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Internal structure of Earth
Internal structure of Earth
Internal structure of Earth
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Geological cross section of Earth, showing the different layers of the interior.

The internal structure of Earth is the layers of the planet Earth, excluding its atmosphere and hydrosphere. The structure consists of an outer silicate solid crust, a highly viscous asthenosphere, and solid mantle, a liquid outer core whose flow generates the Earth's magnetic field, and a solid inner core.

Scientific understanding of the internal structure of Earth is based on observations of topography and bathymetry, observations of rock in outcrop, samples brought to the surface from greater depths by volcanoes or volcanic activity, analysis of the seismic waves that pass through Earth, measurements of the gravitational and magnetic fields of Earth, and experiments with crystalline solids at pressures and temperatures characteristic of Earth's deep interior.

Global properties

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Chemical composition of the upper internal structure of Earth[1]
Chemical element/oxide Chondrite model (1) (%) Chondrite model (2) (%)
MgO 26.3 38.1
Al2O3 2.7 3.9
SiO2 29.8 43.2
CaO 2.6 3.9
FeO 6.4 9.3
Other oxides N/A 5.5
Fe 25.8 N/A
Ni 1.7 N/A
Si 3.5 N/A

Note: In chondrite model (1), the light element in the core is assumed to be Si. Chondrite model (2) is a model of chemical composition of the mantle corresponding to the model of core shown in chondrite model (1).[1]

see caption
A photograph of Earth taken by the crew of Apollo 17 in 1972. A processed version became widely known as The Blue Marble.[2][3]

Measurements of the force exerted by Earth's gravity can be used to calculate its mass. Astronomers can also calculate Earth's mass by observing the motion of orbiting satellites. Earth's average density can be determined through gravimetric experiments, which have historically involved pendulums. The mass of Earth is about 6×1024 kg.[4] The average density of Earth is 5.515 g/cm3.[5]

Layers

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Schematic view of Earth's interior structure.
  1.   continental crust
  2.   oceanic crust
  3.   upper mantle
  4.   lower mantle
  5.   outer core
  6.   inner core

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.[6] The geologic component layers of Earth are at increasing depths below the surface.[6]: 146 

Crust and lithosphere

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Main articles: Earth's crust and Lithosphere
Map of Earth's tectonic plates
Earth's major plates, which are:

Earth's crust ranges from 5 to 70 kilometres (3.1–43.5 mi)[7] in depth and is the outermost layer.[8] The thin parts are the oceanic crust, which underlies the ocean basins (5–10 km) and is mafic-rich[9] (dense iron-magnesium silicate mineral or igneous rock).[10] The thicker crust is the continental crust, which is less dense[11] and is felsic-rich (igneous rocks rich in elements that form feldspar and quartz).[12] The rocks of the crust fall into two major categories – sial (aluminium silicate) and sima (magnesium silicate).[13] It is estimated that sima starts about 11 km below the Conrad discontinuity,[14] though the discontinuity is not distinct and can be absent in some continental regions.[15]

Earth's lithosphere consists of the crust and the uppermost mantle.[16] The crust-mantle boundary occurs as two physically different phenomena. The Mohorovičić discontinuity is a distinct change of seismic wave velocity. This is caused by a change in the rock's density[17] – immediately above the Moho, the velocities of primary seismic waves (P wave) are consistent with those through basalt (6.7–7.2 km/s), and below they are similar to those through peridotite or dunite (7.6–8.6 km/s).[18] 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.[clarification needed]

Many rocks making up Earth's crust formed less than 100 million years ago; however, the oldest known mineral grains are about 4.4 billion years old, indicating that Earth has had a solid crust for at least 4.4 billion years.[19]

Mantle

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Main article: Earth's mantle

[20]

Earth's crust and mantle, Mohorovičić discontinuity between bottom of crust and solid uppermost mantle

Earth's mantle extends to a depth of 2,890 km (1,800 mi), making it the planet's thickest layer.[21] [This is 45% of the 6,371 km (3,959 mi) radius, and 83.7% of the volume - 0.6% of the volume is the crust]. The mantle is divided into upper and lower mantle[22] separated by a transition zone.[23] The lowest part of the mantle next to the core-mantle boundary is known as the D″ (D-double-prime) layer.[24] The pressure at the bottom of the mantle is ≈140 GPa (1.4 Matm).[25] The mantle is composed of silicate rocks richer in iron and magnesium than the overlying crust.[26] Although solid, the mantle's extremely hot silicate material can flow over very long timescales.[27] Convection of the mantle propels the motion of the tectonic plates in the crust. The source of heat that drives this motion is the decay of radioactive isotopes in Earth's crust and mantle combined with the initial heat from the planet's formation[28] (from the potential energy released by collapsing a large amount of matter into a gravity well, and the kinetic energy of accreted matter).

