Planetary interiors

Planetary interiors characteristically have a nested shell structure, with concentric layers of differing physical properties and chemical compositions. Often, planetary materials are compressed into unusual forms by the tremendous pressures deep within planets.

Overview

The interior of Earth is typical in many ways of the interiors of other terrestrial, or “rocky,” planets and serves as a useful starting point for understanding those environments. Earth consists of three layers, differing sharply in composition and physical properties. These include a thin outermost layer called the crust, a thick intermediate layer called the mantle, and a dense central core.

The Earth’s crust consists of two very different materials: granitic rocks averaging about forty kilometers thick beneath the continents and basaltic rocks five to ten kilometers thick beneath the ocean basins. Sea waters fill the ocean basins because, here, the crust is thin and its elevation low. Nonetheless, there is no direct connection between oceans and oceanic crust. The crust of the Earth composes only about 0.3 percent of its mass and about 0.8 percent of its volume. Beneath the Earth’s crust is a dense mantle of iron and magnesium silicate rocks extending to a depth of about 2,885 kilometers. The mantle occupies about five-sixths of the volume of the Earth and accounts for two-thirds of its mass.

At the center of the Earth is a dense core consisting of two parts. The outer core, with a radius of 3,486 kilometers, is slightly more than half the diameter of the Earth and about the size of the planet Mars. This outer core is most likely molten iron and nickel with other elements present in smaller amounts. Fluid movements in the liquid outer core are believed to be the source of the Earth’s magnetic field. The inner core, which is solid but similar to the outer core in composition, has a radius of 1,216 kilometers. The core occupies about one-sixth of the volume of the Earth and accounts for a third of its mass. The inner core accounts for only 0.7 percent of the volume of the Earth and 1 percent of its mass.

The composition of Earth’s interior, as well as that of other planets, is dictated by the chemistry of the solar system. If the solar system formed from a cloud of material similar to the present sun in composition, any solid grains that condensed should have a composition similar to the sun, minus those elements that would have formed gases. In the inner solar system (above the freezing point of water), most hydrogen, nitrogen, and carbon would have formed gases. Oxygen would largely be tied up in water vapor, but much would form oxides with various metallic elements or combine with silicon to form silicate minerals. Some sulfur would be available for iron sulfide minerals such as troilite (FeS). Noble gases (helium, neon, argon, krypton, and xenon) combine rarely with other elements and would be nearly absent from any solid grains. The most abundant elements of the inner planets should be oxygen, magnesium, silicon, sulfur, and iron. Direct chemical analysis of solar system bodies is still largely limited to the Earth, Moon, meteorites, and more recently Mars.

Meteorites are particularly valuable because they provide samples of the materials that condensed in the early solar system. A class of meteorites called chondrites closely matches the theoretically predicted composition of the planets; they are believed to be samples of very primitive solar system meteorites. Chondrites are usually the starting point for studies of the chemical composition of the inner planets. The presumed composition of Earth’s mantle, mostly iron and magnesium silicates, agrees well with the theory that Earth has an overall chondritic composition. Much of the iron in Earth has settled into the central core, together with nickel and perhaps much of Earth’s sulfur.

Pressure deep within the Earth is caused by the weight of overlying rocks. It is easy to calculate once the distribution of mass within Earth is known. The gravitational attraction within the Earth is solely caused by the mass of the Earth between the observer and the core; the gravitational effect of a shell surrounding an observer is zero. In a uniform Earth, gravitational attraction would decrease steadily toward the center, but because the Earth’s core is dense, gravitational attraction actually increases slightly with depth and is about 4 percent higher at the boundary of the core than at the Earth’s surface. Pressure at the base of the Earth’s continental crust is about ten thousand times greater than Earth’s atmospheric pressure. Geologists commonly use a metric unit, the bar, to measure pressure. A bar is convenient because it is nearly equal to the pressure of the Earth’s atmosphere at sea level and thus is easily visualized. A bar is equal to 100,000 pascals. The pressure at the bottom of the mantle is about 1.4 million bars, and the pressure at the Earth’s center is about 3.6 million bars.

Temperatures deep inside the planet are not known as precisely as many other physical quantities. The Earth becomes hotter at an average rate of twenty-five kelvins per kilometer near the surface, reaching about 1,250 kelvins at the base of the continental crust. Thereafter, temperatures increase much more slowly, reaching three thousand to five thousand kelvins in the core. Despite high temperatures in the mantle, it is still solid, because high pressure elevates the melting point of rocks. Portions of the mantle melt at depths of fifty to one hundred kilometers only locally to generate magma for volcanic eruptions. Much of the Earth’s internal heat is caused by radiogenic heating, but some may be primordial remnants from the accretion of the Earth or gravitational energy released when dense material sank to create the core.

