Basics -- Earth'sInterior
Introduction
Tour through the layers of earth
Surface to core, chemically
Crust
Mantle
Core
Surface to core, physically
Lithosphere
Asthenosphere
Upper mesosphere
Lower mesosphere
Outer core
Inner core
Earth's magnetic field originates in the core
Beyond simple layers
How do we know?
Igneous rocks and fault blocks
Seismic waves
Gravity
Moment of inertia
Meteorites
Experiments
Introduction
If we could journey to the center of the earth we would have to travel about 6,400 km (4,000 miles). Along the way to earth's core we would pass layers of rock that can be classified in two different ways, either by their chemistry or their physical behavior.
According to the chemical composition of the rocks, earth's interior can be differentiated into three layers - crust, mantle, and core.
When considering the rocks of earth's interior in terms of their physical behavior, six layers can be differentiated from the surface to the core. The characteristics that distinguish these six different layers are based on the relative strength of a given layer in response to stress and whether it is solid or liquid.
Tour through the layers of earth
The chemical compositions and physical behaviors of the rocks of earth's interior converge in many ways but it is worthwhile to become familiar with the each of the different set of layers individually before exploring how they overlap.
The chemical composition and physical behavior of rock inside the earth relate to each other because the chemical composition of a rock is one of the factors that determines its physical behavior. However, the physical behavior of rock also depends on the pressure and temperature it is subjected to at its depth within the earth. As depth inside the earth increases, the pressure and temperature increase. Some layers in the earth are harder or softer than adjacent layers, even though they have the same composition, because they are at different pressures and temperatures.
Surface to core, chemically
Crust
The tour starts at the surface with earth's crust. Generally speaking, the crust is predominately silicon oxide and aluminum oxide. Continental crust is thicker and less dense than oceanic crust. Earth's crust varies in thickness from less than 5 km (under mid-ocean spreading ridges) to more than 70 km (beneath the highest mountain range).
Mantle
The next layer down chemically is the mantle. The mantle has an ultramafic composition - it contains more iron, magnesium, less aluminum and somewhat less silicon than the crust. The mantle is roughly 2,900 km thick. In terms of volume, the mantle is the largest of earth's three chemical layers.
Core
The final stop on the chemical tour is the core, which is mostly iron and nickel. The core is about 3,500 km thick.
The following table summarizes the chemical layers of the earth.
Chemical Layers of Earth
CrustMantleCore
composition: high Si, Al, & Ocomposition: moderate Si, high Mg & Fecomposition: Fe & Ni
thickness: 5 to 70 kmthickness: 2,900 kmthickness: 3,500 km
Surface to core, physically
Lithosphere
Starting at the surface, the first layer is the lithosphere. We humans, and the other creatures that live on earth, occupy the surface of the lithosphere. The lithosphere is entirely solid except where there are zones of magma beneath volcanoes or in places undergoing magma intrusion. The volume of molten rock in the lithosphere is a tiny fraction, less than 0.1%, of the volume of the entire lithosphere.
The lithosphere itself has two parts. The top part is the crust. The bottom part is the lithospheric mantle. The two components of the lithosphere, in combination, form a relatively strong, rigid layer of rock that covers the earth. Earth's tectonic plates, all of which are in motion relative to each other, make up the lithosphere.
Asthenosphere
Beneath the lithosphere is a relatively weak and ductile layer of the mantle called the asthenosphere. Although the asthenosphere is solid, not liquid, it flows at geological rates, up to several cm (several inches) per year. In other words, the asthenosphere behaves much more plastically than the rigid lithosphere above it does.
The chemical composition of the asthenosphere is about the same as the chemical composition of the overlying lithospheric mantle. Why, then, is the asthenosphere soft and the lithosphere rigid? It is because at the depth of the asthenosphere, temperatures are very close to the melting point of the rock, weakening the rock. In fact, it is thought likely, from indirect evidence, that there is a small percentage of molten rock in the tiny spaces between the minerals of the asthenosphere, which contributes to the soft nature of the rock. However, the solid minerals of the asthenosphere are extensively in contact with each other, forming a material that is solid overall despite the possible presence of a small amount of partial melt.
The asthenosphere is the primary source of most magma. Because the asthenosphere is close to its melting point and may contain everywhere a small proportion of partly molten rock, it does not take much to cause magma to form and separate from the asthenosphere. As explained in the igneous rocks Basics page, melting of the asthenosphere can be caused by addition of fluid, particularly water, or by a decrease in pressure. In subduction zones, slabs of sinking lithosphere release water into the asthenosphere, causing the asthenosphere to melt and produce magma which rises into and through the crust, producing the volcanic arc that is found at every subduction zone. At divergent plate boundaries, the asthenosphere flows upward, or upwells, which reduces the lithostatic pressure enough to cause the rock of the asthenosphere to melt. That is why divergent plate boundaries are volcanic zones. The solidified lavas and intrusions at divergent plate boundaries produce new lithosphere, which fills in and replaces the plate material that spreads away on either side.
Upper mesosphere
Beneath the asthenosphere is the rest of the mantle, the mesosphere. The mesosphere makes up most of the volume of the mantle. The mesosphere is entirely solid. The temperature and pressure of the rock in the mesosphere keep it from breaking; therefore, no earthquakes originate from the mesosphere.
The upper mesosphere is a transition zone in which the rock rapidly becomes denser with depth in response to the increasing lithostatic pressure.
Lower mesosphere
The lower mesosphere starts at a depth of 660 km from earth's surface. At that depth there is an abrupt increase in density. This increase is caused by changes in the crystal structures of the most abundant minerals in the rock. These minerals change from less dense crystal structures above the boundary to more dense crystal structures below the boundary. The lower mesosphere undergoes little density change from its top boundary at 660 km to its base at 2900 km where it meets the outer core.
Outer core
The bottom of the mesosphere is the boundary with the earth's core. The core is about twice as dense as the crust, and about 1.5 times as dense as the mantle. The outer core is liquid, as was discovered when it was first observed that S-waves will not pass through it.
Inner core
The inner core is solid. The inner and outer cores are made of the same iron-rich, metallic composition. The temperature of the inner core is not very much greater than the temperature of the outer core. However, lithostatic pressure keeps increasing with depth and the inner core has the great weight of the rest of the earth pressing in on it. The pressure on the inner core is high enough to keep it in the solid state.
The following table summarizes the physical layers of the earth.
Physical Layers of Earth
LithosphereAsthenosphereUpper MesosphereLower MesosphereOuter CoreInner Core
physical behavior:
rigid, brittle at shallow depthsphysical behavior:
ductilephysical behavior:
rigid, not brittle, rapid increase in density with depthphysical behavior:
denser and more rigid than upper mesospherephysical behavior:
liquidphysical behavior:
rigid, not brittle
thickness:
5 to 200 kmthickness:
100 to 300 kmthickness:
300 to 400 kmthickness:
2,300 kmthickness:
2,300 kmthickness:
1,200 km
Earth's magnetic field originates in the core
The liquid outer core is the source of the earth's magnetic field, as a result of its metallic nature, which means it contains electrons not attached to particular nuclei. Heat is transferred upward to the mantle from the inner core via convective cells, in which the liquid in the outer core flows in looping patterns. The combination of the loose electrons and looping convective flow with the rotation of the earth results in a geodynamo that produces a magnetic field. Because the magnetic field is generated by a dynamically convecting and rotating sphere of liquid, it is unstable. Every now and then, after several hundred thousand to several million years, the earth's magnetic field becomes unstable to the point that it temporarily shuts down. When it restarts, its north and south magnetic poles must inevitably be reversed, according to the physics of magnetic fields produced spontaneously from geodyamos. (For comparison, the magnetic field of the Sun, which is also produces by convecting electrical charges in a rotating sphere, becomes magnetically unstable and reverses its magnetic field on a more regular basis, every 11 years.)
Given that the inner core is a solid metallic sphere, made mostly of iron and nickel, surrounded entirely by liquid, it can be pictured as a giant ball bearing spinning in a pressurized fluid. Detailed studies of earthquake waves passing though the inner core have found evidence that it is spinning - rotating - just slightly faster than the rest of the earth.
Beyond simple layers
The interior of the earth is not simply layered. Some of the layers, particularly the crust and lithosphere, are highly variable in thickness. The boundaries between layers are rough and irregular. Some layers penetrate other layers at certain places. Variations in the thickness of the earth's layers, irregularities in layer boundaries, and interpenetrations of layers, reflect the dynamic nature of the earth.
For example, the lithosphere penetrates deep into the mesosphere at subduction zones. Although it is still a matter of research and debate, there is some evidence that subducted plates may penetrate all the way into the lower mesosphere. If so, plate tectonics is causing extensive mixing and exchange of matter in the earth, from the bottom of the mantle to the top of the crust.
As another example, hot spots may be places where gases and fluids rise from the core-mantle boundary, along with heat. Studies of helium isotopes in hot spot volcanic rocks find evidence that much of the helium comes from deep in the earth, probably from the lower mesosphere.
How do we know?
We humans have no hands-on access to samples of the earth's interior from deeper than the upper mantle. The earth's core is so dense and so deep, it is completely inaccessible. Contrary to a popular misconception, lava does not come from the earth's core. Magma and lava come from only the lithosphere and asthenosphere, the upper 200 km of earth's 6,400 km thickness. Attempts have been made to drill through the crust to reach the mantle, without success. Given the lack of actual pieces of the earth from deeper than the asthenosphere, how do we know about the internal layers of the earth, what they are made of, and what their properties and processes are?
Igneous rocks and fault blocks
There are two sources of rock samples from the lower lithosphere and asthenosphere, igneous rocks and fault blocks. Some igneous rocks contain xenoliths, pieces of solid rock that were adjacent to the body of magma, became incorporated into the magma, and were carried upward in the magma. From xenoliths in plutonic and volcanic igneous rocks, many samples of the lower crust and upper mantle have been identified and studied.
Another source of pieces of the lower crust and upper mantle is fault zones and exposed orogenic zones (root zones of mountains that have been exposed after much uplift and erosion). Some slabs of thrust-faulted rock contain lithospheric mantle rock. In ophiolites (see the ophiolite discussion on the exotic terranes Basics page), ultramafic rock from the mantle part of the lithosphere is a defining attribute. Most ophiolites and thrust-faulted slices of rock that contain pieces of the upper mantle are related to either subduction zones or transform plate boundaries.
Seismic waves
As discussed on the earthquakes Basics page, P-waves and S-waves move through different parts of the earth's interior in different ways. Analogous to how you can see what is in the room around you by interpreting the light that your eyes receive, light that has interacted with the things in the room around you to give it its characteristics, seismologists can interpret recordings of seismic waves to "see" inside the earth. Such imaging of the earth's interior is based on how the different layers of the earth have affected the seismic waves in different ways.
Where seismic waves speed up or slow down, they refract, changing the direction in which they are traveling. Where seismic waves encounter an abrupt boundary between two very different layers, some of the seismic wave energy is reflected, bouncing back at the same angle it struck. The reflections and refractions of seismic waves allow the layers and boundaries within the earth to be located and studied.
Here are some examples of what we have been able to distinguish in the earth's interior from the study of seismic waves and how they travel through the layers of the earth:
The thickness of the crust. This is a measure of the thickness of the crust based on the abrupt increase in speed of seismic waves that occurs when they enter the mantle. The boundary between the crust and mantle, as inferred from the change in the speed of P- and S-waves, is called the Mohorovicic discontinuity, named after the Croatian seismologist who first discerned it; usually it is referred to simply as the Moho. It is mainly from seismic waves that we know how thin oceanic crust is and how thick continental crust is.
The thickness of the lithosphere. Where seismic waves pass down from the lithosphere into the asthenosphere, they slow down. This is because of the lower rigidity and compressibility of the rocks in the layer below the lithosphere. The zone below the lithosphere where seismic waves travel more slowly is called the low velocity zone. The low velocity zone is probably coincident with the asthenosphere.
The boundary between the upper and lower mesosphere (upper and lower mantle). This shows up as an increase in seismic wave speed at a depth of 660 km.
The boundary between the mantle and the core. This is marked by S-waves coming to an abrupt stop, presumably because the outer core is liquid, and a sudden large reduction in the speed of P-waves, as they enter the liquid core where there is no rigidity to contribute to P-wave speed.
The inner core. This was first recognized by refraction of P-waves passing through this part of the core, due to an abrupt increase in their speed, which was not shown by P-waves traveling through only the outer part of the core.
Seismic tomography: imaging slabs and masses at various orientations in the earth, not just in layers. By combining data from many seismometers, three-dimensional images of zones in the earth that have higher or lower seismic wave speeds can be constructed. Seismic tomography shows that in some places there are masses of what may be subducted plates that have penetrated below the asthenosphere into the mesosphere and, in some cases, penetrated into the lower mesosphere, the deepest part of the mantle. In other places, subducted plates appear to have piled up at the base of the upper mesosphere without penetrating into the lower mesosphere.
Gravity
Isaac Newton was the first to calculate the total mass of the earth. This gives us an important constraint on what the earth is made of, because, by dividing the mass of the earth by the volume of the earth, we know the average density of the earth. Whatever the earth is made of, it must add up to the correct amount of mass. Gravity measurements, and the earth's mass, tell us that the interior of the earth must be denser than the crust, because the average density of earth is much higher than the density of the crust.