Due to increasing pressure deeper in the mantle, the lower part flows less easily, though chemical changes within the mantle may also be important. The viscosity of the mantle ranges between 1021 and 1024 pascal-second, increasing with depth.[29] In comparison, the viscosity of water at 300 K (27 °C; 80 °F) is 0.89 millipascal-second [30] and pitch is (2.3 ± 0.5) × 108 pascal-second.[31]

Core[32]

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Main articles: Earth's inner core and Earth's outer core
A diagram of Earth's geodynamo and magnetic field, which could have been driven in Earth's early history by the crystallization of magnesium oxide, silicon dioxide, and iron(II) oxide. Convection of Earth's outer core is displayed alongside magnetic field lines.
A diagram of Earth's geodynamo and magnetic field, which could have been driven in Earth's early history by the crystallization of magnesium oxide, silicon dioxide, and iron(II) oxide

Earth's outer core is a fluid layer about 2,260 km (1,400 mi) in height (i.e. distance from the highest point to the lowest point at the edge of the inner core) [36% of the Earth's radius, 15.6% of the volume] and composed of mostly iron and nickel that lies above Earth's solid inner core and below its mantle.[33] Its outer boundary lies 2,890 km (1,800 mi) beneath Earth's surface. The transition between the inner core and outer core is located approximately 5,150 km (3,200 mi) beneath Earth's surface. Earth's inner core is the innermost geologic layer of the planet Earth. It is primarily a solid ball with a radius of about 1,220 km (760 mi), which is about 19% of Earth's radius [0.7% of volume] or 70% of the Moon's radius.[34][35]

The inner core was discovered in 1936 by Inge Lehmann and is composed primarily of iron and some nickel. Since this layer is able to transmit shear waves (transverse seismic waves), it must be solid. Experimental evidence has at times been inconsistent with current crystal models of the core.[36] 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.[37][38] The laser studies create plasma,[39] 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.[40][41]

Under laboratory conditions a sample of iron–nickel alloy was subjected to the core-like pressure 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.[42][43]

The composition of Earth bears strong similarities to that of certain chondrite meteorites, and even to some elements in the outer portion of the Sun.[44][45] Beginning as early as 1940, scientists, including Francis Birch, built geophysics upon the premise that Earth is like ordinary chondrites, the most common type of meteorite observed impacting Earth. This ignores the less abundant enstatite chondrites, which formed under extremely limited available oxygen, leading to certain normally oxyphile elements existing either partially or wholly in the alloy portion that corresponds to the core of Earth.[citation needed]

Dynamo theory suggests that convection in the outer core, combined with the Coriolis effect, gives rise to Earth's magnetic field. The solid inner core is too hot to hold a permanent magnetic field (see Curie temperature) but probably acts to stabilize the magnetic field generated by the liquid outer core. The average magnetic field in Earth's outer core is estimated to measure 2.5 milliteslas (25 gauss), 50 times stronger than the magnetic field at the surface.[46]

The magnetic field generated by core flow is essential to protect life from interplanetary radiation and prevent the atmosphere from dissipating in the solar wind. The rate of cooling by conduction and convection is uncertain,[47] but one estimate is that the core would not be expected to freeze up for approximately 91 billion years, which is well after the Sun is expected to expand, sterilize the surface of the planet, and then burn out.[48][better source needed]

Seismology

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Main article: Seismology

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.

See also

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References

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Further reading

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Revisions and contributorsEdit on WikipediaRead on Wikipedia

Internal structure of Earth

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The internal structure of Earth is characterized by four primary concentric layers—the crust, mantle, outer core, and inner core—differentiated by their compositions, densities, temperatures, and physical states, which together form a dynamic influencing geological processes like and the planet's . The crust, the outermost and thinnest layer, is a rigid shell of solid rock composed mainly of , varying in thickness from about 5 kilometers beneath ocean basins to 30–70 kilometers under continental landmasses. Beneath the crust lies the mantle, a thick layer extending approximately 2,900 kilometers deep, consisting of dense, semi-solid silicate rocks rich in iron, magnesium, and calcium that flow slowly over geological timescales due to driven by internal heat. The mantle is divided into the rigid upper portion, known as the (which includes the crust and uppermost mantle, about 80 kilometers thick), and the more ductile below it, enabling the movement of tectonic plates. The outer core, a layer of molten iron and alloy roughly 2,200–2,300 kilometers thick, surrounds the mantle's base and generates through its convective motion. At the center is the inner core, a solid of iron and about 1,220–1,250 kilometers in radius, enduring extreme pressures and temperatures up to 5,400°C while remaining solid. This layered model was primarily deduced in the early through the analysis of seismic waves from earthquakes, which refract and reflect differently at boundaries between layers, combined with data on Earth's , , and magnetic properties. These internal features not only maintain Earth's by protecting against solar radiation but also drive surface phenomena such as , earthquakes, and mountain building.