Higher temperatures in the Earth’s deep interior create convection, the rise of hot, light material and the sinking of cooler, denser material. Although the rocks in the mantle are very rigid over short time spans, they can flow slowly over long timescales. Crustal plates consist of the crust, plus part of the mantle beneath. The material of the plates, collectively called the lithosphere, is about one hundred kilometers thick.

The movement of crustal plates on the Earth results in recycling of oceanic crust over a period of about two hundred million years. New oceanic crust is formed as plates spread apart at the mid-ocean ridges, and old oceanic crust returns to the Earth’s interior at oceanic trenches, mostly around the Pacific Rim. As oceanic crust is drawn back into the Earth’s interior, it partially melts and molten material invades the Earth’s crust. This molten material is enriched in elements of largeelectric chargethat do not fit easily into the crystal structures of mantle rocks. During the Earth’s lifetime, the crust has become enriched in potassium, rubidium, silicon, uranium, thorium, and rare Earth elements. As a result the mantle has become depleted. Volatile materials such as water and carbon dioxide have also escaped to the surface.

Physical processes active in Earth’s interior may be at work in other planets as well: internal heating, recrystallization of rocks and other materials to denser forms at great depths, gravitational separation of a dense metallic or metallic sulfide core, internal flow caused by convection, and segregation of particular materials into a crust. These processes are all driven, directly or indirectly, by internal heat. Small planetary bodies have less internal heat than large ones. As the planets accreted from the impact of smaller bodies, impacts released heat. Smaller bodies had less mass to retain heat and therefore cooled more quickly than larger ones. Gravitational separation of a dense core would also release heat, but separation of a core is less likely in a body with weak gravity and a cool interior. Small bodies have smaller amounts of radioactive material for radiogenic heating and smaller thicknesses of insulating mass to retain the heat that is produced. Larger bodies of the inner solar system show considerably more internal activity than smaller ones.

Spectroscopic evidence indicates that some large asteroids have basaltic surfaces, which means that bodies a few hundred kilometers in diameter have undergone enough internal heating at some time in their history to melt rocks. The moon is believed to have a very small core, probably rich in iron or iron sulfide. Mercury is unusually dense for a small planet and probably has a core with a radius three-fourths that of the planet itself. Although both Mercury and the moon are covered with vast volcanic plains, neither has experienced significant volcanic or tectonic activity in the last three billion years, perhaps because their interiors are too cool and rigid to permit much convection or melting.

Venus is only slightly smaller than the Earth, and its interior is probably similar. The surface of Venus shows many manifestations of intense internal activity such as fracturing and folding of the crust and widespread volcanic activity. Nevertheless, Venus lacks the sharp division between two kinds of crust that is so evident on the Earth. Also, Venus lacks obvious topographic features typical of plate tectonics on the Earth: ridges with crustal spreading, fracture zones, or trenches. One possible reason for its different geology may be that Venus has a thin, flexible crust that results from its very high surface temperature (650 kelvins), rather than thick, rigid crustal plates, as on the Earth.

Mars is believed to have a small core of iron or iron sulfide and shows considerably more evidence of recent dynamic internal activity than Mercury or the moon. A gigantic rift valley, Valles Marineris, extends more than five thousand kilometers on Mars. The Tharsis Ridge is close by, crowned with several enormous volcanoes. The relative lack of craters on these features of Mars indicates that they formed much more recently than the ancient volcanic plains on Mercury or Venus, probably about one billion years ago. Nevertheless, the opposite side of Mars consists of ancient, cratered terrain that has been inactive since its early history. Mars never developed a global system of crustal plates as had Earth, probably because the planet cooled enough for a very thick, rigid lithosphere to form.

In the outer solar system, water ice and other volatile materials (collectively termed “ices”) are major constituents of planetary bodies. Water ice evaporates quickly in a vacuum even in a solid state. Only at 175 kelvins or a little sunward of Jupiter was it cold enough for ice to form in a vacuum and survive billions of years. At 150 kelvins, methane hydrate (a solid mixture of methane and water) condenses, followed by ammonia hydrate at 120 kelvins, and argon and pure methane at about sixty-five kelvins. These temperatures are comparable to the surface temperatures on the satellites of Jupiter (135 kelvins), Saturn (one hundred kelvins), Uranus (eighty-five kelvins), and Neptune (fifty kelvins).