Because different parts of the crust, mantle, and core have different thicknesses and densities, the strength of gravity over particular points on earth varies slightly. These variations from the average strength of earth's gravity are called gravity anomalies. Mapping and analyzing gravity anomalies, in some cases by using satellites, and also be measuring the effect of gravity anomalies on the surface shape of the ocean, has given us much insight into subduction zones, mid-ocean spreading ridges, and mountain ranges, including constraints on the depths of their roots.
Moment of inertia
The earth's gravity tells us how much total mass the earth has, but does not tell us how the mass is distributed within the earth. A property known as moment of inertia, which is the resistance (inertia) of an object to changes in its spin (rotation), is determined by exactly how matter is distributed in a spinning object, from its core to its surface. The earth's moment of inertia is measured by its effect on other objects with which it interacts gravitationally, including the Moon, and satellites. Knowing the earth's moment of inertia provides a way of checking and refining our understanding of the mass and density of each of the earth's internal layers.
Meteorites
Studies of meteorites, which are pieces of asteroids that have landed on earth, along with astronomical studies of what the Sun, the other planets, and orbiting asteroids are made of, give us a model for the general chemical composition of objects in the inner solar system, which are made mainly of elements that form rocks and metals, as opposed to the outer planets such as Jupiter, which are made mostly of light, gas-forming elements. The general compositional model of the rocky and metallic part of the solar system has much higher percentages of iron, nickel, and magnesium than is found in the earth's crust.
If the earth's mantle is made of ultramafic rock, as is found in actual samples of the upper mantle in xenoliths and ophiolites, that would account for part of the missing iron, nickel, and magnesium. But much more iron and nickel would still be missing. If the core is made mostly of iron, and abundant nickel as well, it would give the earth an overall composition similar to the composition of other objects in the inner solar system, and similar to the proportions of rock and metal-forming elements measured in the Sun.
A mantle with an ultramafic composition, and a core made mostly of iron plus nickel, would make earth's composition match the composition of the rest of the solar system, and give those layers the right densities to account for the earth's moment of inertia and total mass.
Experiments
Geology, like other sciences, is based on experiment along with observation and theory. earth scientists and physicists have developed experimental methods to study how materials behave at the pressures and temperatures of the earth's interior, including core temperatures and pressures. They can measure such properties as the density, the state of matter (liquid or solid), the rigidity, the compressibility, and the speed at which seismic waves pass through these materials at high pressures and temperatures. These studies allow further refinement of our knowledge of what the interior of the earth is made of and how it behaves. These experiments support the theory that the mantle is ultramafic and the core is mostly iron and nickel, because they show that materials with those compositions have the same density and seismic wave speeds as have been observed in the earth.
The Earth's mantle is divided into two major rheological layers: the rigid lithosphere comprising the uppermost mantle, and the more viscous asthenosphere, separated by the lithosphere-asthenosphere boundary. Lithosphere underlying ocean crust has a thickness of around 100 km, whereas lithosphere underlying continental crust generally has a thickness of 150-200 km.[3] The lithosphere and overlying crust make up tectonic plates, which move over the asthenosphere.
the transition zone (approximately 410–660 km or 250–410 mi), in which wadsleyite (≈ 410–520 km or 250–320 mi) and ringwoodite (≈ 525–660 km or 326–410 mi) are stable
the lower mantle (approximately 660–2,891 km or 410–1,796 mi), in which bridgmanite (≈ 660–2,685 km or 410–1,668 mi) and post-perovskite (≈ 2,685–2,891 km or 1,668–1,796 mi) are stable
The lower ~200 km of the lower mantle constitutes the D" (D-double-prime) layer, a region with anomalous seismic properties. This region also contains LLSVPs and ULVZs.
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".[5][6]
The upper mantle is dominantly peridotite, composed primarily of variable proportions of the minerals olivine, clinopyroxene, orthopyroxene, and an aluminous phase. The aluminous phase is plagioclase in the uppermost mantle, then spinel, and then garnet below ~100 km. Gradually through the upper mantle, pyroxenes become less stable and transform into majoritic garnet.
At the top of the transition zone, olivine undergoes isochemical phase transitions to wadsleyite and ringwoodite. Unlike nominally anhydrous olivine, these high-pressure olivine polymorphs have a large capacity to store water in their crystal structure. This has led to the hypothesis that the transition zone may host a large quantity of water.[7] At the base of the transition zone, ringwoodite decomposes into bridgmanite (formerly called magnesium silicate perovskite), and ferropericlase. Garnet also becomes unstable at or slightly below the base of the transition zone.
The lower mantle is composed primarily of bridgmanite and ferropericlase, with minor amounts of calcium perovskite, calcium-ferrite structured oxide, and stishovite. In the lowermost ~200 km of the mantle, bridgmanite isochemically transforms into post-perovskite.
The chemical composition of the mantle is difficult to determine with a high degree of certainty because it is largely inaccessible. Rare exposures of mantle rocks occur in ophiolites, where sections of oceanic lithosphere have been obducted onto a continent. Mantle rocks are also sampled as xenoliths within basalts or kimberlites.
Composition of the Earth's upper mantle (depleted MORB)[8][9]
Most estimates of the mantle composition are based on rocks that sample only the uppermost mantle. There is debate as to whether the rest of the mantle, especially the lower mantle, has the same bulk composition.[10] The mantle's composition has changed through the Earth's history due to the extraction of magma that solidified to form oceanic crust and continental crust.
In the mantle, temperatures range from approximately 200 °C (392 °F) at the upper boundary with the crust to approximately 4,000 °C (7,230 °F) at the core-mantle boundary.[11] 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.[12]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.[13] The enormous lithostatic pressure exerted on the mantle prevents melting, because the temperature at which melting begins (the solidus) increases with pressure.
The pressure in the mantle increases from a few kbar at the Moho to 1390 kbar (139 GPa) at the core-mantle boundary[11].
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.[14] 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 spotvolcanism. 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).[15]
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.[16] 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.[17]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 strike-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).[18] 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 composing the mantle. Estimates for the viscosity of the upper mantle range between 1019 and 1024Pa·s, depending on depth,[16]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 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 Atlanticseafloor 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.[20][21] A relatively difficult attempt to retrieve samples from the Earth's mantle was scheduled for later in 2007.[22] 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.[23] 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 continentallithosphere.[24]
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.[
What is the Earth’s Mantle Made Of?
Like all the other terrestrial planets, (Mercury, Venus, and Mars) the Earth is made up of many layers. This is the result of it undergoing planetary differentiation, where denser materials sink to the center to form the core while lighter materials form around the outside. Whereas the core is composed primarily of iron and nickel, Earth’s upper layer are composed of silicate rock and minerals.
This region is known as the mantle, and accounts for the vast majority of the Earth’s volume. Movement, or convection, in this layer is also responsible for all of Earth’s volcanic and seismic activity. Information about structure and composition of the mantle is either the result of geophysical investigation or from direct analysis of rocks derived from the mantle, or exposed mantle on the ocean floor.
Definition:
Composed of silicate rocky material with an average thickness of 2,886 kilometres (1,793 mi), the mantle sits between the Earth’s crust and its upper core. The mantle makes up 84% of the Earth by volume, compared to 15% in the core and the remainder being taken up by the crust. While it is predominantly solid, it behaves like a viscous fluid due to the fact that temperatures are close to the melting point in this layer.
The layers of the Earth, a differentiated planetary body. Credit: Wikipedia Commons/Surachit
Our knowledge of the upper mantle, including the tectonic plates, is derived from analyses of earthquake waves; heat flow, magnetic, and gravity studies; and laboratory experiments on rocks and minerals. Between 100 and 200 kilometers below the Earth’s surface, the temperature of the rock is near the melting point; molten rock erupted by some volcanoes originates in this region of the mantle.
Structure and Composition:
The mantle is divided into sections which are based upon results from seismology. These are the upper mantle, which extends from about 7 to 35 km (4.3 to 21.7 mi) from the surface down to a depth of 410 km (250 mi); the transition zone, which extends from 410 t0 660 km (250 – 410 mi); the lower mantle, which reaches from 660 km to a depth of 2,891 km (410 – 1,796 mi); and the the core-mantle boundary, which has a variable thickness (~200 km or 120 mi on average).
In the upper mantle two main zones are distinguished. The innermost of these is the inner asthenosphere, which is composed of plastic flowing rock of that averages about 200 km (120 mi) in thickness. The outer zone is the lowermost part of the lithosphere, which is composed of rigid rock and is about 50 to 120 km (31 to 75 mi) thick.
The upper part of the lithosphere is the Earth’s crust, a thin layer that is about 5 to 75 km (3.1 to 46.6 mi) thick, which is separated from the mantle by the Mohorovicic discontinuity (or “Moho”, which is defined by a sharp increase downward in the speed of earthquake waves).
The internal structure of Earth. Credit: Wikipedia Commons/Kelvinsong
In some places under the ocean, the mantle is actually exposed. 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 Newfoundland and Labrador, Canada, St. John’s Island, Egypt, or the island of Zabargad in the Red Sea.
In terms of its constituent elements, the mantle is made up of 44.8% oxygen, 21.5% silicon, and 22.8% magnesium. There’s also iron, aluminum, calcium, sodium, and potassium. These elements are all bound together in the form of silicate rocks, all of which take the form of oxides. The most common is Silicon dioxide (SiO2) at 48%, followed by Magnesium Oxide (MgO) at 37.8%. Examples of rocks that you might find inside the mantle include: olivine, pyroxenes, spinel, and garnet.
Convection:
Because of the temperature difference between the Earth’s surface and outer core, there is a convective material circulation in the mantle. This consists of the slow, creeping motion of the Earth’s silicate mantle across the surface, carrying heat from the interior of the Earth to the surface. Whereas hot material rises to the surface, cooler, heavier material sinks beneath.
The Pacific Ring of Fire, a string of volcanic regions extending from the South Pacific to South America. Credit: Public Domain
The lithosphere is divided into a number of plates that are continuously being created and consumed at their opposite plate boundaries. Downward motion of material occurs in subduction zones, locations at convergent plate boundaries where one mantle layer moves under another. Accretion occurs as material is added to the growing edges of a plate, associated with seafloor spreading.
This chaotic process is believed to be an integral part of the motion of plates, which in turn gives rise to continental drift. Subducted oceanic crust is also what gives rise to volcanism, as demonstrated by the Pacific Ring of Fire.
Exploration:
Scientific investigations and exploration of the mantle is generally conducted on the seabed due to the relative thickness of the oceanic crust compared to the continental crust. The first attempt at mantle exploration (known as Project Mohole) achieved a deepest penetration of approximately 180 meters (590 feet). It was abandoned in 1966 after repeated failures and cost over-runs.
In 2005, the ocean drilling vessel JOIDES Resolution achieved a borehole that was 1,416 meters (4,646 ft) in depth below the sea floor. In 2007, a team of scientists aboard the UK research ship RRS James Cook conducted a study on an exposed section of mantle located between the Cape Verdr Islands and the Caribbean Sea.
The scientific drilling ship JOIDES Resolution, pictured at sea in 2009. Credit: William Crawford/IODP/TAMU
In recent years, a method of exploring the Earth’s layers was proposed using a small, dense, heat-generating probe. This would melt its way through the crust and mantle and communicate via acoustic signals generated by its penetration of the rocks. The probe would consist of an outer shell of tungsten with a core of cobalt-60, which acts 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. In 2009, a supercomputer application created a simulation that provided new insight into the distribution of mineral deposits from when the mantle developed 4.5 billion years ago.
While the Earth’s mantle has yet to be explored at any significant depth, much has been learned from indirect studies over the past few centuries. As human exploration of the Solar System continues, we are sure to learn more about terrestrial planets, their geological behavior, and their formation.
Is the Earth’s mantle made of liquid magma?
The Earth’s mantle, on which the crust is lying on, is not made of liquid magma. It is not even made of magma. The Earth’s mantle is mostly made of solid rock. The misconception of a liquid mantle arises from expressions like “a subducted tectonic plate sinks into the mantle” or “continental drift”, expressions that implicitly refer to the liquid element.
The Earth’s mantle is mostly solid from the liquid outer core to the crust, but it can creep on the long-term, which surely strengthens the misconception of a liquid mantle.
What is Earth made of?
The Earth is made out of many things. Deep inside Earth, near its center, lies Earth's core which is mostly made up of nickel and iron. Above the core is Earth's mantle, which is made up of rock containing silicon, iron, magnesium, aluminum, oxygen and other minerals. The rocky surface layer of Earth, called the crust, is made up of mostly oxygen, silicon, aluminum, iron, calcium, sodium, potassium and magnesium. Earth's surface is mainly covered with liquid water and its atmosphere is is mainly nitrogen and oxygen, with smaller amounts of carbon dioxide, water vapor and other gases.
The Earth's surface lithosphere rides atop the asthenosphere and the two form the components of the upper mantle. The lithosphere is divided into a number of plates that are continuously being created and consumed at their opposite plate boundaries. Accretion occurs as mantle is added to the growing edges of a plate, associated with seafloor spreading. This hot added material cools down by conduction and convection of heat. At the consumption edges of the plate, the material has thermally contracted to become dense, and it sinks under its own weight in the process of subduction usually at an ocean trench.[3]
The subducted oceanic crust triggers volcanism, although the basic mechanisms are varied. Volcanism may occur due to processes that add buoyancy to partially melted mantle, which would cause upward flow of the partial melt due to decrease in its density. Secondary convection may cause surface volcanism as a consequence of intraplate extension[6] and mantle plumes.[7]
Mantle convection causes tectonic plates to move around the Earth's surface.[8] It seems to have been much more active during the Hadean period, resulting in gravitational sorting of heavier molten iron, nickel, and sulphides to the core and lighter silicate minerals to the mantle.
