Great Oxygenation Event

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O₂ build-up in the Earth's atmosphere. Red and green lines represent the range of the estimates while time is measured in billions of years ago (Ga).
 

Stage 1 (3.85–2.45 Ga): Practically no O₂ in the atmosphere. The oceans were also largely anoxic with the possible exception of O₂ in the shallow oceans.
 

Stage 2 (2.45–1.85 Ga): O₂ produced, rising to values of 0.02 and 0.04 atm, but absorbed in oceans and seabed rock.
 

Stage 3 (1.85–0.85 Ga): O₂ starts to gas out of the oceans, but is absorbed by land surfaces. No significant change in terms of oxygen level.
 

Stages 4 and 5 (0.85–present): Other O₂ reservoirs filled; gas accumulates in atmosphere.

The Great Oxygenation Event, the beginning of which is commonly known in scientific media as the Great Oxidation Event (GOE, also called the Oxygen Catastrophe, Oxygen Crisis, Oxygen Holocaust, Oxygen Revolution, or Great Oxidation) was the biologically induced appearance of dioxygen (O₂) in Earth's atmosphere. Geological, isotopic, and chemical evidence suggests that this major environmental change happened around 2.45 billion years ago (2.45 Ga), during the Siderian period, at the beginning of the Proterozoic eon. The causes of the event remain unclear. As of 2016, the geochemical and biomarker evidence for the development of oxygenic photosynthesis before the Great Oxidation Event has been mostly inconclusive.

Oceanic cyanobacteria, which evolved into coordinated (but not multicellular or even colonial) macroscopic forms more than 2.3 billion years ago (approximately 200 million years before the GOE), are believed to have become the first microbes to produce oxygen by photosynthesis. Before the GOE, any free oxygen they produced was chemically captured by dissolved iron or by organic matter. The GOE started when oxygen produced by the cyanobacteria started escaping into the atmosphere, after other oxygen reservoirs were filled.

Cyanobacteria: responsible for the build-up of oxygen in the Earth's atmosphere

The increased production of oxygen set Earth's original atmosphere off-balance. Free oxygen is toxic to obligate anaerobic organisms, and the rising concentrations may have destroyed most such organisms at the time.

A spike in chromium contained in ancient rock-deposits formed underwater shows the accumulation had been washed off from the continental shelves. Chromium is not easily dissolved and its release from rocks would have required the presence of a powerful acid. One such acid, sulfuric acid (H₂SO₄), might have formed through bacterial reactions with pyrite. Mats of oxygen-producing cyanobacteria can produce a thin layer, one or two millimeters thick, of oxygenated water in an otherwise anoxic environment even under thick ice; before oxygen started accumulating in the atmosphere, these organisms would already have adapted to oxygen. Additionally, the free oxygen would have reacted with atmospheric methane, a greenhouse gas, greatly reducing its concentration and triggering the Huronian glaciation, possibly the longest episode of glaciation in Earth's history and called "snowball Earth".

Eventually, the evolution of aerobic organisms that consumed oxygen established an equilibrium in its availability. Free oxygen has been an important constituent of the atmosphere ever since.

Timing

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The most widely accepted chronology of the Great Oxygenation Event suggests that free oxygen was first produced by prokaryotic and then later eukaryotic organisms that carried out photosynthesis more efficiently, producing oxygen as a waste product. The first oxygen-producing organisms arose long before the GOE, perhaps as early as 3,400 million years ago.

Initially, the oxygen they produced would have quickly been removed from the atmosphere by the chemical weathering of reducing (oxidizable) minerals, most notably iron. This 'mass rusting' led to the deposition of iron(III) oxide in the form of banded-iron formations such as the sediments in Minnesota and Pilbara, Western Australia. The saturation of these mineral sinks, and the resulting persistence of oxygen in the atmosphere, led within 50 million years to the start of the GOE. Oxygen could have accumulated very rapidly: at today's rates of photosynthesis (much greater than those in the Precambrian without land plants), modern atmospheric O₂ levels could be produced in only 2,000 years.