Satellites of all the large outer planets, and perhaps the dwarf planet Pluto and its satellite Charon as well, consist of rocky cores surrounded by mantles of ices. One of the great surprises of planetary exploration has been the degree of internal activity that occurs on objects made largely of ices. A rocky core need not be very large to generate enough radiogenic heat to warm water ice to near or above the melting point, and some larger bodies, especially the large satellites of Jupiter, may have liquid water or icy slush within their mantles. A number of satellites in the outer solar system show evidence that ice has flowed upward from the interior to the surface. This process has been called “ice volcanism,” but it is more nearly akin to vertical glacial flow. Water ice undergoes a remarkable series of changes in crystal structure at high pressure, and these high-pressure forms of ice are present undoubtedly in the interiors of large satellites in the outer solar system.

The icy satellites of the outer solar system circle very massive planets and experience a force that is only of minor importance on Earth: tidal forces. Tidal forces arise because the gravitational attraction between two bodies is greater on their near sides than their far sides. Tidal forces distort bodies into ellipsoidal shapes whose long axes point toward the other attracting body. The moon exerts only a small effect on the solid Earth and a more noticeable effect on the oceans, but Earth’s tidal force has locked the moon’s rotation so that it always shows the same face to Earth. Tidal stresses also produce very small “moonquakes” when the Moon is nearest to Earth in its orbit.

Almost all satellites in the solar system are similarly locked to their planets. If a planet has several large satellites, the satellites also exert tidal forces that tend to twist the satellites into alignment with one another. Under the right conditions, the conflicting tidal forces can generate a significant amount of internal frictional heat. The most spectacular example is Io, Jupiter’s innermost satellite. Frictional heating from tidal forces makes Io’s interior hot enough to cause active volcanism, despite the fact that Io is smaller than Earth’s moon. The ultimate source of energy for this heat is the orbital motion of Jupiter’s satellites and the rotation of Jupiter. Heating of Io is but a negligible energy drain on these bodies. Other satellites, notably Europa, another satellite of Jupiter, the Saturnian satellite Enceladus, and several satellites of Uranus have extensive fracture networks also related to tidal stresses.

Jupiter, Saturn, Uranus, and Neptune, the so-called gas giants, probably formed by accumulating rock and ice cores several times as massive as Earth. These large cores were massive enough to attract and hold large amounts of gases. Jupiter and Saturn, which were sizable enough to retain essentially all of their gases, are similar to the sun in composition. Uranus and Neptune did not become large enough to attract or retain as much hydrogen and are made mostly of denser gases such as ammonia or methane. The deep interiors of all of these planets are very dense gases or fluids. In Jupiter and Saturn, pressures become great enough that hydrogen is believed to have metallic properties, a form of matter predicted in theory and only created in small systems in the laboratory. Fluid movements in the interiors of the gas giants are believed to account for their magnetic fields, in the same manner that fluid motions in Earth’s core produce a magnetic field.

Methods of Study

The deepest mine on the Earth is four kilometers deep, and the deepest well ever drilled is twelve kilometers deep, about one one-thousandth the diameter of the Earth. Direct sampling of the Earth’s interior has not penetrated through the crust, much less reached the core. Geologic processes have brought materials from the deep crust and upper mantle to the surface, from depths perhaps greater than one hundred kilometers. Knowledge of the Earth’s deep interior is derived from a variety of indirect techniques, therefore, and knowledge of the interiors of other planets is based in part on knowledge of the Earth’s interior.

Some information about Earth’s interior can be derived from the surface. Most rocks have densities of 2.7 to 3.0 grams per cubic centimeter. The size of Earth and hence its volume are known from mapping. Earth’s mass is known from its gravitational attraction. The average density of the Earth, 5.5 grams per cubic centimeter, implies that there is dense material deep in the interior. Another observation supports this conclusion. All rotating objects, including the Earth, can be described by a physical quantity called moment of inertia. Moment of inertia plays much the same role in equations of rotational motion that mass plays in equations of linear motion. The Earth’s moment of inertia can be found from observations of the gravitational effects of the moon on the Earth. It does not match the moment of inertia of a uniform body with the size and mass of the Earth; instead, the Earth’s moment of inertia indicates that much of the Earth’s mass is concentrated near its center. Similar studies applied to other planets also provide clues to their internal structure.

