3.1: Earth's Interior

What's Below the Surface

In order to understand the details of plate tectonic theory, it is essential to first understand the layers that make up Earth’s interior. Firsthand information about what is below the surface is very limited; most of what we know is inferred from data of Earth’s gravitational and magnetic fields, the movement of seismic waves through Earth, the geothermal heat measured at Earth’s surface, and inferences from the composition of meteorites. There are two ways in which we divide Earth into layers: by chemical composition and by physical properties (Figure \(\PageIndex{1}\)).

Figure \(\PageIndex{1}\): The two ways in which Earth’s layers can be defined. Left: Earth’s compositional layers include the crust, mantle, and core. Right: Earth’s mechanical layers include the lithosphere, mesosphere, outer core, and inner core. "Two different Views" by Nooreen Meghani is licensed under CC BY-NC-SA.

Layers Defined by Chemical Composition

The Earth is not made up of a single, uniform chemical composition. As the planet formed and continued to evolve, it differentiated into layers based on compositional and density differences, with its heavier elements sinking to its center and lighter elements remaining near its surface. This process resulted in the Earth developing roughly concentric layers with varying compositions, creating its crust, mantle, and core (Figure \(\PageIndex{1}\)).

The Crust

The outermost chemical layer of the Earth is the crust Figure \(\PageIndex{2}\). There are two types of crust: continental and oceanic.

The continental crust has a relatively low density because it has an average composition that is felsic. Although there are a range of different rock types that make up the continental crust, geologists generally refer to it as "granitic" in average composition (granite is a common example of a felsic rock and it has an average density of 2.7 g/cm³). The average thickness of continental crust is 25 to 70 km (16 to 43 mi), and its average elevation is 840 m (2,750 ft) above sea level.

Figure \(\PageIndex{2}\): Differences between continental and oceanic crust.  Continental crust is thicker and less dense compared to the thinner and more dense oceanic crust. "Crust" by USGS Earthquakes Hazards Program is in the public domain.

The oceanic crust has a higher density, reflecting the fact that it is primarily made of the mafic rock “basalt”. The density of basalt is ~3.0 g/cm³. Oceanic crust is also much thinner than continental crust, ranging from 7 to 10 km (4.3 to 6.2 mi), and because it is denser, it is generally found below sea level. Mafic is a chemical term for rock that is lower in silica and has a greater amount of iron and magnesium. The density of oceanic crust also increases as it gets older and colder.

Box \(\PageIndex{1}\): Why are Mountains Tall?

What holds up a mountain? Have you ever wondered why mountains are tall? Most people would say that rock is strong; strong enough to hold itself up, which is sometimes true, but not always. It depends on the size of the mountain or the mountain range. The strength of the rock is only a small part of the story. Mountains are supported in part by the strength of the rock, but mostly by the buoyancy of the rock below them.

First consider the strength of the rock, not the buoyancy of the rock below. As an analogy, consider the cover on a swimming pool to be like the strength of the rock and a house cat to be like a small mountain. As the cat walks across the pool cover it flexes, but continues to support the cat. This holds true for a small mountain. The rock in the lithosphere may slightly flex, but it supports the small mountain. Most mountains or mountain ranges, though, are not small, so the strength of the rock alone is not enough to hold the mountain up.

Then why do mountains even exist? The surface of Earth is anything but uniform in elevation, there are lots of mountains and basins and continents and oceans. These changes in elevation at the surface of Earth are all expressions of the relationship of the rocks in the lithosphere and the asthenosphere trying to achieve gravitational or isostatic equilibrium, a balance between the force of gravity pulling downwards on the rock, the buoyancy of the rock at depth, and the strength of the rock.

The rocks of the lithosphere vary in thickness and density - think of continental crust and oceanic crust as a broad example. The rocks of the continental crust are generally thicker and less dense, while the rocks of the oceanic crust are thinner and denser. Thin or thick, less or more dense, the rocks of the lithosphere actually float on the asthenosphere. The relative height (or lack of height) occurs as a balancing act.

Box Figure \(\PageIndex{1.1}\): Iceberg with submerged portion visible. The submerged portion is larger (more voluminous) that the exposed portion. "Iceberg in the Arctic with its underside exposed" by AWeith via Wikimedia Commons is licensed under CC BY-SA.