Another hypothesis is that oxygen producers did not evolve until a few million years before the major rise in atmospheric oxygen concentration. This is based on a particular interpretation of a supposed oxygen indicator used in previous studies, the mass-independent fractionation of sulfur isotopes. This hypothesis would eliminate the need to explain a lag in time between the evolution of oxyphotosynthetic microbes and the rise in free oxygen.

In either case, oxygen did eventually accumulate in the atmosphere, with two major consequences.
Firstly, it oxidized atmospheric methane (a strong greenhouse gas) to carbon dioxide (a weaker one) and water. This decreased the greenhouse effect of the Earth's atmosphere, causing planetary cooling, and triggered the Huronian glaciation. Starting around 2.4 billion years ago, this lasted 300-400 million years, and may have been the longest ever snowball Earth episode.

Secondly, the increased oxygen concentrations provided a new opportunity for biological diversification, as well as tremendous changes in the nature of chemical interactions between rocks, sand, clay, and other geological substrates and the Earth's air, oceans, and other surface waters. Despite the natural recycling of organic matter, life had remained energetically limited until the widespread availability of oxygen. This breakthrough in metabolic evolution greatly increased the free energy available to living organisms, with global environmental impacts. For example, mitochondria evolved after the GOE, giving organisms the energy to exploit new, more complex morphologies interacting in increasingly complex ecosystems.

Timeline of glaciations, shown in blue.

Time lag theory

There may have been a gap of up to 900 million years between the start of photosynthetic oxygen production and the geologically rapid increase in atmospheric oxygen about 2.5–2.4 billion years ago. Several hypotheses propose to explain this time lag.

Tectonic trigger

2.1 billion year old rock showing banded iron formation

The oxygen increase had to await tectonically driven changes in the Earth, including the appearance of shelf seas, where reduced organic carbon could reach the sediments and be buried. The newly produced oxygen was first consumed in various chemical reactions in the oceans, primarily with iron. Evidence is found in older rocks that contain massive banded iron formations apparently laid down as this iron and oxygen first combined; most present-day iron ore lies in these deposits. Evidence suggests oxygen levels spiked each time smaller land masses collided to form a super-continent. Tectonic pressure thrust up mountain chains, which eroded to release nutrients into the ocean to feed photosynthetic cyanobacteria.

Nickel famine

Early chemosynthetic organisms likely produced methane, an important trap for molecular oxygen, since methane readily oxidizes to carbon dioxide (CO₂) and water in the presence of UV radiation. Modern methanogens require nickel as an enzyme cofactor. As the Earth's crust cooled and the supply of volcanic nickel dwindled, oxygen-producing algae began to out-perform methane producers, and the oxygen percentage of the atmosphere steadily increased. From 2.7 to 2.4 billion years ago, the rate of deposition of nickel declined steadily from a level 400 times today's.

Bistability

Another hypothesis posits a model of the atmosphere that exhibits bistability: two steady states of oxygen concentration. The state of stable low oxygen concentration (0.02%) experiences a high rate of methane oxidation. If some event raises oxygen levels beyond a moderate threshold, the formation of an ozone layer shields UV rays and decreases methane oxidation, raising oxygen further to a stable state of 21% or more. The Great Oxygenation Event can then be understood as a transition from the lower to the upper steady states.

Hydrogen gas

Another theory credits the appearance of cyanobacteria with suppressing hydrogen gas and increasing oxygen.

Some bacteria in the early oceans could separate water into hydrogen and oxygen. Under the Sun's rays, hydrogen molecules were incorporated into organic compounds, with oxygen as a by-product. If the hydrogen-heavy compounds were buried, it would have allowed oxygen to accumulate in the atmosphere.

However, in 2001 scientists realized that the hydrogen would instead escape into space through a process called methane photolysis, in which methane releases its hydrogen in a reaction with oxygen. This could explain why the early Earth stayed warm enough to sustain oxygen-producing lifeforms.

