Structure of Earth
The internal structure of Earth, structure of the solid Earth, or simply structure of Earth refers to concentric spherical layers subdividing the Solid earth, i.e., excluding Earth's atmosphere and hydrosphere. It consists of an outer silicate solid crust, a highly viscous asthenosphere and solid mantle, a liquid outer core whose flow generates the Earth's magnetic field, and a solid inner core.
Scientific understanding of the internal structure of Earth is based on observations of topography and bathymetry, observations of rock in outcrop, samples brought to the surface from greater depths by volcanoes or volcanic activity, analysis of the seismic waves that pass through Earth, measurements of the gravitational and magnetic fields of Earth, and experiments with crystalline solids at pressures and temperatures characteristic of Earth's deep interior.
The structure of Earth can be defined in two ways: by mechanical properties such as rheology, or chemically. Mechanically, it can be divided into lithosphere, asthenosphere, mesospheric mantle, outer core, and the inner core. Chemically, Earth can be divided into the crust, upper mantle, lower mantle, outer core, and inner core. The geologic component layers of Earth are at the following depths below the surface:
|Depth (km)||Chemical layer||Depth (km)||Mechanical layer||Depth (km)||PREM||Depth (km)||General layer|
|0–35†||Crust||0–80*||Lithosphere||0–10||0–80*||… Upper crust||Lithosphere||0–35†||Crust|
|10–20||… Lower crust|
|35–670||Upper mantle||… LID||35†-80*||Lithospheric mantle|
|400–600||... Transition zone||400–670||Transition zone|
|35–670||... Transition zone|
|35–670||600–670||... Transition zone|
|670–2,890||Lower mantle||220–2,890||Mesospheric mantle||670–770||Lower mantle||… Uppermost||670–2,890||Lower Mantle|
|2,740–2,890||... D″ layer|
|2,890–5,150||Outer core||2,890–5,150||Outer core||2,890–5,150||Outer core||2,890–5,150||Outer core|
|5,150–6,370||Inner core||5,150–6,370||Inner core||5,150–6,370||Inner core||5,150–6,370||Inner core|
|* Depth varies locally between 5 and 200 km.
† Depth varies locally between 5 and 70 km.
The layering of Earth has been inferred indirectly using the time of travel of refracted and reflected seismic waves created by earthquakes. The core does not allow shear waves to pass through it, while the speed of travel (seismic velocity) is different in other layers. The changes in seismic velocity between different layers causes refraction owing to Snell's law, like light bending as it passes through a prism. Likewise, reflections are caused by a large increase in seismic velocity and are similar to light reflecting from a mirror.
The Earth's crust ranges from 5–70 kilometres (3.1–43.5 mi) in depth and is the outermost layer. The thin parts are the oceanic crust, which underlie the ocean basins (5–10 km) and are composed of dense (mafic) iron magnesium silicate igneous rocks, like basalt. The thicker crust is continental crust, which is less dense and composed of (felsic) sodium potassium aluminium silicate rocks, like granite. The rocks of the crust fall into two major categories – sial and sima (Suess, 1831–1914). It is estimated that sima starts about 11 km below the Conrad discontinuity (a second order discontinuity). The uppermost mantle together with the crust constitutes the lithosphere. The crust-mantle boundary occurs as two physically different events. First, there is a discontinuity in the seismic velocity, which is most commonly known as the Mohorovičić discontinuity or Moho. The cause of the Moho is thought to be a change in rock composition from rocks containing plagioclase feldspar (above) to rocks that contain no feldspars (below). Second, in oceanic crust, there is a chemical discontinuity between ultramafic cumulates and tectonized harzburgites, which has been observed from deep parts of the oceanic crust that have been obducted onto the continental crust and preserved as ophiolite sequences.
Many rocks now making up Earth's crust formed less than 100 million (1×108) years ago; however, the oldest known mineral grains are about 4.4 billion (4.4×109) years old, indicating that Earth has had a solid crust for at least 4.4 billion years.
Earth's mantle extends to a depth of 2,890 km, making it the planet's thickest layer. The mantle is divided into upper and lower mantle separated by a transition zone. The lowest part of the mantle next to the core-mantle boundary is known as the D″ (D-double-prime) layer. The pressure at the bottom of the mantle is ≈140 GPa (1.4 Matm). The mantle is composed of silicate rocks richer in iron and magnesium than the overlying crust. Although solid, the mantle's extremely hot silicate material can flow over very long timescales. Convection of the mantle propels the motion of the tectonic plates in the crust. The source of heat that drives this motion is the primordial heat left over from the planet's formation renewed by the radioactive decay of uranium, thorium, and potassium in Earth's crust and mantle.
Due to increasing pressure deeper in the mantle, the lower part flows less easily, though chemical changes within the mantle may also be important. The viscosity of the mantle ranges between 1021 and 1024 Pa·s, increasing with depth. In comparison, the viscosity of water is approximately 10−3 Pa·s and that of pitch is 107 Pa·s.