The first and the most important part of the balancing act is to consider how an ice cube floats in a glass of lemonade, or an iceberg floats in the ocean (Box Figure \(\PageIndex{1.1}\)). It is the ice below the surface of the ocean that is important for the iceberg to float. Gravity is trying to pull the iceberg down, but because the ice is less dense than the surrounding seawater, its buoyancy is making it rise. If it were to snow, the iceberg would become heavier and sink a bit more but would still be supported by the ice below the water level because that part of the iceberg is surrounded by denser seawater. On the other hand, if it were warmer, some of the ice might melt and flow into the sea. There would be less ice to support, so the iceberg would rise a little. Either way there is a balance, and isostatic or gravitational equilibrium is eventually achieved.

Now consider rocks. A mountain in the lithosphere floats like an iceberg (Box Figure \(\PageIndex{1.2}\)). It is made of rock rather than ice, and the asthenosphere is also made of rock. But just like ice and seawater, the rock in the lithosphere is less dense than the rock in the asthenosphere. However, unlike seawater, the asthenosphere is solid, but it’s a ductile solid that behaves like an extremely viscous liquid. Think of it as more like peanut butter than water. Peanut butter flows very slowly. If you tip over a jar of peanut butter, it will take hours for peanut butter to flow out of the jar. The rock of the asthenosphere by comparison would take thousands of years. However, because it can flow, even if the flow is very slow, it behaves like a very viscous fluid and the lithosphere can therefore float on the asthenosphere.

Box Figure \(\PageIndex{1.2}\): The isostatic relationship between the crust and the mantle. Left: a mountain and how its weight depresses the asthenosphere below and arrows showing how the asthenosphere moves. Right: how when the mountain is eroded away the lithosphere rebounds and the arrows indicate the motion of the asthenosphere. "Isostatic Relationships" by Steven Earle is licensed under CC BY 4.0.

Now a mountain or a mountain range stands tall, which makes it easy to erode. As the mountain erodes, and material is removed, the land may rise because the rock supporting it from below is adjusting to the removal of the eroded material. But as the land rises, the mountain also rises, so erosion continues, until perhaps there no longer is a mountain to erode. The opposite can also occur. Consider the sediment that eroded from the mountain. If it piles up somewhere else, the mass of sediment may cause the land surface to subside, as the rock below adjusts to the additional weight, and a basin may form. The Sierra Nevada and the Great Valley are California examples.

One last consideration, the isostatic equilibrium of an area is always slowly changing. This is because of the changes occurring to the rock at or near Earth’s surface and because the lithosphere is not exactly like a pool cover. Unlike a pool cover, its strength varies with temperature and depth. To complicate matters, the asthenosphere and lower mantle below are also constantly in motion because of mantle convection.

The Mantle

The mantle is the largest compositional layer by volume, extending from the base of the crust to a depth of about 2,900 km (1,802 mi). Most of what we know about the mantle is inferred from seismic wave analysis and the study of rocky meteorites. Ophiolites and xenoliths are also used to study conditions within the mantle and mantle composition. Ophiolites are pieces of oceanic lithosphere that have been uplifted and exposed at Earth’s surface (Figure \(\PageIndex{3}\)). In California, ophiolites are preserved in the Coast Ranges, the Klamath Mountains, and the foothills of the Sierra Nevada. The changes these rocks undergo provide information on the pressure and temperatures within the mantle. Mantle xenoliths are rocks from depth brought to the surface of Earth by volcanic eruptions. As magma moves upward through the mantle, it can break off pieces of the mantle it is rising through, and those pieces can be preserved as inclusions within the rock that forms as the lava from the volcanic eruption cools. Most xenoliths are made of peridotite, an ultramafic class of igneous rock (see Igneous Rocks). Because of this, geoscientists hypothesize that portions of the mantle are made of peridotite.

Figure \(\PageIndex{3}\): Idealized ophiolite sequence: Ultramafic mantle rocks, overlain by crustal gabbro, sheeted dikes and sills (diabase), pillow basalts, and marine sediments and marine rocks. These sequences are typically around 10 km (6 mi) thick. "ophiolite-cross-section.jpg" by Dexter Perkins is licensed under CC BY-NC-SA 4.0.

Mantle Xenolith 3D Model

The following 3D model is of a mantle xenolith preserved in basalt. As the basalt made its way to the surface, pieces of the mantle (the patches of green), became entrained in the rock.

This model by Ryan Hollister via Sketchfab is licensed under CC BY.

The boundary between the crust and the mantle is identified by a seismic discontinuity called the Mohorovičić Discontinuity (or Moho for short; Figure \(\PageIndex{4}\)). The discontinuity refers to the change in seismic wave velocities, which increase dramatically as seismic waves travel from the crust into the denser mantle. The change in wave direction and speed is caused by significant mineralogical differences between the crust and mantle. Underneath the oceans, the Moho is roughly 5 km (3 mi) below the ocean floor. Under the continents, it is about 30-40 km (18.5-25 miles) below the surface. Near large mountain-building events, or orogenies, the continental Moho depth can be more than double that depth.