Earth cross-section showing location of upper (3) and lower (5) mantle
Earth's temperature vs depth. Dashed curve: layered mantle convection. Solid curve: whole-mantle convection.[7]
A superplume generated by cooling processes in the mantle.[9]
During the late 20th century, there was significant debate within the geophysics community as to whether convection is likely to be "layered" or "whole".[10][11] Although elements of this debate still continue, results from seismic tomography, numerical simulations of mantle convection and examination of Earth's gravitational field are all beginning to suggest the existence of 'whole' mantle convection, at least at the present time. In this model, cold, subducting oceanic lithosphere descends all the way from the surface to the core–mantle boundary (CMB) and hot plumes rise from the CMB all the way to the surface.[12] This picture is strongly based on the results of global seismic tomography models, which typically show slab and plume-like anomalies crossing the mantle transition zone.
Although it is now well accepted that subducting slabs cross the mantle transition zone and descend into the lower mantle, debate about the existence and continuity of plumes persists, with important implications for the style of mantle convection. This debate is linked to the controversy regarding whether intraplate volcanism is caused by shallow, upper-mantle processes or by plumes from the lower mantle.[6]Many geochemistry studies have argued that the lavas erupted in intraplate areas are different in composition from shallow-derived mid-ocean ridge basalts (MORB). Specifically, they typically have elevated Helium-3 – Helium-4 ratios. Being a primordial nuclide, Helium-3 is not naturally produced on earth. It also quickly escapes from earth's atmosphere when erupted. The elevated He-3/He-4 ratio of Ocean Island Basalts (OIBs) suggest that they must be sources from a part of the earth that has not previously been melted and reprocessed in the same way as MORB source has been. This has been interpreted as their originating from a different, less well-mixed, region, suggested to be the lower mantle. Others, however, have pointed out that geochemical differences could indicate the inclusion of a small component of near-surface material from the lithosphere.
On Earth, the Rayleigh number for convection within Earth's mantle is estimated to be of order 107, which indicates vigorous convection. This value corresponds to whole mantle convection (i.e. convection extending from the Earth's surface to the border with the core). On a global scale, surface expression of this convection is the tectonic plate motions, and therefore has speeds of a few cm/a.[13][14][15] Speeds can be faster for small-scale convection occurring in low-viscosity regions beneath the lithosphere, and slower in the lowermost mantle where viscosities are larger. A single shallow convection cycle takes on the order of 50 million years, though deeper convection can be closer to 200 million years.[16]
Currently, whole mantle convection is thought to include broad-scale downwelling beneath the Americas and the Western Pacific, both regions with a long history of subduction, and upwelling flow beneath the central Pacific and Africa, both of which exhibit dynamic topographyconsistent with upwelling.[17] This broad-scale pattern of flow is also consistent with the tectonic plate motions, which are the surface expression of convection in the Earth's mantle and currently indicate degree-2 convergence toward the western Pacific and the Americas, and divergence away from the central Pacific and Africa.[18] The persistence of net tectonic divergence away from Africa and the Pacific for the past 250 Myr indicates the long-term stability of this general mantle flow pattern,[18] and is consistent with other studies [19][20][21] that suggest long-term stability of the LLSVP regions of the lowermost mantle that form the base of these upwellings.
Since the mantle is primarily composed of olivine ((Mg,Fe)2SiO4), the rheological characteristics of the mantle are largely those of olivine. Additionally, due to the varying temperatures and pressures between the lower and upper mantle, a variety of creep processes can occur with dislocation creep dominating in the lower mantle and diffusional creep occasionally dominating in the upper mantle. However, there is a large transition region in creep processes between the upper and lower mantle and even within each section, creep properties can change strongly with location and thus temperature and pressure. In the power law creep regions, the creep equation fitted to data with n = 3–4 is standard.[22]
The strength of olivine not only scales with its melting temperature, but also is very sensitive to water and silica content. The solidus depression by impurities, primarily Ca, Al, and Na, and pressure affects creep behavior and thus contributes to the change in creep mechanisms with location. While creep behavior is generally plotted as homologous temperature versus stress, in the case of the mantle it is often more useful to look at the pressure dependence of stress. Though stress is simple force over area, defining the area is difficult in geology. Equation 1 demonstrates the pressure dependence of stress. Since it is very difficult to simulate the high pressures in the mantle (1MPa at 300–400 km), the low pressure laboratory data is usually extrapolated to high pressures by applying creep concepts from metallurgy.[23]
Most of the mantle has homologous temperatures of 0.65–0.75 and experiences strain rates of per second. Stresses in mantle are dependent on density, gravity, thermal expansion coefficients, temperature differences driving convection, and distance convection occurs over, all of which give stresses around a fraction of 3-30MPa. Due to the large grain sizes (at low stresses as high as several mm), it is unlikely that Nabarro-Herring (NH) creep truly dominates. Given the large grain sizes, dislocation creep tends to dominate. 14 MPa is the stress below which diffusional creep dominates and above which power law creep dominates at 0.5Tm of olivine. Thus, even for relatively low temperatures, the stress diffusional creep would operate at is too low for realistic conditions. Though the power law creep rate increases with increasing water content due to weakening, reducing activation energy of diffusion and thus increasing the NH creep rate, NH is generally still not large enough to dominate. Nevertheless, diffusional creep can dominate in very cold or deep parts of the upper mantle. Additional deformation in the mantle can be attributed to transformation enhanced ductility. Below 400 km, the olivine undergoes a pressure induced phase transformation into spinel and can cause more deformation due to the increased ductility.[23] Further evidence for the dominance of power law creep comes from preferred lattice orientations as a result of deformation. Under dislocation creep, crystal structures reorient into lower stress orientations. This does not happen under diffusional creep, thus observation of preferred orientations in samples lends credence to the dominance of dislocation creep.[24]
A similar process of slow convection probably occurs (or occurred) in the interiors of other planets (e.g., Venus, Mars) and some satellites (e.g., Europa, Enceladus).
There are three great categories of rocks: igneous, sedimentary, and metamorphic. Most of the time, they're simple to tell apart. They are all connected in the endless rock cycle, moving from one form to another and changing shape, texture, and even chemical composition along the way. Igneous rocks form from the cooling of magmaor lava and compose much of the Earth's continental crust and nearly all of the oceanic crust.
Identifying Igneous Rocks
The key concept about all igneous rocks is that they were once hot enough to melt. The following traits are all related to that.
Because their mineral grains grew together tightly as the melt cooled, they are relatively strong rocks.
They're made of primary minerals that are mostly black, white, or gray. Any other colors they may have are pale in shade.
Their textures generally look like something that was baked in an oven. The even texture of coarse-grained granite is familiar from building stones or kitchen counters. Fine-grained lava may look like black bread (including gas bubbles) or dark peanut brittle (including larger crystals).
Origin
Igneous rocks (derived from the Latin word for fire, ignis) can have very different mineral backgrounds, but they all share one thing in common: they formed by the cooling and crystallization of a melt. This material may have been lava erupted at the Earth's surface, or magma (unerupted lava) at depths of up to a few kilometers, known as magma in deeper bodies.
Those three different settings create three main types of igneous rocks. Rock formed of lava is called extrusive, rock from shallow magma is called intrusive, and rock from deep magma is called plutonic. The deeper the magma, the slower it cools, and it forms larger mineral crystals.
Where They Form
Igneous rocks form at four main places on Earth:
At divergent boundaries, like mid-ocean ridges, plates drift apart and form gaps that are filled by magma.
Subduction zones occur whenever a dense oceanic plate is subducted underneath another oceanic or continental plate. Water from the descending oceanic crust lowers the melting point of the above mantle, forming magma that rises to the surface and forms volcanoes.
At continental-continental convergent boundaries, large landmasses collide, thickening and heating the crust to melting.
Hot spots, like Hawaii, form as the crust moves over a thermal plume rising from deep in the Earth. Hot spots form extrusive igneous rocks.
People commonly think of lava and magma as a liquid, like molten metal, but geologists find that magma is usually a mush — a partially-melted fluid loaded with mineral crystals. As it cools, magma crystallizes into a series of minerals, some of which crystallize sooner than others. As the minerals crystallize, they leave the remaining magma with a changed chemical composition. Thus, a body of magma evolves as it cools and also as it moves through the crust, interacting with other rocks.
Once magma erupts as lava, it freezes quickly and preserves a record of its history underground that geologists can decipher. Igneous petrology is a very complex field, and this article is only a bare outline.
Textures
The three types of igneous rocks differ in their textures, starting with the size of their mineral grains.
Extrusive rocks cool quickly (over periods of seconds to months) and have invisible or microscopic grains or an aphanitic texture.
Intrusive rocks cool more slowly (over thousands of years) and have visible grains of small to medium-size, or phaneritic texture.
Plutonic rocks cool over millions of years and can have grains as large as pebbles — even meters across.
Because they solidified from a fluid state, igneous rocks tend to have a uniform fabric without layers, and the mineral grains are packed together tightly. Think of the texture of something you would bake in the oven.
In many igneous rocks, large mineral crystals "float" in a fine-grained groundmass. The large grains are called phenocrysts, and rock with phenocrysts is called a porphyry — that is, it has a porphyritic texture. Phenocrysts are minerals that solidified earlier than the rest of the rock, and they are important clues to the rock's history.
Some extrusive rocks have distinctive textures.
Obsidian, formed when lava hardens quickly, has a glassy texture.
Pumice and scoria are volcanic froth, puffed up by millions of gas bubbles that give them a vesicular texture.
Tuff is a rock made entirely of volcanic ash, fallen from the air or avalanched down a volcano's sides. It has a pyroclastic texture.
Pillow lava is a lumpy formation created by extruding lava underwater.
Basalt, Granite, and More
Igneous rocks are classified by the minerals they contain. The main minerals in igneous rocks are hard, primary ones: feldspar, quartz, amphiboles, and pyroxenes (together called "dark minerals" by geologists), as well as olivine, along with the softer mineral mica. The two best-known igneous rock types are basalt and granite, which have distinctly different compositions and textures.
Basalt is the dark, fine-grained stuff of many lava flows and magma intrusions. Its dark minerals are rich in magnesium (Mg) and iron (Fe), hence basalt is called a "mafic" rock. It can be either extrusive or intrusive.
Granite is the light, coarse-grained rock formed at a depth that is exposed after deep erosion. It is rich in feldspar and quartz (silica) and hence is called a "felsic" rock. Therefore, granite is felsic and plutonic.
Basalt and granite account for the great majority of igneous rocks. Ordinary people, even ordinary geologists, use the names freely. Stone dealers call any plutonic rock "granite." But igneous petrologists use many more names. They generally talk about basaltic and granitic or granitoid rocks among themselves and out in the field, because it takes laboratory work to determine an exact rock type according to the official classifications. True granite and true basalt are narrow subsets of these categories.
A few of the less common igneous rock types can be recognized by non-specialists. For instance, a dark-colored plutonic mafic rock, the deep version of basalt, is called gabbro. A light-colored intrusive or extrusive felsic rock, the shallow version of granite, is called felsite or rhyolite. And there is a suite of ultramafic rocks with even more dark minerals and even less silica than basalt. Peridotite is the foremost of those.
Where Igneous Rocks Are Found
The deep seafloor (the oceanic crust) is made almost entirely of basaltic rocks, with peridotite underneath in the mantle. Basalts are also erupted above the Earth's great subduction zones, either in volcanic island arcs or along the edges of continents. However, continental magmas tend to be less basaltic and more granitic.
The continents are the exclusive home of granitic rocks. Nearly everywhere on the continents, no matter what rocks are on the surface, you can drill down and reach granitoid eventually. In general, granitic rocks are less dense than basaltic rocks, and thus the continents float higher than the oceanic crust on top of the ultramafic rocks of the Earth's mantle. The behavior and histories of granitic rock bodies are among geology's deepest and most intricate mysteries.
The Mantle
The Mantle is the second layer of the Earth. It is the biggest and takes up 84 percent of the Earth. In this section you will learn and more about how hot the mantle is, what it is made of, and someinteresting facts about the Mantle.
Sections
The mantle is divided into two sections. The Asthenosphere, the bottom layer of the mantle made of plastic like fluid and The Lithosphere the top part of the mantle made of a cold dense rock.
Temperature
The average temperature of the mantle is 3000° Celsius. The temperature of the mantle will become much hotter as you get closer to the Inner Core
The mantle is composed of silicates of iron and magnesium, sulphides and oxides of silicon and magnesium.
Thickness
The mantle is about 2900 km thick. It is the largest layer of the Earth, taking up 84% of the Earth.
Convection Currents
Convection currents happen inside the mantle and are caused by the continuous circular motion of rocks in the lithosphere being pushed down by hot molasses liquid from the asthenosphere. The rocks then melt and float up as molasses liquid because it is less dense and the rocks float down because it is more dense.
Crust
The Crust is our home, yet it is also not our home. The very top of the crust is where we live on but deeper down it is all dense rock and metal ores. In this section you will learn about what the Crust is made of, the temperature, the thickness and a few interesting facts about the crust.
The Crust is composed of mainly granite, basalt, and dioriterocks
Thickness
The Crust's thickness can vary from wherever you are. From a continent to the edge of the crust is about 60 km. From the bottom of the ocean to the edge of the crust is about 10 km.