Late evolution of oxy-photosynthesis theory

The oxygen indicator might have been misinterpreted. During the proposed lag era in the previous theory, there was a change in sediments from mass-independently fractionated (MIF) sulfur to mass-dependently fractionated (MDF) sulfur. This was assumed to show the appearance of oxygen in the atmosphere, since oxygen would have prevented the photolysis of sulfur dioxide, which causes MIF. However, the change from MIF to MDF of sulfur isotopes may instead have been caused by an increase in glacial weathering, or the homogenization of the marine sulfur pool as a result of an increased thermal gradient during the Huronian glaciation period (which in this interpretation was not caused by oxygenation).

Role in mineral diversification

The Great Oxygenation Event triggered an explosive growth in the diversity of minerals, with many elements occurring in one or more oxidized forms near the Earth's surface. It is estimated that the GOE was directly responsible for more than 2,500 of the total of about 4,500 minerals found on Earth today. Most of these new minerals were formed as hydrated and oxidized forms due to dynamic mantle and crust processes.

Origin of eukaryotes

It has been proposed that a local rise in oxygen levels due to cyanobacterial photosynthesis in ancient microenvironments was highly toxic to the surrounding biota, and that this selective pressure drove the evolutionary transformation of an archaeal lineage into the first eukaryotes. Oxidative stress involving production of reactive oxygen species (ROS) might have acted in synergy with other environmental stresses (such as ultraviolet radiation and/or desiccation) to drive selection in an early archaeal lineage towards eukaryosis. This archaeal ancestor may already have had DNA repair mechanisms based on DNA pairing and recombination and possibly some kind of cell fusion mechanism. The detrimental effects of internal ROS (produced by endosymbiont proto-mitochondria) on the archaeal genome could have promoted the evolution of meiotic sex from these humble beginnings. Selective pressure for efficient DNA repair of oxidative DNA damages may have driven the evolution of eukaryotic sex involving such features as cell-cell fusions, cytoskeleton-mediated chromosome movements and emergence of the nuclear membrane. Thus the evolution of eukaryotic sex and eukaryogenesis were likely inseparable processes that evolved in large part to facilitate DNA repair. Constant pressure of endogenous ROS has been proposed to explain the ubiquitous maintenance of meiotic sex in eukaryotes.

Structure of the Earth

From Wikipedia, the free encyclopedia

 

Structure of the Earth

The internal structure of the Earth is layered in spherical shells: an outer silicate solid crust, a highly viscous asthenosphere and mantle, a liquid outer core that is much less viscous than the mantle, and a solid inner core. Scientific understanding of the internal structure of the 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 the Earth, measurements of the gravitational and magnetic fields of the Earth, and experiments with crystalline solids at pressures and temperatures characteristic of the Earth's deep interior.

Mass

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.

Structure

Earth's radial density distribution according to the preliminary reference earth model (PREM).

 

Earth's gravity according to the preliminary reference earth model (PREM). Comparison to approximations using constant and linear density for Earth's interior.

 

Mapping the interior of Earth with earthquake waves.

 

Schematic view of the interior of Earth. 1. continental crust – 2. oceanic crust – 3. upper mantle – 4. lower mantle – 5. outer core – 6. inner core – A: Mohorovičić discontinuity – B: Gutenberg Discontinuity – C: Lehmann–Bullen discontinuity.

 

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		Layer
Kilometres	Miles
0–60	0–37	Lithosphere (locally varies between 5 and 200 km)
0–35	0–22	… Crust (locally varies between 5 and 70 km)
35–60	22–37	… Uppermost part of mantle
35–2,890	22–1,790	Mantle
210-270	130-168	… Upper mesosphere (upper mantle)
660–2,890	410–1,790	… Lower mesosphere (lower mantle)
2,890–5,150	1,790–3,160	Outer core
5,150–6,360	3,160–3,954	Inner core

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.

Crust

The 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.

Mantle

World map showing the position of the Moho.

