting and subsequent metamorphic or erosional processes,—or, in places, both. Crustal rocks can thus be broadly classified as igneous, metamorphic, or sedimentary, according to their individual histories. The crust is comprised of rocks of all ages from the oldest Hadean (more than 4000 Ma old) to the youngest modern lavas.

The crust that makes up the continents differs in origin, structure, and composition from the crust beneath the oceans. On average the continental crust is of granitoid composition, whereas the oceanic crust is basaltic. Beneath the continents the crust has an average thickness of about 38 km. However, in areas of continental extension, such as the Rhine Graben in Germany and the East African Rift, the crust is locally thinner by 10–15 km than the crust on either side of the rift zone. Beneath features such as the Andes, the Alps, the Himalayas, or Tibet, the continental crust is much thicker than normal, often far exceeding 50 km. In some respects the continental crust can be regarded as the light ‘scum’ that floats on the Earth's denser mantle, in the same way that icebergs or logs float on water. The total volume of continental crust has increased through Earth history. The regions termed ‘cratons’ which now form the centres of several of the continents (e.g. North America) have crust dating from the Archaean (pre-2500 Ma). Younger rocks then surround and overlie these ancient cratons as new material has been accreted to the continent over time. Models of ‘crustal growth rate’ (which are based on isotopic ratios) indicate that the continental crust formed gradually through much of the Archaean with an increased growth rate in the late Archaean, since when there has been a gradual increase. Over 70 per cent of the present surficial area of the continents was formed more than 450 Ma ago. The continual processes of erosion and deposition of sediments mean that a good deal of ‘recycling’ takes place in crustal rocks. There is a continuing loss of some sediment into the mantle at subduction zones, although the ‘conveyor-belt’ system of plate tectonics ensures that most sediment is added to the accretionary wedge of the overriding plate. New crust in the form of volcanic rocks derived directly from partial melting of the mantle is continually added to the continents.

In contrast to the continental crust, the oceanic crust is young, thin, and chemically magnesium-rich. All the oceanic crust has been formed since the Jurassic, and only fragments of mid-Jurassic crust remain. The average thickness of the oceanic crust is 7 km. Oceanic crust is formed as a result of decompression melting in the mantle at shallow depths beneath the mid-ocean ridges. As a result, the oceanic crust is basaltic and is uniform in composition. Oceanic basalts are generally termed MORB (Mid-Ocean Ridge Basalts). Some of the rising magma that forms the oceanic crust is erupted at the sea bed, but much more solidifies without erupting, to form the characteristic stratified layering of the oceanic crust. The broad lithological layering is inferred from seismic P-wave velocities. Layers 2 and 3 are commonly subdivided in terms of the details of lithology and physical properties (Table 2).

The uppermost parts of the crust have been sampled directly by drilling. The international Ocean Drilling Program (ODP), a major co-operative programme of drilling in oceanic regions, has provided detailed information on the fine structure of the oceanic crust and has answered many questions about the details of the formation of oceanic regions. On the continents there are just two deep boreholes that penetrate to mid-crustal levels: one in Germany (KTB) and the other on the Kola peninsula in Russia. Despite the scarcity of drill information, a variety of geophysical techniques are used to great advantage to determine the gross overall structure of the continental crust in different tectonic regions, as well as some of its fine structure. Gravity surveys enable models of possible underlying density structures to be established; and as the density of rocks is broadly dependent upon their composition, gravity measurements can be used to infer lithology. Electrical and magnetic surveys enable models of the electrical and magnetic properties of the crust and uppermost mantle to be determined. Mineral composition, porosity, and permeability are additional factors controlling the electrical conductivity and magnetic susceptibility of rocks. However, seismic methods provide the most detailed and unique images of the structure of the crust (both continental and oceanic).

Table 2 Gross layering of the oceanic crust

	Lithology	Average thickness	Seismic
(km)	P-wave
velocity
(km s−1)
Layer 1	sediments	about 0.5	about 2
(variable)	(depends upon age)
Layer 2	basalts	2.1±0.6	2.5–6.6
fractured lavas
Layer 3	dykes, gabbro,	5.0±0.8	6.6–7.6
cumulates