Temperature
The Crust's temperature is different throughout the entire crust. The temperatures start at about 200°Celsius and can rise up to 400° Celsius.
The Always Moving Layer
In the Mantle there are Convection Currents which will be defined more in The Mantle section. These huge currents are causing the crust to constantly move. These movements will cause earthquakes and volcanoes to erupt. The moving of the crust is also known as The Theory Of Plate Tectonics.
The Lithosphere
The Lithosphere is the top half of the mantle. It is a cold dense rock. This rock will be forced into the liquid in the asthenosphere causing it to melt and be pushed back up again causing the melted rock to cool off and return to normal rock. This is called Convection Currents.
The Asthenosphere is the bottom half of the mantle. It is a hot plastic liquid that is a liquid molasses like substance. This hot molasses substance will be forced into the lithosphere causing the liquid to cool off into rock, only to be forced back into the asthenosphere causing the rock to melt. This is called Convection Currents.
The Outer Core
The Outer Core is the second to last layer of the Earth. It is a magma like liquid layer that surrounds the Inner Core and creates Earth's magnetic field. In this section you will learn about how Earth's magnetic field is created, how hot it is, how thick the Outer Core is and a few interesting facts about the Outer Core.
Temperature
The Outer Core is about 4000-5000 degrees Celsius. The Inner Core is so hot it causes all the metal in the Outer Core to melt into liquid magma.
The Outer Core is composed of iron and some nickel. There is very few rocks and iron and nickel ore left in the Outer Core because of the Inner Core melting all the metal into liquid magma
Thickness
The Outer Core is about 2200 km thick. It is the second largest layer and made entirely out of liquid magma.
Magnetism
Because the outer core moves around the inner core, Earth's magnetism is created.
The Inner Core
The Inner Core is the final layer of the Earth. It is a solid ball made of metal. To learn about what metal the Inner Core is made of, read this section about the Inner Core. You can also learn how hot the Inner Core is, how thick it is and some interesting facts about the Inner Core.
Thickness
The Inner Core is about 1250 km thickand is the second smallest layer of the Earth. Although it is one of the smallest, the Inner Core is also the hottest layer.
The Inner Core is a solid ball composed of an element named NiFe. Ni for Nickel and Fe for Ferrum also known as Iron.
Temperature
The Inner Core is about 5000-6000 degrees Celsius. It melts all metal ores in the Outer Core causing it to turn into liquid magma.
Theory Of Continental Drift
Alfred Wegner, A German Scientist, one day called his girlfriend and told her that the East edge of Africa looks like the Southwest/West edge of South America. He then let everyone know about his discovery, starting the Theory of Continental Drift.
Pangaea
Alfred Wegner was determined to prove to the world that the world was once a super continent called Pangea. He estimated that Pangaea existed around 245 million years ago around the Triassic Period. He believed that all continents fit together like puzzle pieces but eventually drifted apart creating today's continent formation.
Evidence Of Continental Drift
As the years went by, some scientists said that Pangaea could not have happened. The thought that every continent was at one time together is insane! Alfred knew he was right, he just had to prove it.
Plant And Animal Fossils
Plant fossils and animal fossils were a big part of Alfred's evidence. He found plant fossils that could only be found on land, in the middle of the ocean. He also discovered that there were the same animal fossils in North America and in Russia that could not swim, in both parts of the world.
Rock Formations
Alfred found that the rock in the Appalachian Mountains in North America is the same kind of rock found in a mountain range in Britain and Norway.
Earthquakes
Believe it or not, the ground is moving under you right this second! Every day the crust's Tectonic Plates are moving in all different directions, causing them to force one another under them, rip apart, or rip off chunks as they move past each other. This section is on earthquakes, everything you will need to know about how they work.
Seismology
Seismology is the study of earthquakes. Seism, the Romanian word for earthquake and ology meaning the study of.
Seismographs
The word seismograph derives from the greek words seismós meaning shaking, or in other words, earthquakes and gráphō,meaning draw. Seismographs measure and record the movement of the crust. After measuring and recording from at least 3 seismographs the epicentre of the earthquake can be found. The Richter Scale though, measures the Magnitude of an earthquake. Magnitude is the power of the earthquake.
Seismic Wave
Seismic waves are the waves released by a rock or plate after they break. When a rock or seismic wave plate breaks, it releases energy called seismic waves. There are three types of Seismic Waves each with is own unique characteristics.
Primary Waves
Primary waves are the fastest and least felt of all earthquakes. Although they are the least felt they can make dishes rattle, and make a warning sign as to the bigger waves to come. These waves can also travel through any state of matter on earth.
Secondary Waves
Secondary waves are slower then Primary waves and again is used as a warning sign. Unlike Primary waves, Secondary waves can not travel through solids or liquids.
Surface Waves
Surface waves are the slowest out of all three waves but they are also the strongest. These huge waves travel in a rolling motion ripping up the ground a causing buildings to crumble.
Ring Of Fire
The Ring Of Fire is where a lot of the plate boundaries are. Because of this there are many earthquakes going on and lots of volcanoes being formed.
Normal Fault
Normal faults occur at divergent boundaries. When this type of fault happens one of the two plates moving apart will suddenly slide under the earth causing an earthquake.
Reverse Fault
Reverse faults occur at convergent boundaries. This type of fault causes the rock/plate above the fault line to be forced up causing more earthquakes.
Strike Slip Fault
Strike Slip Faults occur at transform boundaries and is caused by edges of a plate sticking out as the plates slide past each other. When the plates get caught on one another they push and pull until eventually the excess rock breaks off causing mass amount of pressure to be released making earthquakes.
Volcanoes
Volcanoes are mountains with craters at the top formed on Hot Spots or Divergent and Subduction Zones. Below that crater is a tube. Within that tube is hot magma flowing through to the top of the volcano. Gases within the volcano heat up, expand and finally cause the volcano to erupt. In this section you will learn about volcanoes, different types and some information about each of them.
Volcanology
Volcanology is the scientific study of volcanoes. Volcan is a word that comes from the Spanish language meaning "Volcano". Ology means "The study of".
Inside Volcano
From the outside a volcano may look like a simple mountain but the inside of a volcano is much different. Inside a volcano is a long tube leading up towards the top filled with lava. This is called a vent. The area in which the magma is held is called a magma reservoir. There are really only two parts to a volcano, the vent and magma reservoir. Although there are only two parts, there are different volcanoes to.
Shield Volcano
A shield volcano is a volcano that is huge in size, but not height. Take the Hawaiian islands for example, the islands are Shield Volcanoes. These volcanoes are mostly produced on plates with Hot Spots below them. Shield Volcanoes are made of lava flows that create a smooth sloped surface. Unlike the other volcanoes, shield volcanoes ooze magma rather than erupt violently.
Cinder Cone Volcanoes
Cinder cone volcanoes are the simplest type of volcano. The cinder cone has a bowl shaped crater and rarely rises more than 1 thousand feet. It is a steep sided volcano made of layers of ash and cinder that is crushed during eruption causing the tiny pieces to fall rapidly towards the ground, therefore creating more steep and rough ground. Unlike other volcanoes the cinder cone has only one vent leading towards the top of the volcano and has a very violent eruption that shoots out ash and lava.
Composite Volcano
Composite volcanoes are made of alternating layers of ash and lava. These volcanoes are a mix of the other two types of volcanoes, therefore it shoots out both lava and ash, along with oozing out magma. Composite volcanoes can have a very violent eruption or a calm ooze of magma. These volcanoes are the biggest in the world. Unlike other volcanoes, the composite has many vents rather than one.
Tsunamis
Earthquakes, believe it or not, can be even more destructive than they already are. When an earthquake appears out at sea the force of the tectonic waves is so powerful it will create a gigantic wave in the water. Not one of those waves that a surfer rides on, this wave is one that floods entire towns and causes many deaths. Tsunamis move towards the shore and slow down as they hit the shore causing water to flood a city and drag many people out to sea.
Tectonic Plate Boundaries
Tectonic plates are the plates beneath us in the mantle that are constantly moving. There are 3 types of tectonic plate boundaries. Divergent, Convergent, and Transform. These plate movements are caused by convection currents in the Mantle.
The Theory Of Plate Tectonics
German Scientist, Alfred Wegner, came up with The Theory Of Continental Drift. This theory soon became the Theory Of Plate Tectonics. The word Tectonic came from the Greek word tektonikos, basically meaning "construction". The word tectonics is talking about the general structure of the Earth.
Convergent Boundaries
A convergent boundary is a type of boundary where two plates meet together and start to push against one another. There are three types of convergent boundaries each with its own consequences.
Oceanic-Continental Convergence
The first type of convergent boundary is Oceanic-Continetal Convergence. This type of convergent boundary happens where an oceanic plate and a continental plate push together causing the oceanic plate to be forced under the continental plate into the mantle because the oceanic plate is thinner. This is called a subduction. When the oceanic plate is pushed under, it melts and turns into hot magma which burns its way through the continental plate, creating a volcano and causing many earthquakes.
Oceanic-Oceanic Convergence
Thenext type is Oceanic-Oceanic Convergence. This type of convergent boundary happens wheretwo oceanic plates push against one another, causing the colder, denser, older plate to buckle up and sink into the mantle. Hot magma comes from where the plate sank, creating new crust.
Continental-Continental Convergence
The final type of convergent boundary is Continental-Continental Convergence. This type of boundary happens where two continental plates collide and push up creating mountain ranges. Like colliding icebergs resist downward motion, the same thing happens with colliding continental plates, instead of moving down, they move up.
Divergent Boundaries
Divergent boundaries are plates that move away from one another causing underwater volcanoes and the spreading of the seafloor.
The Spreading Seafloor
Whendivergent plates move they move away from each other. When two oceanic plates move away from each other the seafloor starts to spread apart making more and more and more ground at the bottom of the sea.
Why do they move apart?
Deep in the mantle are huge convection currents. When they move upthe plates start to move apart because of the rising convection currents pushing on the bottom of the asthenosphere.
Transform Boundary
A Transform Boundary is a type of boundary that slides past one another instead of colliding or moving apart.
Transform Faults
Transform faults are found on the seafloor and are caused by surges of magma rising to make a ridge that is discontinued after its formed. While this happens the plates move in different directions sliding against each other making earthquakes, big and small. The most famous fault is the San Andreas fault
Minerals
Minerals are chemical elements or compounds that form a rock. There are many minerals in the world. There are so many that some minerals end up looking like each other! Geologists, Scientists who study rocks and minerals, use different tests to identify each mineral. These tests are: Hardness, Crystal Structure, Lustre, Colour, Streak, and Cleavage/Fracture.
Hardenss
Hardness is a mineral test involving Moh's Hardness Scale. The hardness scale measures a minerals resistance to scratching from 1-10, 1 being the softest and 10 being the hardest. Talc is one of the softest materials in the world, its so soft you can scratch it with your fingernail. Diamond is the hardest, only able to be scratched by diamond itself.
Crystal Structure
There are 7 Major Crystal Systems in the world. Crystal systems tell us about the shape of a mineral. The 7 systems are: Cubic, Tetragonal, Orthorhombic, Monoclinic, Triclinic, Hexagonal, and Trigonal. Cubic is in the shape of a cube, Tetragonal is a rectangle, Orthorhombic is a smaller type of rectangle, Monoclinic is even smaller, Triclinic is the smallest rectangle, Hexagonal is in the shape of a hexagon and Trigonal is a mix between a cubic and hexagonal shape.
Lustre
Lustre is a minerals shine. A minerals shine is put into to groups of lustre. Metallic lustre is where a mineral shines like polished metal and Non Metallic lustre is where a mineral does not shine like polished metal. Some examples of non metallic lustre are Waxy: where the mineral feels like wax, Pearly: where the mineral looks like pearl, Silky: where the mineral has a shine like silk, Greasy: it looks as if its covered in oil, Glassy: it has a reflection like glass, and Adamantine: the mineral shines like a diamond.
Color
Color is probably one of the least important mineral tests because there are so many minerals they all look the same! Color, although not very useful, tells us a minerals colour.
Streak
The streak of a mineral is the mark it will leave on an unglazed porcelain tile. The streak will help us identify different minerals because some minerals leave no streak, some leave a white streak and some will leave a coloured streak.
Cleavage And Fracture
Cleavage and Fracture is the test that tells us how a mineral breaks. If a mineral has cleavage it will break in nice smooth edges. If it has fracture then the mineral will break in rough jagged edges. An example of a mineral with cleavage is Mica (My-KA). This mineral breaks in smooth flat planes. An example of fracture is quartz. This mineral breaks in rough jagged edges.
Depends which part of the mantle you are talking about. The outer mantle consists of olivine and pyroxene (both coarse dark green crystalls) with flecs of black chromite. Deeper into the mantle the pressure and temperature conditions cause the minerals to recrystallize to perovskite and spinel. Most probably dark brown or greyish, possibly with red garnet. Deeper still, is so deep that nobody has ever sampled it, and we cannot be sure what the minerals will be. So the short answer is deep green or brownish crystalline rock.
Incidentally, you can visit the Sultanate of Oman where an anomalous raft of mantle material is exposed at the ground surface in what is known as the Semail ophiolite.