Figure \(\PageIndex{4}\): Depiction of seismic waves emanating from an earthquake (red star). Some waves travel through the crust to the seismic station (at ~6 km/s), while others go down into the mantle (where they travel at ~8 km/s) and are bent upward toward the surface, reaching the station before the ones that traveled only through the crust. "Seismic Wave Propagation in the Crust an Mantle" by Steven Earle is licensed under CC BY 4.0.

The Core

While there is no direct evidence about the composition of the core, the relative prevalence of chemical elements in our Solar System among other observations suggest the core to be made primarily of iron (Fe) and nickle (Ni) alloy.

Box \(\PageIndex{2}\): Discovering Earth's Internal Layers

Ancient cultures have long imagined caves and volcanoes provided openings into a mysterious and marvelous underworld filled with chasms, rivers, and fantastic beasts. In more recent times, writers like Jules Verne tapped into this natural understanding with his adventure novel, Journey to the Center of the Earth. But even before Verne’s time, observations of Earth’s gravitational and magnetic fields began to reveal an Earth with a very different interior. The scientific effort to understand Earth’s interior is an interesting history of using observations made at Earth’s surface to infer the properties of its interior.

In 1687, Isaac Newton outlined a procedure for determining the density of Earth. Density, or mass per unit volume of a material, is what determines the gravitational potential of an object, such as a planet. Astronomers were interested in determining the density of Earth to improve their understanding of the Solar System and planetary motion. It wasn’t until 1738 that Pierre Bouguer and Charles-Marie de La Condamine attempted to use Newton’s procedure at Mount Chimborazo in Ecuador to calculate the density of Earth and were able to prove that Earth is not hollow. This experiment was refined and repeated in 1774 by Charles Hutton, who was able to calculate that Earth’s interior must consist of material with a higher density than the average density of Earth’s surface rock. Then, in 1798, Henry Cavendish performed the first experiment specifically designed to calculate the average density of Earth. He found it to be 5.48 g/cm³, which closely approximates the current established value of 5.51 g/cm³. So by 1798, the description of Earth as a planet had changed from that of a hollow Earth to that of a solid Earth filled with material more dense than the density of surface rock.

Stepping away from gravity measurements to magnetic measurements is the next step towards understanding Earth as a planet. Earth’s magnetic field was first measured by Charles Friedrich Gauss in 1832. He discovered that the magnetism of surface rocks was insufficient to account for Earth’s strong magnetic field and therefore he inferred that the interior of Earth must contain more iron than Earth’s surface rock. Iron is very dense compared to most rock, so more iron inside the Earth could also help account for the discrepancy between the density of surface rock and the higher density needed inside Earth to produce Earth’s gravitational field.

Geologists, planetary astronomers, and physicists therefore postulated that Earth must have an internal core containing significant amounts of iron. Without more information, there was however no technique that would allow them to calculate the distribution of iron within Earth, other than it must be near the center. All of this changed with the invention of the modern seismometer in 1880 by John Milne. This is because seismic waves can be used to infer physical properties of materials, like rocks, and they can also infer the location of changes in those physical properties within rock (see Earthquakes for a description of seismic waves).

In order for seismic waves to infer the physical properties of rock, knowledge about how temperature, pressure, and composition affect the physical properties of rock is also needed. Changes in temperature and pressure can cause rock recrystallization or a partial melting of the rock, either of which changes the physical properties of the rock (see Igneous Rocks and Metamorphic Rocks). In general, as pressure increases, rock becomes stiffer, and seismic wave velocities increase. But, as temperature increases, rock becomes weaker, and seismic wave velocities decrease. Therefore, the interplay of composition, pressure, and temperature determines the seismic wave velocity at any location within Earth. Melting or partial melting decreases P-wave velocities, while partial melting attenuates the strength of S-waves; if the material melts, S-waves are stopped. With the invention of the modern seismometer however seismic waves and changes in seismic waves could be recorded for analysis.