Seismic reflection profiling can yield images of the fine structure over small areas. It is the mainstay of the oil industry's search for hydrocarbons in sedimentary rocks. The method is used to image the structure of the crystalline crust and on occasion even the uppermost mantle. National programmes, such as COCORP in the USA, Lithoprobe in Canada, and BIRPS in the UK, have been very successful. Seismic reflection profiling provides images of any structures or features that send back a reflection. Reflections arise only when there is a change in seismic velocity or density. Reflections should thus be recorded from a sediment–igneous contact or from the base of a pluton, but totally homogenous material, such as a salt diapir or the interior of a granite pluton, will be reflection-free. Wide-angle seismic reflection profiling and seismic refraction are the other controlled-source seismic methods that are used to image the crust and uppermost mantle. While these methods do not generally provide such fine detail as reflection profiling, they have the significant advantage that they yield the seismic velocity structure. The velocity–depth structural models can then be interpreted in terms of lithology. The continental crust is not stratified like the oceanic crust and so does not have a characteristic seismic velocity structure. Nevertheless, the uppermost 10 km or so of the crystalline crust beneath the sedimentary cover generally has a P-wave velocity of 6.0–6.3 km s⁻¹, and beneath that the velocity is normally in excess of 6.5 km s⁻¹. According to the tectonic history of an area and its complexity, there may be low-velocity zones in the crust, or lower crustal material with a velocity in excess of 7 km s⁻¹.

The boundary between the crust and the underlying mantle is called the Mohorovicićdiscontinuity (abbreviated to Moho). It is named after Andrya Mohorovicić, who first delineated the boundary in 1909. The normal seismic P-wave velocity of the uppermost mantle is 8.1 km s⁻¹, but considerable variation is observed locally. In regions where the mantle is hotter, such as along the axis of the mid-ocean ridge system, the seismic velocity is reduced. In contrast, cold dense regions can have elevated seismic velocities.

The mantle and deep interior

The controlled-source seismic methods used to determine crustal and shallow mantle structures are not suitable for determining seismic velocities deep within the Earth. Instead, methods utilizing earthquakes as the energy source and networks of seismic recording stations are used to calculate the travel times of seismic waves. These travel times are then used to calculate the variation of seismic velocities with depth in the Earth. Seismic velocities in the mantle are also determined by using the dispersion of surface waves (i.e. the change in the character of the wave with increasing distance, which is due to the fact that waves of different frequency travel with different velocities). Figure 2 shows the seismic velocity structure for the whole Earth. This was determined using travel-time data, as well as the periods of the Earth's free oscillations and its mass and moment of inertia.

The Earth can also be classified by the way in which heat is transferred through it. The lithospheric plates are the Earth's outermost near-rigid, cool ‘skin’. These are the plates that move about on the Earth's surface and along whose edges much of the seismicity and volcanism occur. Conduction is the main mechanism of heat transfer through the lithosphere. As the plates are about 100 km thick, the lithosphere comprises both the crust and the uppermost part of the mantle. The mantle beneath the lithosphere is hotter and, although behaving as a solid on a short timescale, is able to flow on a geological timescale. This means that convection is the mechanism for heat transfer through the sub-lithospheric mantle. There has been considerable scientific debates as to whether or not the upper and lower mantle convect as two separate systems. That they may be separate is suggested by four lines of argument: (a) the increase in seismic velocities and density at 670 km; (b) the maximum depth of earthquakes at subduction zones is 670 km; (c) some subducting plates break off at 670 km; and (d) some geochemical models suggest that the upper mantle is depleted (e.g. by partial melting and extraction of crust) and has been separate from the lower mantle for much of the Earth's history. Detailed three-dimensional seismic images of the velocity structures in the mantle indicate, however, that some plates are subducted into the lower mantle, thus indicating that the mantle may not be completely stratified.

Both the seismic P-wave and S-wave velocities increase with depth through the mantle. The P-wave velocity increases from a value of 8.1 km s⁻¹ at the top of the mantle to 13.7 km s⁻¹ at the core–mantle boundary (CMB). There are, however, several major irregularities superimposed on the steady increase of both seismic velocities and density with depth. There is a low-velocity zone for S-waves in the upper mantle which extends to depths of approximately 220 km. This low-velocity zone, which has been well defined by surface wave dispersion data, is generally known as the asthenosphere (from asthenia, Greek ‘weak’ or ‘sick’). Beneath the asthenosphere, the seismic velocities and density increase steadily down to 400 km depth. There, and again at 670 km depth, all increase sharply. The velocities then increase steadily though the lower mantle to the basal 200 km of the mantle, where the rate of increase is much reduced. This is also a region that produces increased scatter in the amplitude and travel times of seismic waves, indicating that it is a zone of considerable inhomogeneity.