The mantle is the mostly-solid bulk of Earth’s interior. The mantle lies between Earth’s dense, super-heated core and its thin outer layer, the crust. The mantle is about 2,900 kilometers (1,802 miles) thick, and makes up a whopping 84% of Earth’s total volume. As Earth began to take shape about 4.5 billion years ago, iron and nickel quickly separated from otherrocks and minerals to form the core of the new planet. The moltenmaterial that surrounded the core was the early mantle. Over millions of years, the mantle cooled. Water trapped inside minerals erupted with lava, a process called “outgassing.” As more water was outgassed, the mantle solidified. The rocks that make up Earth’s mantle are mostly silicates—a wide variety of compounds that share a silicon and oxygen structure. Common silicates found in the mantle include olivine, garnet, and pyroxene. The other major type of rock found in the mantle is magnesium oxide. Other mantle elements include iron, aluminum, calcium, sodium, and potassium. The temperature of the mantle varies greatly, from 1000° Celsius (1832° Fahrenheit) near itsboundary with the crust, to 3700° Celsius (6692° Fahrenheit) near its boundary with the core. In the mantle, heat and pressuregenerally increase with depth. Thegeothermal gradient is a measurement of this increase. In most places, the geothermal gradient is about 25° Celsius per kilometer of depth (1° Fahrenheit per 70 feet of depth). The viscosity of the mantle also varies greatly. It is mostly solid rock, but more viscous at tectonic plate boundaries and mantle plumes. Mantle rocks there are soft and able to move plastically (over the course of millions of years) at great depth and pressure. The transfer of heat and material in the mantle helps determine the landscape of Earth. Activity in the mantle drives plate tectonics, contributing to volcanoes, seafloor spreading, earthquakes, andorogeny (mountain-building). The mantle is divided into several layers: the upper mantle, the transition zone, the lower mantle, and D” (D double-prime), the strange region where the mantle meets the outer core.
Upper Mantle The upper mantle extends from the crust to a depth of about 410 kilometers (255 miles). The upper mantle is mostly solid, but its moremalleable regions contribute totectonic activity. Two parts of the upper mantle are often recognized as distinctregions in Earth’s interior: the lithosphere and the asthenosphere. LithosphereThe lithosphere is the solid, outer part of the Earth, extending to a depth of about 100 kilometers (62 miles). The lithosphere includes both the crust and the brittle upper portion of the mantle. The lithosphere is both the coolest and the most rigid of Earth’s layers. The most well-known feature associated with Earth’s lithosphere is tectonic activity. Tectonic activity describes the interaction of the huge slabs of lithosphere calledtectonic plates. The lithosphere is divided into 15 major tectonic plates: the North American, Caribbean, South American, Scotia, Antarctic, Eurasian, Arabian, African, Indian, Philippine, Australian, Pacific, Juan de Fuca, Cocos, and Nazca. The division in the lithosphere between the crust and the mantle is called the Mohorovicic discontinuity, or simply the Moho. The Moho does not exist at a uniform depth, because not all regions of Earth are equally balanced in isostatic equilibrium.Isostasy describes the physical, chemical, and mechanical differences that allow the crust to “float” on the sometimes more malleable mantle. The Moho is found at about 8 kilometers (5 miles) beneath the ocean and about 32 kilometers (20 miles) beneath continents. Different types of rocks distinguishlithospheric crust and mantle. Lithospheric crust is characterized by gneiss (continental crust) and gabbro (oceanic crust). Below the Moho, the mantle is characterized by peridotite, a rock mostly made up of the minerals olivine and pyroxene. AsthenosphereThe asthenosphere is the denser, weaker layer beneath the lithospheric mantle. It lies between about 100 kilometers (62 miles) and 410 kilometers (255 miles) beneath Earth’s surface. The temperature and pressure of the asthenosphere are so high that rocks soften and partly melt, becoming semi-molten. The asthenosphere is much moreductile than either the lithosphere or lower mantle. Ductility measures a solid material’s ability to deform or stretch under stress. The asthenosphere is generally more viscous than the lithosphere, and the lithosphere-asthenosphere boundary (LAB) is the point wheregeologists and rheologists—scientists who study the flow of matter—mark the difference in ductility between the two layers of the upper mantle. The very slow motion of lithospheric plates “floating” on the asthenosphere is the cause of plate tectonics, a process associated with continental drift, earthquakes, the formation of mountains, and volcanoes. In fact, the lava that erupts from volcanicfissures is actually the asthenosphere itself, melted intomagma. Of course, tectonic plates are not really floating, because the asthenosphere is not liquid. Tectonic plates are only unstable at their boundaries and hot spots.
Transition Zone From about 410 kilometers (255 miles) to 660 kilometers (410 miles) beneath Earth’s surface, rocks undergo radicaltransformations. This is the mantle’s transition zone. In the transition zone, rocks do not melt or disintegrate. Instead, theircrystalline structure changes in important ways. Rocks become much, much more dense. The transition zone prevents large exchanges of material between the upper and lower mantle. Some geologists think that the increased density of rocks in the transition zone prevents subducted slabs from the lithosphere from falling further into the mantle. These huge pieces of tectonic plates stall in the transition zone for millions of years before mixing with other mantle rock and eventually returning to the upper mantle as part of the asthenosphere, erupting as lava, becoming part of the lithosphere, or emerging as new oceanic crust at sites of seafloor spreading. Some geologists and rheologists, however, think subducted slabs can slip beneath the transition zone to the lower mantle. Other evidence suggests that the transition layer ispermeable, and the upper and lower mantle exchange some amount of material. WaterPerhaps the most important aspectof the mantle’s transition zone is itsabundance of water. Crystals in the transition zone hold as much water as all the oceans on Earth’s surface. Water in the transition zone is not “water” as we know it. It is not liquid, vapor, solid, or even plasma. Instead, water exists as hydroxide.Hydroxide is an ion of hydrogen and oxygen with a negative charge. In the transition zone, hydroxide ions are trapped in the crystalline structure of rocks such as ringwoodite and wadsleyite. These minerals are formed from olivine at very high temperatures and pressure. Near the bottom of the transition zone, increasing temperature and pressure transform ringwoodite and wadsleyite. Their crystal structures are broken and hydroxide escapes as “melt.” Melt particles flow upwards, toward minerals that can hold water. This allows the transition zone to maintain a consistent reservoir of water. Geologists and rheologists think that water entered the mantle from Earth’s surface during subduction.Subduction is the process in which a dense tectonic plate slips or melts beneath a more buoyant one. Most subduction happens as an oceanic plate slips beneath a less-dense plate. Along with the rocks and minerals of the lithosphere, tons of water and carbon are also transported to the mantle. Hydroxide and water are returned to the upper mantle, crust, and even atmosphere through mantle convection, volcanic eruptions, and seafloor spreading.
Lower Mantle The lower mantle extends from about 660 kilometers (410 miles) to about 2,700 kilometers (1,678 miles) beneath Earth’s surface. The lower mantle is hotter and denser than the upper mantle and transition zone. The lower mantle is much less ductile than the upper mantle and transition zone. Although heat usually corresponds to softening rocks, intense pressure keeps the lower mantle solid. Geologists do not agree about the structure of the lower mantle. Some geologists think that subducted slabs of lithosphere have settled there. Other geologists think that the lower mantle is entirely unmoving and does not even transfer heat by convection.
D Double-Prime (D’’) Beneath the lower mantle is a shallow region called D'', or “d double-prime.” In some areas, D’’ is a nearly razor-thin boundary with the outer core. In other areas, D’’ has thick accumulations of iron and silicates. In still other areas, geologists and seismologists have detected areas of huge melt. The unpredictable movement of materials in D’’ is influenced by the lower mantle and outer core. The iron of the outer core influences the formation of a diapir, a dome-shaped geologic feature (igneous intrusion) where more fluidmaterial is forced into brittle overlying rock. The iron diapir emits heat and may release a huge, bulging pulse of either material or energy—just like a Lava Lamp. This energy blooms upward, transferring heat to the lower mantle and transition zone, and maybe even erupting as a mantle plume. At the base of the mantle, about 2,900 kilometers (1,802 miles) below the surface, is the core-mantle boundary, or CMB. This point, called the Gutenberg discontinuity, marks the end of the mantle and the beginning of Earth’s liquid outer core.
Mantle Convection Mantle convection describes the movement of the mantle as it transfers heat from the white-hot core to the brittle lithosphere. The mantle is heated from below, cooled from above, and its overall temperature decreases over long periods of time. All these elements contribute to mantle convection. Convection currents transfer hot, buoyant magma to the lithosphere at plate boundaries and hot spots.Convection currents also transfer denser, cooler material from the crust to Earth’s interior through the process of subduction. Earth's heat budget, which measures the flow of thermalenergy from the core to the atmosphere, is dominated by mantle convection. Earth’s heat budget drives most geologic processes on Earth, although its energy output is dwarfed by solar radiation at the surface. Geologists debate whether mantle convection is “whole” or “layered.”Whole-mantle convectiondescribes a long, long recycling process involving the upper mantle, transition zone, lower mantle, and even D’’. In this model, the mantle convects in a single process. A subducted slab of lithosphere may slowly slip into the upper mantle and fall to the transition zone due to its relative density and coolness. Over millions of years, it may sink further into the lower mantle. Convection currents may then transport the hot, buoyant material in D’’ back through the other layers of the mantle. Some of that material may even emerge as lithosphere again, as it is spilled onto the crust through volcanic eruptions or seafloor spreading. Layered-mantle convectiondescribes two processes. Plumes of superheated mantle material may bubble up from the lower mantle and heat a region in the transition zone before falling back. Above the transition zone, convection may be influenced by heat transferred from the lower mantle as well as discreteconvection currents in the upper mantle driven by subduction and seafloor spreading. Mantle plumes emanating from the upper mantle may gush up through the lithosphere as hot spots. Mantle PlumesA mantle plume is an upwelling of superheated rock from the mantle. Mantle plumes are the likely cause of “hot spots,” volcanic regions not created by plate tectonics. As a mantle plume reaches the upper mantle, it melts into a diapir. This molten material heats the asthenosphere and lithosphere, triggering volcanic eruptions. These volcanic eruptions make a minor contribution to heat loss from Earth’s interior, although tectonic activity at plate boundaries is the leading cause of such heat loss. The Hawaiian hot spot, in the middle of the North Pacific Ocean, sits above a likely mantle plume. As the Pacific plate moves in a generally northwestern motion, the Hawaiian hot spot remains relatively fixed. Geologists think this has allowed the Hawaiian hot spot to create a series of volcanoes, from the 85-million-year-old Meiji Seamount near Russia’s Kamchatka Peninsula, to the Loihi Seamount, a submarine volcano southeast of the “Big Island” of Hawaii. Loihi, a mere 400,000 years old, will eventually become the newest Hawaiian island. Geologists have identified two so-called “superplumes.” These superplumes, or large low shear velocity provinces (LLSVPs), have their origins in the melt material of D’’. The Pacific LLSVP influences geology throughout most of the southern Pacific Ocean (including the Hawaiian hot spot). The African LLSVP influences the geology throughout most of southern and western Africa. Geologists think mantle plumes may be influenced by many different factors. Some may pulse, while others may be heated continually. Some may have a single diapir, while others may have multiple “stems.” Some mantle plumes may arise in the middle of a tectonic plate, while others may be “captured” by seafloor spreading zones. Some geologists have identified more than a thousand mantle plumes. Some geologists think mantle plumes don’t exist at all. Until tools and technology allow geologists to more thoroughly explore the mantle, the debate will continue.
Exploring the Mantle The mantle has never been directly explored. Even the mostsophisticated drilling equipmenthas not reached beyond the crust. Drilling all the way down to the Moho (the division between the Earth's crust and mantle) is an important scientific milestone, but despite decades of effort, nobody has yet succeeded. In 2005, scientists with the Integrated Ocean Drilling Project drilled 1,416 meters (4,644 feet) below the North Atlantic seafloor and claimed to have come within just 305 meters (1,000 feet) of the Moho. XenolithsMany geologists study the mantle by analyzing xenoliths. Xenoliths are a type of intrusion—a rock trapped inside another rock. The xenoliths that provide the most information about the mantle are diamonds. Diamonds form under very unique conditions: in the upper mantle, at least 150 kilometers (93 miles) beneath the surface. Above depth and pressure, the carbon crystallizes as graphite, not diamond. Diamonds are brought to the surface in explosive volcanic eruptions, forming “diamond pipes” of rocks called kimberlites and lamprolites. The diamonds themselves are of less interest to geologists than the xenoliths some contain. These intrusions are minerals from the mantle, trapped inside the rock-hard diamond. Diamond intrusions have allowed scientists to glimpse as far as 700 kilometers (435 miles) beneath Earth’s surface—the lower mantle. Xenolith studies have revealed that rocks in the deep mantle are most likely 3-billion-year old slabs of subducted seafloor. The diamond intrusions include water, oceansediments, and even carbon. Seismic WavesMost mantle studies are conducted by measuring the spread of shock waves from earthquakes, calledseismic waves. The seismic waves measured in mantle studies are called body waves, because these waves travel through the body of the Earth. The velocity of body waves differs with density, temperature, and type of rock. There are two types of body waves: primary waves, or P-waves, and secondary waves, or S-waves. P-waves, also called pressure waves, are formed by compressions. Sound waves are P-waves—seismic P-waves are just far too low a frequency for people to hear.S-waves, also called shear waves, measure motion perpendicular to the energy transfer. S-waves are unable to transmit through fluids or gases. Instruments placed around the world measure these waves as they arrive at different points on the Earth’s surface after an earthquake. P-waves (primary waves) usually arrive first, while s-waves arrive soon after. Both body waves “reflect” off different types of rocks in different ways. This allows seismologists to identify different rocks present in Earth’s crust and mantle far beneath the surface. Seismic reflections, for instance, are used to identify hidden oil deposits deep below the surface. Sudden, predictable changes in the velocities of body waves are called “seismic discontinuities.” The Moho is a discontinuity marking the boundary of the crust and upper mantle. The so-called “410-kilometer discontinuity” marks the boundary of the transition zone. The Gutenberg discontinuity is more popularly known as the core-mantle boundary (CMB). At the CMB, S-waves, which can’t continue in liquid, suddenly disappear, and P-waves are strongly refracted, or bent. This alerts seismologists that the solid and molten structure of the mantle has given way to the fiery liquid of the outer core. Mantle MapsCutting-edge technology has allowed modern geologists and seismologists to produce mantle maps. Most mantle maps display seismic velocities, revealing patterns deep below Earth’s surface. Geoscientists hope that sophisticated mantle maps can plot the body waves of as many as 6,000 earthquakes withmagnitudes of at least 5.5. These mantle maps may be able to identify ancient slabs of subducted material and the precise position and movement of tectonic plates. Many geologists think mantle maps may even provide evidence for mantle plumes and their structure.