Physical Characteristics

Size, Mass, and Density

Earth's equatorial radius measures 6,378 km, while its polar radius is 6,357 km, resulting in an oblate spheroid shape due to rotational flattening. The mean radius, calculated as the volumetric average, is 6,371 km. The total mass of Earth is 5.972×10245.972 \times 10^{24} kg. This value is derived from Newton's law of universal gravitation, which relates surface gravitational acceleration gg to the planet's mass MM, mean radius RR, and the gravitational constant GG via the equation g=GMR2g = \frac{G M}{R^2}, rearranged to solve for M=gR2GM = \frac{g R^2}{G}. Here, g≈9.81g \approx 9.81 m/s² (standard surface value), R=6.371×106R = 6.371 \times 10^6 m, and G=6.67430×10−11G = 6.67430 \times 10^{-11} m³ kg⁻¹ s⁻². Earth's average density is 5.51 g/cm³, significantly higher than the typical density of surface rocks, which ranges from 2.7 g/cm³ for to 3.0 g/cm³ for . This contrast indicates a denser interior, with heavier materials concentrated toward the center. The (PREM), developed by Dziewonski and Anderson in 1981, provides a standard radial profile based on seismic data, normal mode periods, and constraints, showing increasing from about 2.7 g/cm³ at the surface to over 13 g/cm³ at the core-mantle boundary. The for , defined as CMR2=0.3307\frac{C}{M R^2} = 0.3307 where CC is the polar , MM is the total , and RR is the mean , reveals a non-uniform distribution with denser material toward . This value, lower than 0.4 for a uniform , confirms the planet's internal differentiation.

Shape and Gravitational Field

The deviates from a perfect , adopting the of an oblate spheroid primarily due to its , which generates centrifugal forces that counteract more strongly at the than at the poles. This rotational effect causes an , with the equatorial approximately 21 kilometers greater than the polar , resulting in a factor of about 1/298.25. The balance between these centrifugal forces and shapes the planet's overall figure, influencing the distribution of and the internal stress field, though and mantle remain largely hydrostatic under this equilibrium. The global exhibits variations, or anomalies, that arise from uneven mass distributions within and on the , detectable through measurements such as those from the Gravity Recovery and Climate Experiment (GRACE) mission and its follow-on GRACE-FO. Launched in 2002 and operating until 2017, GRACE's twin satellites tracked changes in Earth's by measuring microscopic variations in their separation due to mass redistributions; GRACE-FO, launched in 2018, continues these observations, revealing anomalies on scales from regional to global, including those linked to ice melt, ocean currents, and deeper mantle flows. These data highlight how rotational dynamics and surface processes perturb the field, providing indirect evidence of internal mass shifts without direct sampling. The represents the surface of Earth's gravity field that coincides with mean , serving as a reference for measurements worldwide. Undulations in the , relative to an idealized ellipsoidal shape, reach amplitudes of up to about 100 meters, reflecting underlying heterogeneities in and crust that alter local gravitational pull. These irregularities stem from variations in internal , such as those from previous discussions on bulk properties, and are quantified through satellite gravimetry. To account for oblateness in gravitational calculations, the acceleration g(r)g(r) is modified from the simple point-mass form g(r)=GMr2g(r) = \frac{GM}{r^2}, where GG is the , MM is Earth's mass, and rr is the radial distance. The primary adjustment incorporates the second-degree zonal harmonic coefficient J2≈1.0826×10−3J_2 \approx 1.0826 \times 10^{-3}, which captures the equatorial flattening's effect on the via the expansion V(r,ϕ)=GMr[1−J2(Rer)23sin⁡2ϕ−12]V(r, \phi) = \frac{GM}{r} \left[ 1 - J_2 \left( \frac{R_e}{r} \right)^2 \frac{3 \sin^2 \phi - 1}{2} \right], where ReR_e is the equatorial radius and ϕ\phi is the ; differentiation yields the adjusted gg. This coefficient, derived from orbit analyses, is essential for precise modeling of the field's latitudinal variations.