Earth's mantle extends to a depth of 2,890 km, making it the thickest layer of Earth. The mantle is divided into upper and lower mantle. The upper and lower mantle are separated by the transition zone. The lowest part of the mantle next to the core-mantle boundary is known as the D″ (pronounced dee-double-prime) layer. The pressure at the bottom of the mantle is ≈140 GPa (1.4 Matm). The mantle is composed of silicate rocks that are rich in iron and magnesium relative to the overlying crust. Although solid, the high temperatures within the mantle cause the silicate material to be sufficiently ductile that it can flow on very long timescales. Convection of the mantle is expressed at the surface through the motions of tectonic plates. As there is intense and increasing pressure as one travels deeper into the mantle, the lower part of the mantle flows less easily than does the upper mantle (chemical changes within the mantle may also be important). The viscosity of the mantle ranges between 10²¹ and 10²⁴ Pa·s, depending on depth.^[6] In comparison, the viscosity of water is approximately 10⁻³ Pa·s and that of pitch is 10⁷ Pa·s. The source of heat that drives plate tectonics is the primordial heat left over from the planet’s formation as well as the radioactive decay of uranium, thorium, and potassium in Earth’s crust and mantle.

Core

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 ${\displaystyle {\tfrac {9}{5}}}$ 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 critical of 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, 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 crus. 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.

Recent speculation suggests that the innermost part of the core is enriched in gold, platinum and other siderophile elements.

The matter that comprises Earth is connected in fundamental ways to matter of certain chondrite meteorites, and to matter of outer portion of the Sun. There is good reason to believe that Earth is, in the main, like a chondrite meteorite. 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, while totally ignoring another, albeit less abundant type, called enstatite chondrites. The principal difference between the two meteorite types is that enstatite chondrites formed under circumstances of 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 strength in Earth's outer core is estimated to be 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; however, more recent studies in 2011 found this hypothesis to be inconclusive. Options remain for the core which may be oscillatory in nature or a chaotic system. In August 2005 a team of geophysicists announced in the journal Science that, according to their estimates, Earth's inner core rotates approximately 0.3 to 0.5 degrees per year faster relative to the rotation of the surface.

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.

Mantle (geology)

From Wikipedia, the free encyclopedia

The mantle is a layer inside a terrestrial planet and some other rocky planetary bodies. For a mantle to form, the planetary body must be large enough to have undergone the process of planetary differentiation by density. The mantle is bounded on the bottom by the planetary core and on top by the crust. The terrestrial planets (Earth, Venus, Mars and Mercury), the Moon, two of Jupiter's moons (Io and Europa) and the asteroid Vesta each has a mantle made of silicate rock. Interpretation of spacecraft data suggests that at least two other moons of Jupiter (Ganymede and Callisto), as well as Titan and Triton, each have a mantle made of ice or other solid volatile substances.

Earth's mantle

The internal structure of Earth

The interior of Earth, similar to the other terrestrial planets, is divided into layers of different composition. The mantle is a layer between the crust and the outer core. Earth's mantle is a silicate rocky shell with an average thickness of 2,886 kilometres (1,793 mi). The mantle makes up about 84% of Earth's volume. It is predominantly solid but in geological time it behaves as a viscous fluid. The mantle encloses the hot core rich in iron and nickel, which makes up about 15% of Earth's volume. Past episodes of melting and volcanism at the shallower levels of the mantle have produced a thin crust of crystallized melt products near the surface. Information about the structure and composition of the mantle has been obtained from geophysical investigation and from direct geoscientific analyses of Earth mantle-derived xenoliths and mantle that has been exposed by mid-oceanic ridge spreading.

Two main zones are distinguished in the upper mantle: the inner asthenosphere composed of plastic flowing rock of varying thickness, on average about 200 km (120 mi) thick, and the lowermost part of the lithosphere composed of rigid rock about 50 to 120 km (31 to 75 mi) thick. A thin crust, the upper part of the lithosphere, surrounds the mantle and is about 5 to 75 km (3.1 to 46.6 mi) thick. Recent analysis of hydrous ringwoodite from the mantle suggests that there is between one and three times as much water in the transition zone between the lower and upper mantle than in all the world's oceans combined.