The Earth's surface lithosphere rides atop the asthenosphere and the two form the components of the upper mantle. The lithosphere is divided into a number of plates that are continuously being created and consumed at their opposite plate boundaries. Accretion occurs as mantle is added to the growing edges of a plate, associated with seafloor spreading. This hot added material cools down by conduction and convection of heat. At the consumption edges of the plate, the material has thermally contracted to become dense, and it sinks under its own weight in the process of subduction usually at an ocean trench.[3]
The subducted oceanic crust triggers volcanism, although the basic mechanisms are varied. Volcanism may occur due to processes that add buoyancy to partially melted mantle, which would cause upward flow of the partial melt due to decrease in its density. Secondary convection may cause surface volcanism as a consequence of intraplate extension[6] and mantle plumes.[7]
Mantle convection causes tectonic plates to move around the Earth's surface.[8] It seems to have been much more active during the Hadeanperiod, resulting in gravitational sorting of heavier molten iron, nickel, and sulphides to the core and lighter silicate minerals to the mantle.
Earth cross-section showing location of upper (3) and lower (5) mantle
Earth's temperature vs depth. Dashed curve: layered mantle convection. Solid curve: whole-mantle convection.[7]
A superplume generated by cooling processes in the mantle.[9]
During the late 20th century, there was significant debate within the geophysics community as to whether convection is likely to be "layered" or "whole".[10][11] Although elements of this debate still continue, results from seismic tomography, numerical simulations of mantle convection and examination of Earth's gravitational field are all beginning to suggest the existence of 'whole' mantle convection, at least at the present time. In this model, cold, subducting oceanic lithosphere descends all the way from the surface to the core–mantle boundary (CMB) and hot plumes rise from the CMB all the way to the surface.[12] This picture is strongly based on the results of global seismic tomography models, which typically show slab and plume-like anomalies crossing the mantle transition zone.
Although it is now well accepted that subducting slabs cross the mantle transition zone and descend into the lower mantle, debate about the existence and continuity of plumes persists, with important implications for the style of mantle convection. This debate is linked to the controversy regarding whether intraplate volcanism is caused by shallow, upper-mantle processes or by plumes from the lower mantle.[6]Many geochemistry studies have argued that the lavas erupted in intraplate areas are different in composition from shallow-derived mid-ocean ridge basalts (MORB). Specifically, they typically have elevated Helium-3 – Helium-4 ratios. Being a primordial nuclide, Helium-3 is not naturally produced on earth. It also quickly escapes from earth's atmosphere when erupted. The elevated He-3/He-4 ratio of Ocean Island Basalts (OIBs) suggest that they must be sources from a part of the earth that has not previously been melted and reprocessed in the same way as MORB source has been. This has been interpreted as their originating from a different, less well-mixed, region, suggested to be the lower mantle. Others, however, have pointed out that geochemical differences could indicate the inclusion of a small component of near-surface material from the lithosphere.
On Earth, the Rayleigh number for convection within Earth's mantle is estimated to be of order 107, which indicates vigorous convection. This value corresponds to whole mantle convection (i.e. convection extending from the Earth's surface to the border with the core). On a global scale, surface expression of this convection is the tectonic plate motions, and therefore has speeds of a few cm/a.[13][14][15] Speeds can be faster for small-scale convection occurring in low-viscosity regions beneath the lithosphere, and slower in the lowermost mantle where viscosities are larger. A single shallow convection cycle takes on the order of 50 million years, though deeper convection can be closer to 200 million years.[16]
Currently, whole mantle convection is thought to include broad-scale downwelling beneath the Americas and the Western Pacific, both regions with a long history of subduction, and upwelling flow beneath the central Pacific and Africa, both of which exhibit dynamic topography consistent with upwelling.[17] This broad-scale pattern of flow is also consistent with the tectonic plate motions, which are the surface expression of convection in the Earth's mantle and currently indicate degree-2 convergence toward the western Pacific and the Americas, and divergence away from the central Pacific and Africa.[18] The persistence of net tectonic divergence away from Africa and the Pacific for the past 250 Myr indicates the long-term stability of this general mantle flow pattern,[18] and is consistent with other studies [19][20][21] that suggest long-term stability of the LLSVP regions of the lowermost mantle that form the base of these upwellings.
Since the mantle is primarily composed of olivine ((Mg,Fe)2SiO4), the rheological characteristics of the mantle are largely those of olivine. Additionally, due to the varying temperatures and pressures between the lower and upper mantle, a variety of creep processes can occur with dislocation creep dominating in the lower mantle and diffusional creep occasionally dominating in the upper mantle. However, there is a large transition region in creep processes between the upper and lower mantle and even within each section, creep properties can change strongly with location and thus temperature and pressure. In the power law creep regions, the creep equation fitted to data with n = 3–4 is standard.[22]
The strength of olivine not only scales with its melting temperature, but also is very sensitive to water and silica content. The solidus depression by impurities, primarily Ca, Al, and Na, and pressure affects creep behavior and thus contributes to the change in creep mechanisms with location. While creep behavior is generally plotted as homologous temperature versus stress, in the case of the mantle it is often more useful to look at the pressure dependence of stress. Though stress is simple force over area, defining the area is difficult in geology. Equation 1 demonstrates the pressure dependence of stress. Since it is very difficult to simulate the high pressures in the mantle (1MPa at 300–400 km), the low pressure laboratory data is usually extrapolated to high pressures by applying creep concepts from metallurgy.[23]
Most of the mantle has homologous temperatures of 0.65–0.75 and experiences strain rates of per second. Stresses in mantle are dependent on density, gravity, thermal expansion coefficients, temperature differences driving convection, and distance convection occurs over, all of which give stresses around a fraction of 3-30MPa. Due to the large grain sizes (at low stresses as high as several mm), it is unlikely that Nabarro-Herring (NH) creep truly dominates. Given the large grain sizes, dislocation creep tends to dominate. 14 MPa is the stress below which diffusional creep dominates and above which power law creep dominates at 0.5Tm of olivine. Thus, even for relatively low temperatures, the stress diffusional creep would operate at is too low for realistic conditions. Though the power law creep rate increases with increasing water content due to weakening, reducing activation energy of diffusion and thus increasing the NH creep rate, NH is generally still not large enough to dominate. Nevertheless, diffusional creep can dominate in very cold or deep parts of the upper mantle. Additional deformation in the mantle can be attributed to transformation enhanced ductility. Below 400 km, the olivine undergoes a pressure induced phase transformation into spinel and can cause more deformation due to the increased ductility.[23] Further evidence for the dominance of power law creep comes from preferred lattice orientations as a result of deformation. Under dislocation creep, crystal structures reorient into lower stress orientations. This does not happen under diffusional creep, thus observation of preferred orientations in samples lends credence to the dominance of dislocation creep.[24]
Earth’s D” Layer
The D” layer, the lowermost portion of the mantle, sits just above the molten iron-rich outer core. Seismic observations have revealed a region with an intriguingly complex signature. This relatively thin layer, varying around 250 km in thickness, may hold the key to understanding how the core and mantle interact. The D” layer may also be where deep mantle plumes originate and where subducting slabs terminate. Some of the most puzzling seismic features include the splitting of shear-waves travelling through this layer and the presence of ultralow-velocity zones (ULVZ). ULVZs are thin (5- 40 km thick) patches in which the compressional and shear wave velocities are depressed by 5-10% and 10-30%, respectively, relative to the neighboring region.
We found that silicate post-perovskite (ppv), the major constituent in the D” layer, can be highly enriched in iron (Mao et al,PNAS 2004b, Mao et al,PNAS 2005), which has a dramatic effect on its physical and chemical properties (Mao et al, GRL 2006; Mao et al, AGU Monograph 2007). The ability of post-perovskite to absorb iron may explain the anomalously low velocities in ULVZs (Mao et al, Science 2006). More recently we studied the elastic anisotropy of this phase and found that it can exhibit large shear-wave splitting (Mao et al, PEPI 2009), consistent with seismic observations and theoretical calculations. We have been working on transport measurements of the thermal conductivity which is a key property for studying Earth’s thermal evolution and internal dynamics (Goncharov et al, PEPI 2010), and investigating how iron enrichment affects Fe-Mg partitioning and structural variation among a number of possible post-perovskite phases. The goal of these studies is to compile a comprehensive set of materials properties for each individual phase as well as for mineral assemblages, in order to deepen our understanding of how this boundary layer plays such a central role in global dynamics and the evolution of Earth.
The internal structure of the Earth is layered in spherical shells: an outer silicate solid crust, a highly viscousasthenosphereand 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.
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.
Earth's radial density distribution according to the preliminary reference earth model (PREM).[2]
Earth's gravity according to the preliminary reference earth model (PREM).[2] Comparison to approximations using constant and linear density for Earth's 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:[3]
Depth (km)
Layer
0–80
Lithosphere (locally varies between 5 and 200 km)
0–35
... Crust (locally varies between 5 and 70 km)
35–2,890
Mantle
80–220
... Asthenosphere
410–660
... Transition zone
35–660
... Upper mantle
660–2,890
... Lower mantle
2,740–2,890
... D″ layer
2,890–5,150
Outer core
5,150–6,360
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.
The Earth's crust ranges from 5–70 kilometres (3.1–43.5 mi)[4] in depth and is the outermost layer.[5] The thin parts are the oceanic crust, which underlie the ocean basins (5–10 km) and are composed of dense (mafic) iron magnesium silicateigneous 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.[6]
Earth's mantle extends to a depth of 2,890 km, making it the thickest layer of Earth.[7] The mantle is divided into upper and lower mantle,[8] which are separated by the transition zone.[9] The lowest part of the mantle next to the core-mantle boundary is known as the D″ (pronounced dee-double-prime) layer.[10] The pressure at the bottom of the mantle is ≈140 GPa (1.4 Matm).[11]The mantle is composed of silicate rocks that are rich in iron and magnesium relative to the overlying crust.[12] Although solid, the high temperatures within the mantle cause the silicate material to be sufficiently ductile that it can flow on very long timescales.[13]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 1021and 1024Pa·s, depending on depth.[14] In comparison, the viscosity of water is approximately 10−3Pa·s and that of pitch is 107Pa·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.[15]
The average density of Earth is 5.515 g/cm3.[16] 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.[17] 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.[18] Seismic measurements show that the core is divided into two parts, a "solid" inner core with a radius of ≈1,220 km[19]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.[20]
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.[21] 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.[22][23] The laser studies create plasma,[24] 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.[25][26]
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.[27][28]
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 matter that comprises Earth is connected in fundamental ways to matter of certain chondrite meteorites, and to matter of outer portion of the Sun.[30][31] 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.[32][33]
Recent evidence has suggested that the inner core of Earth may rotate slightly faster than the rest of the planet;[34] however, more recent studies in 2011[which?] found this hypothesis to be inconclusive. Options remain for the core which may be oscillatory in nature or a chaotic system.[citation needed] 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.[35][36]
There is more to the Earth than what we can see on the surface. In fact, if you were able to hold the Earth in your hand and slice it in half, you'd see that it has multiple layers. But of course, the interior of our world continues to hold some mysteries for us. Even as we intrepidly explore other worlds and deploy satellites into orbit, the inner recesses of our planet remains off limit from us.
However, advances in seismology have allowed us to learn a great deal about the Earth and the many layers that make it up. Each layer has its own properties, composition, and characteristics that affects many of the key processes of our planet. They are, in order from the exterior to the interior – the crust, the mantle, the outer core, and the inner core. Let's take a look at them and see what they have going on.
Like all terrestrial planets, the Earth's interior is differentiated. This means that its internal structure consists of layers, arranged like the skin of an onion. Peel back one, and you find another, distinguished from the last by its chemical and geological properties, as well as vast differences in temperature and pressure.
Our modern, scientific understanding of the Earth's interior structure is based on inferences made with the help of seismic monitoring. In essence, this involves measuring sound waves generated by earthquakes, and examining how passing through the different layers of the Earth causes them to slow down. The changes in seismic velocity cause refraction which is calculated (in accordance with Snell's Law) to determine differences in density.
These are used, along with 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, to determine what Earth's layers looks like. In addition, it is understood that the differences in temperature and pressure are due to leftover heat from the planet's initial formation, the decay of radioactive elements, and the freezing of the inner core due to intense pressure.