Experimental work on olivine, (Mg, Fe)₂SiO₄, a major constituent mineral of the mantle, has shown that its atoms undergo a phase change at pressures equivalent to depths of about 400 and 670 km. Phase changes like these do not involve any change in mineral composition; rather the atoms are reorganized into more closely packed crystalline lattice structures. Between about 390 and 450 km, the olivine is changed into a spinel structure, which results in a 10 per cent density increase. This change from olivine to spinel is exothermic, which means that the change involves a release of energy in the form of heat. The other major mantle minerals, the pyroxenes, also undergo a phase change at these depths, to garnet. The change that occurs at about 670 km is from spinel to post-spinel structures, perovskite and magnesium oxide. This change also results in a density increase of 10 per cent, but it is endothermic, which means that heat is required for the change to take place. It is thought that these phase changes may control the cessation of seismicity at 670 km depth as well as maintaining the pattern of convection in the mantle as two separate systems: upper mantle and lower mantle. The convection regime may undergo periodic flushing events when large volumes of accumulated upper mantle material descend from the base of the upper mantle into the lower mantle. The base of the lower mantle may in effect be a graveyard for subducted plates.

Some old terminology remains in describing the mantle. The whole region between 400 and 670 km depth is called the mantle transition zone. Beneath this, from 670 to 2700 km depth, the mantle is sometimes referred to as the D9 shell. The basal 200 km of the mantle (from 2700 to 2900 km depth) is called the D0 shell. The reduction of the seismic velocity gradient in the D0 shell, and its heterogeneity, may be due to the shell acting as a thermal boundary layer through which heat is conducted rather than convected as it is through the rest of the mantle. It may also be in part caused by vigorous chemical interaction between the silicate mantle and the iron core.

The core

The core was discovered by R. D. Oldham in 1906 and was accurately delineated as being at 2900 km depth by Beno Gutenberg in 1912. The core–mantle boundary, or CMB, is also known as the Gutenberg discontinuity. The core is physically and chemically distinct from the mantle. In composition it is predominantly iron with small amounts of other elements. Work on tides enabled Sir Harold Jeffreys to establish in 1926 that the outer core must be fluid. A decade later, in 1936, Inge Lehmann (1888–1993) was able to show that there was a solid inner core at the very centre of the Earth. She did this by using seismic energy from an earthquake in New Zealand that was recorded in Europe after having passed through the centre of the Earth. The outer core–inner core boundary is called the Lehmann discontinuity in her honour.

At the CMB the P-wave velocity drops from 13.71 km s⁻¹ to 8.06 km s⁻¹ and the S-wave velocity drops from 7.26 km s⁻¹ to zero, while the density increases from 5567 to 9903 kg m⁻³. This is consistent with the outer core being liquid (S-waves cannot be transmitted through a liquid). Within the outer core the P-wave velocity increases steadily, reaching 10.36 km s⁻¹ at the outer core–inner core boundary. At that boundary the P-wave velocity increases from 10.36 to 11.03 km s⁻¹, the S-wave velocity from zero to 3.50 km s⁻¹, and the density from 12 166 to 12 764 kg m⁻³. The P-wave velocity, S-wave velocity, and density then increase only slightly through the inner core, reaching values of 11.26 km s⁻¹, 3.66 km s⁻¹, and 13 088 kg m⁻³ respectively at the centre of the Earth.

The composition of the core is hard to verify: there are no core samples to be studied. Instead we have to rely on ingenuity and analogue. The relative abundances of elements in the Sun and in meteorites indicate that the core should be predominantly iron, and in bulk could approximate to Fe₂O with a small proportion of nickel. The seismic velocity and density structure, together with experiments conducted in the laboratory at high pressure and temperature, imply that the inner core may be almost pure iron. The outer core is an iron alloy with about 10 per cent of lighter elements, the most likely candidates being oxygen, sulphur, nickel, and silicon. Experiments have shown that liquid iron and iron alloys react strongly with solid iron and magnesium silicates. The seismic complexity of the CMB can thus be explained in terms of the chemical reactions taking place there. This may be the most chemically active part of the planet.

That the outer core is liquid and the inner core solid is consistent with all seismological observations as well as studies on tides and the Earth's rotation, which both require a liquid core. The liquid outer core is the source of the Earth's magnetic field. It acts as a giant spherical dynamo, in which less dense rising convection currents of liquid iron also carry electric currents. The interaction of these electric currents with the Earth's magnetic field then results in an enhancement of that magnetic field. This is called a self-exciting dynamo, and can occur in the Earth only because the outer core is liquid, convecting, and, being iron-rich, conducts electricity.

C. Mary R. Fowler

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