In some places under the ocean the mantle is actually exposed on the surface of Earth. There are also a few places on land where mantle rock has been pushed to the surface by tectonic activity, most notably the Tablelands region of Gros Morne National Park in the Canadian province of Newfoundland and Labrador and Zabargad Island (St. John's Island) in the Red Sea. (Also Macquarie Island, Saint Peter and Saint Paul Archipelago, Troodos Ophiolite, Lizard Complex, Semail Ophiolite, and other ophiolites)

Structure

The mantle is divided into sections which are based upon results from seismology. These layers (and their thicknesses/depths) are the following:

• the upper mantle (starting at the Moho, or base of the crust around 7 to 35 km (4.3 to 21.7 mi) downward to 410 km (250 mi)),
• the transition zone (410–660 km or 250–410 mi),
• the lower mantle (660–2,891 km or 410–1,796 mi),
• anomalous core–mantle boundary with a variable thickness (on average ~200 km (120 mi) thick).

The top of the mantle is defined by a sudden increase in seismic velocity, which was first noted by Andrija Mohorovičić in 1909; this boundary is now referred to as the Mohorovičić discontinuity or "Moho". The uppermost mantle plus overlying crust are relatively rigid and form the lithosphere, an irregular layer with a maximum thickness of perhaps 200 km (120 mi).

Below the lithosphere the upper mantle becomes notably more plastic. In some regions below the lithosphere, the seismic shear velocity is reduced; this so-called low-velocity zone (LVZ) extends down to a depth of several hundred km. Inge Lehmann discovered a seismic discontinuity at about 220 km (140 mi) depth; although this discontinuity has been found in other studies, it is not known whether the discontinuity is ubiquitous.

The transition zone is an area of great complexity; it physically separates the upper and lower mantle.

Very little is known about the lower mantle apart from that it appears to be relatively seismically homogeneous. The D" layer at the core–mantle boundary separates the mantle from the core.

In 2015, research using gravitational data from GRACE satellites and the long wavelength nonhydrostatic geoid indicated viscosity increases by a factor of ten to 150 about 1,000 kilometres (620 mi) below earth's surface; separate research also indicates sinking tectonic plates stall at this depth, leading Robert van der Hilst to speculate "In term's of structure and dynamics, 1,000 kilometers could be more important" (than the currently accepted 660 km depth upper—lower division). The lower mantle also contains some discontinuous zones, called "thermochemical piles" which have been interpreted as either thermally differentiated, upwellings bringing warmer material towards the surface, or as chemically differentiated material.

A principal source of the heat that drives plate tectonics is the radioactive decay of uranium, thorium, and potassium in Earth’s crust and mantle.

Characteristics

The mantle differs substantially from the crust in its mechanical properties as the direct consequence of the difference in composition (expressed as different mineralogy). The distinction between crust and mantle is based on chemistry, rock types, rheology and seismic characteristics. The crust is a solidification product of mantle derived melts, expressed as various degrees of partial melting products during geologic time. Partial melting of mantle material is believed to cause incompatible elements to separate from the mantle, with less dense material floating upward through pore spaces, cracks, or fissures, that would subsequently cool and solidify at the surface. Typical mantle rocks have a higher magnesium to iron ratio and a smaller proportion of silicon and aluminium than the crust. This behavior is also predicted by experiments that partly melt rocks thought to be representative of Earth's mantle.

Mapping the interior of the Earth with earthquake waves.