Earth's outer core is a fluid layer about 2,400 km (1,500 mi) thick and composed of mostly iron and nickel that lies above Earth's solid inner core and below its mantle. Its outer boundary lies 2,890 km (1,800 mi) beneath Earth's surface. The transition between the inner core and outer core is located approximately 5,150 km (3,200 mi) beneath the Earth's surface. Earth's inner core is the innermost geologic layer of the planet Earth. It is primarily a solid ball with a radius of about 1,220 km (760 mi), which is about 20% of Earth's radius or 70% of the Moon's radius.
The average density of Earth is 5.515 g/cm3. Because the average density of surface material is only around 3.0 g/cm3, we must conclude that denser materials exist within Earth's core. This result has been known since the Schiehallion experiment, performed in the 1770s. Charles Hutton in his 1778 report concluded that the mean density of the Earth must be about that of surface rock, concluding that the interior of the Earth must be metallic. Hutton estimated this metallic portion to occupy some 65% of the diameter of the Earth. Hutton's estimate on the mean density of the Earth was still about 20% too low, at 4.5 g/cm3. Henry Cavendish in his torsion balance experiment of 1798 found a value of 5.45 g/cm3, within 1% of the modern value. Seismic measurements show that the core is divided into two parts, a "solid" inner core with a radius of ≈1,220 km and a liquid outer core extending beyond it to a radius of ≈3,400 km. The densities are between 9,900 and 12,200 kg/m3 in the outer core and 12,600–13,000 kg/m3 in the inner core.
The inner core was discovered in 1936 by Inge Lehmann and is generally believed to be composed primarily of iron and some nickel. Since this layer is able to transmit shear waves (transverse seismic waves), it must be solid. Experimental evidence has at times been inconsistent with current crystal models of the core. Other experimental studies show a discrepancy under high pressure: diamond anvil (static) studies at core pressures yield melting temperatures that are approximately 2000 K below those from shock laser (dynamic) studies. The laser studies create plasma, and the results are suggestive that constraining inner core conditions will depend on whether the inner core is a solid or is a plasma with the density of a solid. This is an area of active research.
In early stages of Earth's formation about 4.6 billion years ago, melting would have caused denser substances to sink toward the center in a process called planetary differentiation (see also the iron catastrophe), while less-dense materials would have migrated to the crust. The core is thus believed to largely be composed of iron (80%), along with nickel and one or more light elements, whereas other dense elements, such as lead and uranium, either are too rare to be significant or tend to bind to lighter elements and thus remain in the crust (see felsic materials). Some have argued that the inner core may be in the form of a single iron crystal.
Under laboratory conditions a sample of iron–nickel alloy was subjected to the corelike pressures by gripping it in a vise between 2 diamond tips (diamond anvil cell), and then heating to approximately 4000 K. The sample was observed with x-rays, and strongly supported the theory that Earth's inner core was made of giant crystals running north to south.
The liquid outer core surrounds the inner core and is believed to be composed of iron mixed with nickel and trace amounts of lighter elements.
The composition of the Earth bears strong similarities to that of certain chondrite meteorites, and even to some elements in the outer portion of the Sun. Beginning as early as 1940, scientists, including Francis Birch, built geophysics upon the premise that Earth is like ordinary chondrites, the most common type of meteorite observed impacting Earth. This ignores the less abundant enstatite chondrites, which formed under extremely limited available oxygen, leading to certain normally oxyphile elements existing either partially or wholly in the alloy portion that corresponds to the core of Earth.
Dynamo theory suggests that convection in the outer core, combined with the Coriolis effect, gives rise to Earth's magnetic field. The solid inner core is too hot to hold a permanent magnetic field (see Curie temperature) but probably acts to stabilize the magnetic field generated by the liquid outer core. The average magnetic field in Earth's outer core is estimated to measure 25 Gauss (2.5 mT), 50 times stronger than the magnetic field at the surface.
Recent evidence has suggested that the inner core of Earth may rotate slightly faster than the rest of the planet; in 2005 a team of geophysicists estimated that Earth's inner core rotates approximately 0.3 to 0.5 degrees per year faster. However, more recent studies in 2011[which?] did not support this hypothesis. Other possible motions of the core be oscillatory or chaotic.
The current scientific explanation for Earth's temperature gradient is a combination of heat left over from the planet's initial formation, decay of radioactive elements, and freezing of the inner core.
The force exerted by Earth's gravity can be used to calculate its mass. Astronomers can also calculate Earth's mass by observing the motion of orbiting satellites. Earth's average density can be determined through gravimetric experiments, which have historically involved pendulums. The mass of Earth is about 6×1024 kg.
- Geological history of Earth
- Lehmann discontinuity
- Rain-out model
- Travel to the Earth's center
- Solid earth
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