History of Study:
Since ancient times, human beings have sought to understand the formation and composition of the Earth. The earliest known cases were unscientific in nature – taking the form of creation myths or religious fables involving the gods. However, between classical antiquity and the medieval period, several theories emerged about the origin of the Earth and its proper makeup.
Most of the ancient theories about Earth tended towards the "Flat-Earth" view of our planet's physical form. This was the view in Mesopotamian culture, where the world was portrayed as a flat disk afloat in an ocean. To the Mayans, the world was flat, and at it corners, four jaguars (known as bacabs) held up the sky. The ancient Persians speculated that the Earth was a seven-layered ziggurat (or cosmic mountain), while the Chinese viewed it as a four-side cube.
By the 6th century BCE, Greek philosophers began to speculate that the Earth was in fact round, and by the 3rd century BCE, the idea of a spherical Earth began to become articulated as a scientific matter. During the same period, the development of a geological view of the Earth also began to emerge, with philosophers understanding that it consisted of minerals, metals, and that it was subject to a very slow process of change.
However, it was not until the 16th and 17th centuries that a scientific understanding of planet Earth and its structure truly began to advance. In 1692, Edmond Halley (discoverer of Halley's Comet) proposed what is now known as the "Hollow-Earth" theory. In a paper submitted to Philosophical Transactions of Royal Society of London, he put forth the idea of Earth consisting of a hollow shell about 800 km thick (~500 miles).
Between this and an inner sphere, he reasoned there was an air gap of the same distance. To avoid collision, he claimed that the inner sphere was held in place by the force of gravity. The model included two inner concentric shells around an innermost core, corresponding to the diameters of the planets Mercury, Venus, and Mars respectively.
Halley's construct was a method of accounting for the values of the relative density of Earth and the Moon that had been given by Sir Isaac Newton, in his Philosophiæ Naturalis Principia Mathematica (1687) – which were later shown to be inaccurate. However, his work was instrumental to the development of geography and theories about the interior of the Earth during the 17th and 18th centuries.
Another important factor was the debate during the 17th and 18th centuries about the authenticity of the Bible and the Deluge myth. This propelled scientists and theologians to debate the true age of the Earth, and compelled the search for evidence that the Great Flood had in fact happened. Combined with fossil evidence, which was found within the layers of the Earth, a systematic basis for identifying and dating the Earth's strata began to emerge.
The development of modern mining techniques and growing attention to the importance of minerals and their natural distribution also helped to spur the development of modern geology. In 1774, German geologist Abraham Gottlob Werner published Von den äusserlichen Kennzeichen der Fossilien (On the External Characters of Minerals) which presented a detailed system for identifying specific minerals based on external characteristics.
In 1741, the National Museum of Natural History in France created the first teaching position designated specifically for geology. This was an important step in further promoting knowledge of geology as a science and in recognizing the value of widely disseminating such knowledge. And by 1751, with the publication of the Encyclopédie by Denis Diderot, the term "geology" became an accepted term.
By the 1770s, chemistry was starting to play a pivotal role in the theoretical foundation of geology, and theories began to emerge about how the Earth's layers were formed. One popular idea had it that liquid inundation, like the Biblical Deluge, was responsible for creating all the geological strata. Those who accepted this theory became known popularly as the Diluvianists or Neptunists.
Another thesis slowly gained currency from the 1780s forward, which stated that instead of water, strata had been formed through heat (or fire). Those who followed this theory during the early 19th century referred to this view as Plutonism, which held that the Earth formed gradually through the solidification of molten masses at a slow rate. These theories together led to the conclusion that the Earth was immeasurably older than suggested by the Bible.
In the early 19th century, the mining industry and Industrial Revolution stimulated the rapid development of the concept of the stratigraphic column – that rock formations were arranged according to their order of formation in time. Concurrently, geologists and natural scientists began to understand that the age of fossils could be determined geologically (i.e. that the deeper the layer they were found in was from the surface, the older they were).
Computer simulation of the Earth’s field in a period of normal polarity between reversals. Credit: science.nasa.gov
During the imperial period of the 19th century, European scientists also had the opportunity to conduct research in distant lands. One such individual was Charles Darwin, who had been recruited by Captain FitzRoy of the HMS Beagle to study the coastal land of South America and give geological advice.
Darwin's discovery of giant fossils during the voyage helped to establish his reputation as a geologist, and his theorizing about the causes of their extinction led to his theory of evolution by natural selection, published in On the Origin of Species in 1859.
During the 19th century, the governments of several countries including Canada, Australia, Great Britain and the United States funded geological surveying that would produce geological maps of vast areas of the countries. By this time, the scientific consensus established the age of the Earth in terms of millions of years, and the increase in funding and the development of improved methods and technology helped geology to move farther away from dogmatic notions of the Earth's age and structure.
By the early 20th century, the development of radiometric dating (which is used to determine the age of minerals and rocks), provided the necessary the data to begin getting a sense of the Earth's true age. By the turn of the century, geologists now believed the Earth to be 2 billion years old, which opened doors for theories of continental movement during this vast amount of time.
In 1912, Alfred Wegener proposed the theory of Continental Drift, which suggested that the continents were joined together at a certain time in the past and formed a single landmass known as Pangaea. In accordance with this theory, the shapes of continents and matching coastline geology between some continents indicated they were once attached together.
Research into the ocean floor also led directly to the theory of Plate Tectonics, which provided the mechanism for Continental Drift. Geophysical evidence suggested lateral motion of continents and that oceanic crust is younger than continental crust. This geophysical evidence also spurred the hypothesis of paleomagnetism, the record of the orientation of the Earth's magnetic field recorded in magnetic minerals.
Model of a flat Earth, with the continents modeled in a disk-shape and Antarctica as an ice wall. Credit: Wikipedia Commons
Then there was the development of seismology, the study of earthquakes and the propagation of elastic waves through the Earth or through other planet-like bodies, in the early 20th century. By measuring the time of travel of refracted and reflected seismic waves, scientists were able to gradually infer how the Earth was layered and what lay deeper at its core.
For example, in 1910, Harry Fielding Ried put forward the "elastic rebound theory", based on his studies of the 1906 San Fransisco earthquake. This theory, which stated that earthquakes occur when accumulated energy is released along a fault line, was the first scientific explanation for why earthquakes happen, and remains the foundation for modern tectonic studies.
Then in 1926, English scientist Harold Jeffreys claimed that below the crust, the core of the Earth is liquid, based on his study of earthquake waves. And then in 1937, Danish seismologist Inge Lehmann went a step further and determined that within the earth's liquid outer core, there is a solid inner core.
By the latter half of the 20th century, scientists developed a comprehensive theory of the Earth's structure and dynamics had formed. As the century played out, perspectives shifted to a more integrative approach, where geology and Earth sciences began to include the study of the Earth's internal structure, atmosphere, biosphere and hydrosphere into one.
This was assisted by the development of space flight, which allowed for Earth's atmosphere to be studied in detail, as well as photographs taken of Earth from space. In 1972, the Landsat Program, a series of satellite missions jointly managed by NASA and the U.S. Geological Survey, began supplying satellite images that provided geologically detailed maps, and have been used to predict natural disasters and plate shifts.
Layers:
The Earth can be divided into one of two ways – mechanically or chemically. Mechanically – or rheologically, meaning the study of liquid states – it can be divided into the lithosphere, asthenosphere, mesospheric mantle, outer core, and the inner core. But chemically, which is the more popular of the two, it can be divided into the crust, the mantle (which can be subdivided into the upper and lower mantle), and the core – which can also be subdivided into the outer core, and inner core.
The inner core is solid, the outer core is liquid, and the mantle is solid/plastic. This is due to 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, the nickel-iron alloys and silicates are cool enough to be solid. In the upper mantle, the silicates are generally solid but localized regions of melt exist, leading to limited viscosity.
In contrast, the lower mantle is under tremendous pressure and therefore has a lower viscosity than the upper mantle. The metallic nickel–iron outer core is liquid because of the high temperature. However, the intense pressure, which increases towards the inner core, dramatically changes the melting point of the nickel–iron, making it solid.
The Earth’s Tectonic Plates. Credit: msnucleus.org
The differentiation between these layers is due to processes that took place during the early stages of Earth's formation (ca. 4.5 billion years ago). At this time, melting would have caused denser substances to sink toward the center while less-dense materials would have migrated to the crust. The core is thus believed to largely be composed of iron, along with nickel and some lighter elements, whereas less dense elements migrated to the surface along with silicate rock.
The Earth’s layers (strata) shown to scale. Credit: pubs.usgs.gov
Crust:
The crust is the outermost layer of the planet, the cooled and hardened part of the Earth that ranges in depth from approximately 5-70 km (~3-44 miles). This layer makes up only 1% of the entire volume of the Earth, though it makes up the entire surface (the continents and the ocean floor).
The thinner parts are the oceanic crust, which underlies the ocean basins at a depth of 5-10 km (~3-6 miles), while the thicker crust is the continental crust. Whereas the oceanic crust is composed of dense material such as iron magnesium silicate igneous rocks (like basalt), the continental crust is less dense and composed of sodium potassium aluminum silicate rocks, like granite.
The uppermost section of the mantle (see below), together with the crust, constitutes the lithosphere – an irregular layer with a maximum thickness of perhaps 200 km (120 mi). Many rocks now making up Earth's crust formed less than 100 million (1×108) years ago. However, the oldest known mineral grains are 4.4 billion (4.4×109) years old, indicating that Earth has had a solid crust for at least that long.
Upper Mantle:
The mantle, which makes up about 84% of Earth's volume, is predominantly solid, but behaves as a very viscous fluid in geological time. The upper mantle, which starts at the "Mohorovicic Discontinuity" (aka. the "Moho" – the base of the crust) extends from a depth of 7 to 35 km (4.3 to 21.7 mi) downwards to a depth of 410 km (250 mi). The uppermost mantle and the overlying crust form the lithosphere, which is relatively rigid at the top but becomes noticeably more plastic beneath.
Compared to other strata, much is known about the upper mantle, thanks to seismic studies and direct investigations using mineralogical and geological surveys. Movement in the mantle (i.e. convection) is expressed at the surface through the motions of tectonic plates. Driven by heat from deeper in the interior, this process is responsible for Continental Drift, earthquakes, the formation of mountain chains, and a number of other geological processes.
The mantle is also chemically distinct from the crust, in addition to being different in terms of rock types and seismic characteristics. This is due in large part to the fact that the crust is made up of solidified products derived from the mantle, where the mantle material is partially melted and viscous. This causes incompatible elements to separate from the mantle, with less dense material floating upward and solidifying at the surface.
Illustration of Edmond Halley’s model of a Hallow Earth, one that was made up of concentric spheres. Credit: Wikipedia Commons/Rick Manning
The crystallized melt products near the surface, upon which we live, are typically known to have a lower magnesium to iron ratio and a higher proportion of silicon and aluminum. These changes in mineralogy may influence mantle convection, as they result in density changes and as they may absorb or release latent heat as well.
In the upper mantle, temperatures range between 500 to 900 °C (932 to 1,652 °F). Between the upper and lower mantle, there is also what is known as the transition zone, which ranges in depth from 410-660 km (250-410 miles).
Lower Mantle:
The lower mantle lies between 660-2,891 km (410-1,796 miles) in depth. Temperatures in this region of the planet can reach over 4,000 °C (7,230 °F) at the boundary with the core, vastly exceeding the melting points of mantle rocks. However, due to the enormous pressure exerted on the mantle, viscosity and melting are very limited compared to the upper mantle. Very little is known about the lower mantle apart from that it appears to be relatively seismically homogeneous.
The internal structure of Earth. Credit: Wikipedia Commons/Kelvinsong
Outer Core:
The outer core, which has been confirmed to be liquid (based on seismic investigations), is 2300 km thick, extending to a radius of ~3,400 km. In this region, the density is estimated to be much higher than the mantle or crust, ranging between 9,900 and 12,200 kg/m3. The outer core is believed to be composed of 80% iron, along with nickel and some other lighter elements.
Denser elements, like lead and uranium, are either too rare to be significant or tend to bind to lighter elements and thus remain in the crust. The outer core is not under enough pressure to be solid, so it is liquid even though it has a composition similar to that of the inner core. The temperature of the outer core ranges from 4,300 K (4,030 °C; 7,280 °F) in the outer regions to 6,000 K (5,730 °C; 10,340 °F) closest to the inner core.
Because of its high temperature, the outer core exists in a low viscosity fluid-state that undergoes turbulent convection and rotates faster than the rest of the planet. This causes eddy currents to form in the fluid core, which in turn creates a dynamo effect that is believed to influence Earth's magnetic field. The average magnetic field strength in Earth's outer core is estimated to be 25 Gauss (2.5 mT), which is 50 times the strength of the magnetic field measured on Earth's surface.
Inner Core:
The growing importance of mining in the 17th and 18th centuries, particularly for precious metals, led to further developments in geology and Earth sciences. Credit: minerals.usgs.gov
Like the outer core, the inner core is composed primarily of iron and nickel and has a radius of ~1,220 km. Density in the core ranges between 12,600-13,000 kg/m3, which suggests that there must also be a great deal of heavy elements there as well – such as gold, platinum, palladium, silver and tungsten.
The temperature of the inner core is estimated to be about 5,700 K (~5,400 °C; 9,800 °F). The only reason why iron and other heavy metals can be solid at such high temperatures is because their melting temperatures dramatically increase at the pressures present there, which ranges from about 330 to 360 gigapascals.