Mantle rocks shallower than about 410 km (250 mi) depth consist mostly of olivine, pyroxenes, spinel-structure minerals, and garnet; typical rock types are thought to be peridotite, dunite (olivine-rich peridotite), and eclogite. Between about 400 km (250 mi) and 650 km (400 mi) depth, olivine is not stable and is replaced by high pressure polymorphs with approximately the same composition: one polymorph is wadsleyite (also called beta-spinel type), and the other is ringwoodite (a mineral with the gamma-spinel structure). Below about 650 km (400 mi), all of the minerals of the upper mantle begin to become unstable. The most abundant minerals present, the silicate perovskites, have structures (but not compositions) like that of the mineral perovskite followed by the magnesium/iron oxide ferropericlase. The changes in mineralogy at about 400 and 650 km (250 and 400 mi) yield distinctive signatures in seismic records of the Earth's interior, and like the moho, are readily detected using seismic waves. These changes in mineralogy may influence mantle convection, as they result in density changes and they may absorb or release latent heat as well as depress or elevate the depth of the polymorphic phase transitions for regions of different temperatures. The changes in mineralogy with depth have been investigated by laboratory experiments that duplicate high mantle pressures, such as those using the diamond anvil.

Composition of Earth's mantle in weight percent

Element	Amount		Compound	Amount
O	44.8
Mg	22.8		SiO2	46
Si	21.5		MgO	37.8
Fe	5.8		FeO	7.5
Ca	2.3		Al2O3	4.2
Al	2.2		CaO	3.2
Na	0.3		Na2O	0.4
K	0.03		K2O	0.04
Sum	99.7		Sum	99.1

The inner core is solid, the outer core is liquid, and the mantle solid/plastic. This is because of the relative melting points of the different layers (nickel–iron core, silicate crust and mantle) and the increase in temperature and pressure as depth increases. At the surface both nickel–iron alloys and silicates are sufficiently cool to be solid. In the upper mantle, the silicates are generally solid (localised regions with small amounts of melt exist); however, as the upper mantle is both hot and under relatively little pressure, the rock in the upper mantle has a relatively low viscosity. In contrast, the lower mantle is under tremendous pressure and therefore has a higher viscosity than the upper mantle. The metallic nickel–iron outer core is liquid because of the high temperature, despite the high pressure. As the pressure increases, the nickel–iron inner core becomes solid because the melting point of iron increases dramatically at these high pressures.

Temperature

In the mantle, temperatures range between 500 to 900 °C (932 to 1,652 °F) at the upper boundary with the crust to over 4,000 °C (7,230 °F) at the boundary with the core. The geothermal gradient of the mantle increases rapidly in the thermal boundary layers at the top and bottom of the mantle, and increases gradually through the interior of the mantle. Although the higher temperatures far exceed the melting points of the mantle rocks at the surface (about 1200 °C for representative peridotite), the mantle is almost exclusively solid. The enormous lithostatic pressure exerted on the mantle prevents melting, because the temperature at which melting begins (the solidus) increases with pressure.

Movement

This figure is a snapshot of one time-step in a model of mantle convection. Colors closer to red are hot areas and colors closer to blue are cold areas. In this figure, heat received at the core–mantle boundary results in thermal expansion of the material at the bottom of the model, reducing its density and causing it to send plumes of hot material upwards. Likewise, cooling of material at the surface results in its sinking.

Because of the temperature difference between the Earth's surface and outer core and the ability of the crystalline rocks at high pressure and temperature to undergo slow, creeping, viscous-like deformation over millions of years, there is a convective material circulation in the mantle. Hot material upwells, while cooler (and heavier) material sinks downward. Downward motion of material occurs at convergent plate boundaries called subduction zones. Locations on the surface that lie over plumes are predicted to have high elevation (because of the buoyancy of the hotter, less-dense plume beneath) and to exhibit hot spot volcanism. The volcanism often attributed to deep mantle plumes is alternatively explained by passive extension of the crust, permitting magma to leak to the surface (the "Plate" hypothesis).

The convection of the Earth's mantle is a chaotic process (in the sense of fluid dynamics), which is thought to be an integral part of the motion of plates. Plate motion should not be confused with continental drift which applies purely to the movement of the crustal components of the continents. The movements of the lithosphere and the underlying mantle are coupled since descending lithosphere is an essential component of convection in the mantle. The observed continental drift is a complicated relationship between the forces causing oceanic lithosphere to sink and the movements within Earth's mantle.