Samples of the Earth’s mantle are found in two geologic settings: as parts of ophiolites (fragments of oceanic crust incorporated into mountain ranges) and as inclusions in kimberlite pipes, volcano-like vents that appear to have brought rocks (and occasionally diamonds) from the mantle to the surface with great speed and violence. The composition of mantle rocks—predominantly iron and magnesium silicates—agrees with the theory that the Earth is largely composed of chondritic raw material and their mechanical properties match the properties of the upper mantle as deduced from seismic studies.

One measure of solar-system bodies can provide great insight into their composition, even for bodies not sampled directly: bulk density, which is the mass of a body divided by its volume. The volume of a body can be computed readily once the diameter is known, and diameter can be obtained with high precision with spacecraft imagery. The same spacecraft that obtains imagery can also provide an accurate mass determination, through the gravitational effect of the body on the path of the spacecraft. Bodies with bulk densities of three to four grams per cubic centimeter, such as the moon, Mars, and Io, are probably composed mostly of silicates of the type found in the Earth’s mantle.

The only solid that is abundant in the solar system and less dense than silicate rocks is water ice, with a density of 0.9 gram per cubic centimeter. Small solid bodies with densities of one to three grams per cubic centimeter are very likely made of varying proportions of silicate rock and water ice. Most of the satellites in the outer solar system are of this composition. The very large outer planets also have low bulk densities (Saturn has a bulk density of only 0.7 gram per cubic centimeter). These planets are known, from spectroscopic evidence as well as direct imaging by spacecraft, to be composed mostly of dense gases, probably with solid cores.

Context

Until seismic methods became available, the only way of knowing anything of the Earth’s interior was by analyzing its exterior properties, such as mass, shape, and gravitational attraction. During the eighteenth and early nineteenth centuries, accurate determination of the shape of the Earth was at the foremost frontiers of research in science. Many now-indispensable advanced mathematical methods were stimulated by study of the Earth’s shape and gravitation.

Seismic studies of the Earth’s interior began in the late nineteenth century. John Milne, an English engineer working in Japan, refined the crude seismographs then in existence to new heights of sensitivity. By the 1880s, seismographs were capable of detecting earthquakes at distances of thousands of kilometers. In 1896, Richard Dixon Oldham argued from astronomical and geologic evidence that Earth must have a dense iron core. By 1906, he presented evidence for a core by showing that seismic waves passing through the center of the Earth traveled more slowly than they did in the mantle. In 1909, the Yugoslavian seismologist Andrija Mohorovičić noted that seismic waves for nearby earthquakes showed an abrupt change in velocity beyond a certain distance, which he interpreted as evidence for a thin outer crust. By 1912, Beno Gutenberg had determined accurately the dimensions of the core. In 1936, the Danish seismologist Inge Lehmann presented evidence for the existence of the inner core. Until the 1920s, most seismographs were mechanical, curious blends of massive weights and delicate lever mechanisms to obtain enough sensitivity and magnification to record faint signals. Since that time, electronic signal amplification has vastly improved the sensitivity of seismographs.

One of the great conceptual advances in the understanding of the Earth’s interior was the hypothesis of continental drift, proposed by the German meteorologist Alfred Lothar Wegener in 1912. Beginning in the late 1950s, evidence began to accumulate that the active processes for crustal motion on Earth were actually concentrated in the ocean basins. Between 1960 and 1975, a modified concept of continental drift, or plate tectonics, became the generally accepted view among geologists.

In the 2020s the emphasis of scientific study began to change from an almost sole focus on understanding the Earth's composition to applying this knowledge toward understanding other celestial objects. These include other planets in our solar system, and even exoplanets in remote galaxies.

The study of the interiors of other planets was not possible until spacecraft could provide accurate dimensions and masses for these objects, as well as views of their surfaces that might provide clues to internal processes. The Apollo astronauts placed seismographs on the Moon between 1969 and 1972. Spent third stages of the Saturn V rocket and the discarded ascent stage of the Lunar Module were purposely impacted on the moon to create seismic events of known magnitude for surface experiments to detect. These impacts helped to calibrate the experiments, and also to decipher subsurface structure of the moon. Seismic waves caused by natural meteor impacts have also helped probe the lunar interior.

One of the most dramatic insights into planetary interiors was provided by American scientists Stanton Peale, Patrick Cassen, and Ray Reynolds. They calculated the amount of tidal heating that might occur on Io and found that Io could get hot enough to sustain volcanic activity. They published this hypothesis in the journal Science less than a month before the Voyager spacecraft returned the first photographs of volcanic eruptions on Io, the first active volcanism ever seen outside Earth.

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