Continuous recordings of earthquake waves from the 1897 Assam, India earthquake allowed Richard Dixon Oldham in 1906 to identify the arrival of different types of seismic waves. Using a network of seismometers in Europe, and analyzing the waves recorded on seismograms from different locations, Oldham was able to show that Earth must have a core based upon how P-waves and S-waves had moved through Earth. He was also the first to try and estimate the size of Earth’s core. This supported the work of Emil Wiechert, who in 1897 tried to synthesize the information known about Earth from gravity, magnetic, and seismic data and had postulated that Earth must have an iron core, surrounded by a rocky shell or mantle (Mantle is an old English word for covering). Therefore by 1906, scientists had determined that not only was the Earth solid rather than hollow, but that it also has a layered internal structure.

Other geoscientists were also using seismic wave interpretation techniques and in 1909 Andrija Mohorovičić (pronounced mo-ho-ro-vee-cheech; audio pronunciation) studying seismograms from a local earthquake in Croatia identified a boundary between what is now known as the crust and the mantle, which was named the Mohorovičić Discontinuity in his honor. This important boundary is commonly referred to as the Moho. Discontinuity is a term used for a location where the physical properties of a material change across a short distance. A discontinuity may either be a distinct surface or a transition zone. The Moho is defined by an increase in P-wave velocity from approximately 6 km/s to 8 km/s (Figure \(\PageIndex{4}\)), which occurs as P-waves move through crustal rock into mantle rock of a different composition.

By 1914, Beno Gutenberg was able to use changes in seismic wave velocities to map the interior of Earth as three layers, an outermost crust, a mantle, and a core (Figure \(\PageIndex{1}\)). The core-mantle boundary, abbreviated CMB, is also called the Gutenberg Discontinuity. At the CMB, P-wave velocity decreases and P-waves are refracted, whereas S-waves are stopped. This creates shadow zones where direct arrivals of P-waves and S-waves are not recorded (Box Figure \(\PageIndex{2}\)).

Box Figure \(\PageIndex{2}\): Patterns of seismic wave travel through Earth’s mantle and core. S-waves can’t travel through the liquid outer core, so they leave a shadow on Earth’s far side. P-waves do travel through the core, but P-wave refraction bends seismic waves away from P-wave shadow zones. "Seismic Shadow Zones with Explanation" by Karla Panchuk (2021), a derivative of "Seismic Shadow Zones", is licensed under CC BY 4.0.

Because S-waves are stopped at the CMB, the core of Earth was inferred to be liquid. The importance of finding a liquid core was that it could also explain the Earth’s magnetic field, not because of the total amount of iron in the core, but because of the motion of the iron in the core. Moving liquid iron in the core creates the geodynamo that generates Earth’s magnetic field, as was proposed in the late 1940s by Walter Elsasser and Edward Bullard.

This concept of a simple liquid core changed in 1936 when, Inge Lehmann, analyzing P-wave data recorded in Europe from the 1929 Murchison earthquake in New Zealand, found a solid inner core inside the liquid outer core, as shown by the different shadow zones for P-waves and S-waves (Box Figure \(\PageIndex{2}\)). The boundary between the solid inner core and the liquid outer core is called the Lehmann Discontinuity.

Seismologists continue to map the interior of Earth based on variations of seismic wave velocities, and this information became part of both the framework of the development of the theory of plate tectonics and a more complete model for the structure within Earth. Figure \(\PageIndex{1}\) therefore illustrates how a system of layers based on the behavior of materials, or rocks, was delineated by the work of numerous seismologists and geophysicists since the 1930s.

In the 1970s, with both the advent of digital seismic recording systems and supercomputing power becoming common, the technique of seismic tomography has allowed geodynamicists to create 3-D images of the relatively subtle lateral variations in seismic velocity within the mantle, using methods like those used in medical CT scans. This has further delineated velocity differences within the mantle which lead to the development and refinement of an alternative system to describe Earth based upon physical properties and the layers – lithosphere, asthenosphere, and lower mantle.

Layers Defined by Physical Properties

The Earth can also be divided into five physical layers based on how each layer responds to stress. While there is some overlap in the chemical and physical designations of layers, specifically the core-mantle boundary, there are significant differences between the two systems. (Figure \(\PageIndex{1}\))

Lithosphere

Lithos is Greek for “stone”, and the lithosphere is the outermost physical layer of Earth; in fact, the “plates” of plate tectonics are actually lithospheric plates.

The lithosphere consists of all of the crust and the uppermost part of Earth's mantle. This layer is defined by the brittle behavior of rock under stress. The lithosphere is grouped into two types: oceanic and continental. The oceanic lithosphere is thinner and relatively rigid. It ranges in thickness from nearly zero at mid-ocean ridges, to as much as 120 km (75 mi) away from them. The continental lithosphere is generally thicker ranging from 40 to 280 km (25 to 174 mi).