Because the inner core is not rigidly connected to the Earth's solid mantle, the possibility that it rotates slightly faster or slower than the rest of Earth has long been considered. By observing changes in seismic waves as they passed through the core over the course of many decades, scientists estimate that the inner core rotates at a rate of one degree faster than the surface. More recent geophysical estimates place the rate of rotation between 0.3 to 0.5 degrees per year relative to the surface.
Recent discoveries also suggest that the solid inner core itself is composed of layers, separated by a transition zone about 250 to 400 km thick. This new view of the inner core, which contains an inner-inner core, posits that the innermost layer of the core measures 1,180 km (733 miles) in diameter, making it less than half the size of the inner core. It has been further speculated that while the core is composed of iron, it may be in a different crystalline structure that the rest of the inner core.
What's more, recent studies have led geologists to conjecture that the dynamics of deep interior is driving the Earth's inner core to expand at the rate of about 1 millimeter a year. This occurs mostly because the inner core cannot dissolve the same amount of light elements as the outer core.
Artist’s illustration of Earth’s core, inner core, and inner-inner core. Credit: Huff Post Science
The freezing of liquid iron into crystalline form at the inner core boundary produces residual liquid that contains more light elements than the overlying liquid. This in turn is believed to cause the liquid elements to become buoyant, helping to drive convection in the outer core.
This growth is therefore likely to play an important role in the generation of Earth's magnetic field by dynamo action in the liquid outer core. It also means that the Earth's inner core, and the processes that drive it, are far more complex than previously thought!
Yes indeed, the Earth is a strange and mysteries place, titanic in scale as well as the amount of heat and energy that went into making it many billions of years ago. And like all bodies in our universe, the Earth is not a finished product, but a dynamic entity that is subject to constant change. And what we know about our world is still subject to theory and guesswork, given that we can't examine its interior up close.
As the Earth's tectonic plates continue to drift and collide, its interior continues to undergo convection, and its core continues to grow, who knows what it will look like eons from now? After all, the Earth was here long before we were, and will likely continue to be long after we are gone.
D″ layer is a thermal boundary layerin which heat is transferred from the core to the mantle.
The Earth
The Earth, our planet, is the third planet from the Sun and the only habitable planet in our solar system. The oldest material in the solar system is around 4.5 billion years old. Since the Earth was likely formed during the formation of the solar system, it is also 4.5 billion years old (Gya). After the Earth formed (accreted) it differentiated, meaning its interior separated into layers. The Earth’s interior is made of a spherical shells sort of like an onion. The radius of the Earth is 6371 km. It has an average density of 5.5 g/cm3.
Figure 1: Structure of the Earth's Interior
The structure of the Earth’s upper mantle can be derived from seismic waves. The main layers are the crust, the mantle and the core.
Let’s break them down into each section:
The Crust
The crust is the uppermost layer of the planet. It is between 5 and 80km thick. There are two types of crust, oceanic crust (found beneath the oceans) and continental crust. The oceanic crust is only 5-10km thick and made up mostly of basalt. The continental crust can be much thicker, up to 80km, and is made of less dense rocks such as silicate. The major element components of the crust are oxygen and silicon (Si) and thus the mantle is often referred to as the silicate mantle.
The Mantle
The mantle lies beneath the crust to a depth of about 2900km. The mantle has many layers within the upper and lower mantle. The upper layer is the lithosphere below which is the asthenosphere. The transition zone is the layer between the upper and lower mantle distinguished by the 410 km and 660 km discontinuities, as revealed by seismic evidence. The lower mantle has a D” layer just above the core-mantle boundary.
The Core
The Core is made up of two layers, the inner core and outer core. Seismic evidence tells us that the inner core is solid while the outer core is liquid. The inner core has a radius of 1 216 km and the total radius of the core is 3486 km. The core is composed mostly of iron (80%) and some nickel. The density of the inner ’solid’ core is between 9.9-12.2 g/cm3 and the outer core’s density is between 12.6-13 g/cm3.
It is through our study of the Earth that we can gain insight into the structure and composition of other planets. We also gather details from asteroids that fall to Earth from space. Using these asteroids we are able to learn about how the solar system looked when it was first forming. This helps us understand how the Earth and other planets looks when they were formed. By understanding what make up our planets, we can learn about why they behave in different ways.
The Dynamics of Earth
Convection is a term describing the flow of heat in a fluid that is driven by buoyancy derived from horizontal density gradients. Density gradients in the mantle are largely derived from horizontal temperature gradients (and also chemical/compositional horizontal gradients). In the thermal boundary layers (across which the temperature varies continuously from the surface value to the mean mantle temperature) this buoyancy causes instabilities, allowing fluid to leave the boundary layer and rise or fall throughout the system interior.
The mantle is a visco-elastic solid, meaning it behaves both viscously and elastically in response to a stress. The viscous nature of the mantle is evident in the slow creep of the mantle manifesting itself as plate tectonics on the surface of the Earth. The elastic nature of mantle rock is evident in the seafloor flexure around ocean island chains (e.g. Kearey, 2009). By assuming a fully elastic crustal layer overlying a fluid, the height of flexure in response to a load can be determined. These theoretical values can be compared with oceanic crust response to seafloor mounts to determine the elastic response of the mantle.
Heat is removed from the interior of a planet by thermal conduction as well as subsolidus convection. Subsolidus convection occurs from diffusion or dislocation creep in a solid material. The temperature difference between the interior and cooling surface of a planet maintains the thermal gradient necessary for convection.
Heat is the main source of energy driving convection in the mantle. Heat in the mantle is derived from internal sources (radioactive decay of the el- ements uranium, thorium and potassium), heat released from the core and secular cooling of the planet as a whole (residual heat left over from planetary formation and a higher production of radioactive heating in the past).
Mantle convection manifests itself at the surface of the Earth. Mid-ocean ridges correspond to the site of passive upwelling mantle material while ocean trenches correspond to the location of convective downwellings (subduction). The cycle of upwelling and downwelling convection helps recycle lithosphere into the mantle, producing new lithosphere at ridges and removing it at subduction zones. Figure 2 shows a depiction of a mantle convection cell, with a hot upwelling plume (red) and cold subducting slab (blue). Figure 1 shows a depiction of mantle convection with an upwelling plume, passive upwelling at a mid-ocean ridge and subducting slabs (downwellings). It also shows large-low shear velocity provinces, ultra-low-velocity-zones and areas of post-perovskite.
Figure 2: Convection cell in the mantle. Hot upwelling plume in red and cold downwelling (subducting) slab in blue.
The Earth’s top and bottom thermal boundary layers are influenced by strong chemical heterogeneity. Insert figure ‘convectioncell’ here The study of mantle convection is very important for understanding how the Earth functions, from how heat is lost from the interior to how the continental crust is replenished.
In this article (geography section), we discuss the interior of the earth. Understanding the basic structure of earth is very important to learn higher concepts well. Also, the origin of many phenomena like earthquakes, volcanoes, tsunami etc are linked with the structure of earth’s interior.
What should you understand about the interior of the earth?
It is not possible to know about the earth’s interior by direct observations because of the huge size and the changing nature of its interior composition.
It is an almost impossible distance for the humans to reach till the centre of the earth (The earth’s radius is 6,370 km).
Through mining and drilling operations we have been able to observe the earth’s interior directly only up to a depth of few kilometers.
The rapid increase in temperature below the earth’s surface is mainly responsible for setting a limit to direct observations inside the earth.
But still, through some direct and indirect sources, the scientists have a fair idea about how the earth’s interior look like.
Sources of Information about the interior of the earth
Direct Sources:
Rocks from mining area
Volcanic eruptions
Indirect Sources
By analyzing the rate of change of temperature and pressure from the surface towards the interior.
Meteors, as they belong to the same type of materials earth is made of.
Gravitation, which is greater near poles and less at the equator.
Gravity anomaly, which is the change in gravity value according to the mass of material, gives us information about the materials in the earth’s interior.
Magnetic sources.
Seismic Waves: the shadow zones of body waves (Primary and secondary waves) give us information about the state of materials in the interior.
It is the outermost solid part of the earth, normally about 8-40 kms thick.
It is brittle in nature.
Nearly 1% of the earth’s volume and 0.5% of earth’s mass are made of the crust.
The thickness of the crust under the oceanic and continental areas are different. Oceanic crust is thinner (about 5kms) as compared to the continental crust (about 30kms).
Major constituent elements of crust are Silica (Si) and Aluminium (Al) and thus, it is often termed as SIAL (Sometimes SIAL is used to refer Lithosphere, which is the region comprising the crust and uppermost solid mantle, also).
The mean density of the materials in the crust is 3g/cm3.
The discontinuity between the hydrosphere and crust is termed as the Conrad Discontinuity.
Mantle
The portion of the interior beyond the crust is called as the mantle.
The discontinuity between the crust and mantle is called as the Mohorovich Discontinuity or Moho discontinuity.
The mantle is about 2900kms in thickness.
Nearly 84% of the earth’s volume and 67% of the earth’s mass is occupied by the mantle.
The major constituent elements of the mantle are Silicon and Magnesium and hence it is also termed as SIMA.
The density of the layer is higher than the crust and varies from 3.3 – 5.4g/cm3.
The uppermost solid part of the mantle and the entire crust constitute the Lithosphere.
The asthenosphere (in between 80-200km) is a highly viscous, mechanically weak and ductile, deforming region of the upper mantle which lies just below the lithosphere.
The asthenosphere is the main source of magma and it is the layer over which the lithospheric plates/ continental plates move (plate tectonics).
The discontinuity between the upper mantle and the lower mantle is known as Repetti Discontinuity.
The portion of the mantle which is just below the lithosphere and asthenosphere, but above the core is called as Mesosphere.
Core
It is the innermost layer surrounding the earth’s centre.
The core is separated from the mantle by Guttenberg’s Discontinuity.
It is composed mainly of iron (Fe) and nickel (Ni) and hence it is also called as NIFE.
The core constitutes nearly 15% of earth’s volume and 32.5% of earth’s mass.
The core is the densest layer of the earth with its density ranges between 9.5-14.5g/cm3.
The Core consists of two sub-layers: the inner core and the outer core.
The inner core is in solid state and the outer core is in the liquid state (or semi-liquid).
The discontinuity between the upper core and the lower core is called as Lehmann Discontinuity.
Barysphere is sometimes used to refer the core of the earth or sometimes the whole interior.
Temperature, Pressure and Density of the Earth’s Interior
Temperature
A rise in temperature with increase in depth is observed in mines and deep wells.
These evidence along with molten lava erupted from the earth’s interior supports that the temperature increases towards the centre of the earth.
The different observations show that the rate of increase of temperature is not uniform from the surface towards the earth’s centre. It is faster at some places and slower at other places.
In the beginning, this rate of increase of temperature is at an average rate of 10C for every 32m increase in depth.
While in the upper 100kms, the increase in temperature is at the rate of 120C per km and in the next 300kms, it is 200C per km. But going further deep, this rate reduces to mere 100C per km.
Thus, it is assumed that the rate of increase of temperature beneath the surface is decreasingtowards the centre (do not confuse rate of increase of temperature with increase of temperature. Temperature is always increasing from the earth’s surface towards the centre).
The temperature at the centre is estimated to lie somewhere between 30000C and 50000C, may be that much higher due to the chemical reactions under high-pressure conditions.
Even in such a high temperature also, the materials at the centre of the earth are in solid state because of the heavy pressure of the overlying materials.
Pressure
Just like the temperature, the pressure is also increasing from the surface towards the centre of the earth.
It is due to the huge weight of the overlying materials like rocks.
It is estimated that in the deeper portions, the pressure is tremendously high which will be nearly 3 to 4 million times more than the pressure of the atmosphere at sea level.
At high temperature, the materials beneath will melt towards the centre part of the earth but due to heavy pressure, these molten materials acquire the properties of a solid and are probably in a plastic state.
Density
Due to increase in pressure and presence of heavier materials like Nickel and Iron towards the centre, the density of earth’s layers also gets on increasing towards the centre.
The average density of the layers gets on increasing from crust to core and it is nearly 14.5g/cm3 at the very centre.
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per second. Stresses in mantle are dependent on density, gravity, thermal expansion coefficients, temperature differences driving convection, and distance convection occurs over, all of which give stresses around a fraction of 3-30MPa. Due to the large grain sizes (at low stresses as high as several mm), it is unlikely that Nabarro-Herring (NH) creep truly dominates. Given the large grain sizes, dislocation creep tends to dominate. 14 MPa is the stress below which diffusional creep dominates and above which power law creep dominates at 0.5Tm of olivine. Thus, even for relatively low temperatures, the stress diffusional creep would operate at is too low for realistic conditions. Though the power law creep rate increases with increasing water content due to weakening, reducing activation energy of diffusion and thus increasing the NH creep rate, NH is generally still not large enough to dominate. Nevertheless, diffusional creep can dominate in very cold or deep parts of the upper mantle. Additional deformation in the mantle can be attributed to transformation enhanced ductility. Below 400 km, the olivine undergoes a pressure induced phase transformation into spinel and can cause more deformation due to the increased ductility.[23] Further evidence for the dominance of power law creep comes from preferred lattice orientations as a result of deformation. Under dislocation creep, crystal structures reorient into lower stress orientations. This does not happen under diffusional creep, thus observation of preferred orientations in samples lends credence to the dominance of dislocation creep.[24]