Although there is a tendency to larger viscosity at greater depth, this relation is far from linear and shows layers with dramatically decreased viscosity, in particular in the upper mantle and at the boundary with the core. The mantle within about 200 km (120 mi) above the core–mantle boundary appears to have distinctly different seismic properties than the mantle at slightly shallower depths; this unusual mantle region just above the core is called D″ ("D double-prime"), a nomenclature introduced over 50 years ago by the geophysicist Keith Bullen. D″ may consist of material from subducted slabs that descended and came to rest at the core–mantle boundary and/or from a new mineral polymorph discovered in perovskite called post-perovskite.

Earthquakes at shallow depths are a result of stick-slip faulting; however, below about 50 km (31 mi) the hot, high pressure conditions ought to inhibit further seismicity. The mantle is considered to be viscous and incapable of brittle faulting. However, in subduction zones, earthquakes are observed down to 670 km (420 mi). A number of mechanisms have been proposed to explain this phenomenon, including dehydration, thermal runaway, and phase change. The geothermal gradient can be lowered where cool material from the surface sinks downward, increasing the strength of the surrounding mantle, and allowing earthquakes to occur down to a depth of 400 km (250 mi) and 670 km (420 mi).

The pressure at the bottom of the mantle is ~136 GPa (1.4 million atm). Pressure increases as depth increases, since the material beneath has to support the weight of all the material above it. The entire mantle, however, is thought to deform like a fluid on long timescales, with permanent plastic deformation accommodated by the movement of point, line, and/or planar defects through the solid crystals comprising the mantle. Estimates for the viscosity of the upper mantle range between 10¹⁹ and 10²⁴ Pa·s, depending on depth, temperature, composition, state of stress, and numerous other factors. Thus, the upper mantle can only flow very slowly. However, when large forces are applied to the uppermost mantle it can become weaker, and this effect is thought to be important in allowing the formation of tectonic plate boundaries.

Exploration

Exploration of the mantle is generally conducted at the seabed rather than on land because of the relative thinness of the oceanic crust as compared to the significantly thicker continental crust.
The first attempt at mantle exploration, known as Project Mohole, was abandoned in 1966 after repeated failures and cost over-runs. The deepest penetration was approximately 180 m (590 ft). In 2005 an oceanic borehole reached 1,416 metres (4,646 ft) below the sea floor from the ocean drilling vessel JOIDES Resolution.

On 5 March 2007, a team of scientists on board the RRS James Cook embarked on a voyage to an area of the Atlantic seafloor where the mantle lies exposed without any crust covering, midway between the Cape Verde Islands and the Caribbean Sea. The exposed site lies approximately three kilometres beneath the ocean surface and covers thousands of square kilometres. A relatively difficult attempt to retrieve samples from the Earth's mantle was scheduled for later in 2007. The Chikyu Hakken mission attempted to use the Japanese vessel Chikyū to drill up to 7,000 m (23,000 ft) below the seabed. This is nearly three times as deep as preceding oceanic drillings.

A novel method of exploring the uppermost few hundred kilometres of the Earth was proposed in 2005, consisting of a small, dense, heat-generating probe which melts its way down through the crust and mantle while its position and progress are tracked by acoustic signals generated in the rocks. The probe consists of an outer sphere of tungsten about one metre in diameter with a cobalt-60 interior acting as a radioactive heat source. It was calculated that such a probe will reach the oceanic Moho in less than 6 months and attain minimum depths of well over 100 km (62 mi) in a few decades beneath both oceanic and continental lithosphere.

Exploration can also be aided through computer simulations of the evolution of the mantle. In 2009, a supercomputer application provided new insight into the distribution of mineral deposits, especially isotopes of iron, from when the mantle developed 4.5 billion years ago.