The lithosphere is not continuous. It is broken into segments called tectonic plates. A plate boundary is where two plates meet and move relative to each other. Plate boundaries are where we see plate tectonics in action—mountain building, ocean basin formation, earthquakes, and volcanic activity.

Asthenosphere

The asthenosphere is the layer below the lithosphere. Astheno- means lacking strength, and the most distinctive property of the asthenosphere is movement. Unlike the rigid lithosphere, this layer, which is made of rock that in places can be partially melted, moves and flows. Unlike the lithosphere which consists of multiple plates, the asthenosphere is relatively unbroken. Seismologists determined this by analyzing seismic waves that pass through this layer. The depth to the asthenosphere is temperature dependent and generally extends to depths no more than 200 km (124 mi). The asthenosphere is closer to Earth’s surface in areas where lateral forces are pulling Earth’s surface apart, such as mid-ocean ridges and continental rifts. It is much deeper underneath mountains and the centers of lithospheric plates.

Lower Mantle (Mesosphere)

The lower mantle, while still ductile, is less mobile than the asthenosphere. The top of the lower mantle is located at a depth of approximately 660 km (410 mi) below Earth’s surface. The lower mantle is subjected to very high pressures and temperatures, which is why it is considerably less mobile than the asthenosphere.

Box \(\PageIndex{3}\): Mantle Minerals

The extreme conditions within the mantle create a transition zone in the uppermost part of the lower mantle where minerals continuously change into various forms or pseudomorphs. Scientists identify this zone by changes in seismic velocity (the speed of seismic waves) and occasional physical barriers to movement. Below this upper transitional zone, the lower mantle is relatively uniform until it reaches the core.

Perovskite silicates (e.g. Bridgmanite, (Mg, Fe)SiO₃) are thought to be the main component of the lower mantle, making them the most common minerals in or on Earth.

The following video describes the abundant mineral perovskite:

Inner and Outer Core

The deepest two layers within Earth are the outer core and inner core. While they have the same composition, they behave in distinctly different ways. The outer core is the only entirely liquid layer within Earth and the inner core is a rigid solid layer. The outer core starts at a depth of 2,900 km (1,800 mi) and extends to 5,100 km (3,200 mi), making it ~2,200 km (1,400 mi) thick. The solid inner core’s radius is ~1,200 km (750 mi). The combined radius of the inner and outer cores is 3,400 km (2,100 mi), which makes it larger than Earth’s moon.

Earth’s liquid outer core is critically important in maintaining a breathable atmosphere and other environmental conditions favorable for life at or near Earth’s surface. The motion of the liquid outer core is what creates Earth’s strong magnetic field. If circulation in the outer core were to stop, Earth’s magnetic field would weaken and Earth’s atmosphere would be destroyed by the solar wind. Ions from the sun’s corona, the solar wind, can energize and strip away gas molecules from the atmosphere. In fact, this was the premise of the 2003 fictional film “The Core”. Should Earth’s liquid core stop circulating, its surface would cool and more solar radiation would reach the planet’s surface making conditions for life more difficult. This is what happened, and continues to happen, on Mars.

It may seem like a contradiction that the deepest and hottest part of Earth is solid because at such high temperatures metals such as iron and nickel should be molten. However this area within Earth not only has the highest temperature, it also has the highest pressure. The high pressure forces these metals to remain solid in spite of the high temperatures. From the center of Earth outward, both the pressure and the temperature decrease. Eventually you reach a pressure that is no longer high enough to force the metals in the core into the solid state, while the temperature is still more than enough for them to melt. Therefore, they melt and you transition from a solid inner core to a liquid outer core.

Video: Layers of the Earth - What are they? How were they found? (Educational)

The following video provides a nice summary of this section on Earth's interior.

References

1. Danson, Edwin (2006). Weighing the World. Oxford University Press.
2. Geomagnetism Frequently Asked Questions. (2022, March 9). National Centers for Environmental Information (NCEI). https://www.ncei.noaa.gov/products/geomagnetism-frequently-asked-questions
3. Historical Geology – A free online textbook for Historical Geology courses. (2020, January 17). https://opengeology.org/historicalgeology/
4. Pasyanos, M. E. (15 May 2008). "Lithospheric Thickness Modeled from Long Period Surface Wave Dispersion" (PDF). Retrieved 2014-04-25.
5. Stacey, F. D. (1977). Physics of the Earth, 2nd ed., John Wiley and Sons.
6. Turcotte, D. L. & Schubert, G. (2014). Geodynamics, 3rd ed., Cambridge University Press.