Structural Layers

Crust and Lithosphere

The Earth's crust represents the outermost layer of the planet, forming a thin, brittle shell that varies significantly in thickness and composition between oceanic and continental regions. Oceanic crust is typically 5 to 10 kilometers thick and primarily composed of basaltic rocks, which are mafic in nature and enriched in magnesium and iron. In contrast, continental crust averages 30 to 70 kilometers in thickness and consists mainly of granitic or andesitic rocks, which are felsic and rich in silica (SiO₂), with an average composition approximating andesite at about 60% SiO₂. These differences arise from the distinct geological processes that form each type, with oceanic crust generated at mid-ocean ridges and continental crust built through long-term accretion and differentiation. The boundary between the crust and the underlying mantle is marked by the , commonly known as the Moho, where seismic P-wave velocities exhibit a distinct jump from approximately 6.7 km/s in the crust to 8.1 km/s in the mantle. This discontinuity, first identified in , signifies a sharp change in rock density and composition, separating the silica-rich crustal materials from the more magnesium- and iron-rich ultramafic rocks of the . The Moho's depth aligns with crustal thickness variations, occurring at 5-10 km beneath oceans and 30-70 km under continents, providing a key marker for the planet's layered structure. Encompassing the crust and the uppermost portion of the mantle, the lithosphere forms the Earth's rigid outer shell, typically 100 to 200 kilometers thick, where rocks behave brittlely due to lower temperatures and pressures. This mechanical layer contrasts with the underlying by its strength and coherence, enabling it to support tectonic stresses without significant deformation. The is divided into several large plates that float on the semi-fluid asthenosphere below, carrying both oceanic and continental crust as they move. For instance, the Pacific Plate, one of the largest lithospheric plates, spans much of the basin and includes both oceanic crust and fragments of continental margins, influencing global seismic and volcanic activity through its interactions at plate boundaries.

Mantle

The mantle is the thickest layer of Earth's interior, extending from an average depth of approximately 30 km beneath the crust to about 2,900 km at the core-mantle boundary. It constitutes roughly 84% of Earth's volume and is primarily composed of rocks, with a increasing from around 3.3 g/cm³ in the to about 5.6 g/cm³ in the due to compression and phase changes. The mantle is divided into the , spanning from the base of the crust to approximately 660 km depth, and the , from 660 km to 2,900 km; this division is marked by significant seismic discontinuities associated with phase transitions. The bulk composition of the mantle is inferred to be ultramafic, dominated by , which consists primarily of (approximately 40-50%), (around 40%), and lesser amounts of and other minerals in the upper regions. This pyrolite model, proposed as a primitive mantle composition, assumes a chondritic of elements and is supported by analyses of mantle-derived xenoliths and mid-ocean ridge basalts. Debates persist regarding whether the mantle convects as a single whole-mantle system or as layered reservoirs separated by chemical or rheological barriers, with seismic evidence suggesting some degree of layering due to subducted accumulation, though whole-mantle mixing is favored in many dynamic models. Phase transitions play a critical role in the mantle's structure, driven by increasing pressure and temperature with depth. At around 410 km, the olivine-dominated upper mantle undergoes a transformation to wadsleyite (beta-spinel structure), followed by ringwoodite (gamma-spinel) deeper in the transition zone; these changes cause seismic velocity increases of 5-10% for P-waves. At approximately 660 km, ringwoodite dissociates into bridgmanite (perovskite) and ferropericlase (magnesiowüstite), marking the upper-lower mantle boundary with a velocity jump and density increase of about 8%. Near the core-mantle boundary, at depths of roughly 2,500-2,900 km, bridgmanite further transitions to post-perovskite, a layered silicate phase that may influence seismic anisotropy in the lowermost mantle. Additionally, two vast large low-shear-velocity provinces (LLSVPs) occupy much of the lowermost mantle, potentially remnants of primordial material. Recent research as of November 2025 connects these structures to Earth's early differentiation and the delivery of volatiles essential for life. Despite its solid state, the mantle exhibits ductile , deforming plastically over geological timescales through mechanisms like dislocation creep and diffusion creep, influenced by , , grain size, and water content. varies significantly with depth, estimated at about 10^{21} Pa·s in the due to and hydration weakening, increasing to around 10^{22}-10^{23} Pa·s in the deeper where higher pressures suppress deformation. The , a low-viscosity zone approximately 100-200 km thick located at depths of 100-400 km, corresponds to a seismic where P- and S-wave velocities decrease by 3-5% owing to elevated temperatures and minor melt fractions (1-2%), facilitating decoupling and motion of the overlying lithospheric plates.

Core

The Earth's core consists of two distinct regions: a liquid outer core and a solid inner core, both primarily composed of iron-nickel alloys with lighter elements. The outer core extends from approximately 2,900 km to 5,150 km depth, with a thickness of about 2,250 km, and is made of a molten iron-nickel alloy containing 85–90% iron, 5–10% , and lighter elements such as and oxygen. This fluid layer, with temperatures ranging from roughly 4,000 K near the core-mantle boundary to 5,700 K at its inner boundary, generates through convective action. The core-mantle boundary, known as the at ~2,900 km depth, marks a sharp increase in density and seismic velocity, separating the silicate mantle from the metallic core; associated ultra-low velocity zones (ULVZs) are patchy regions at this interface where seismic waves slow by 10–50%, likely due to or chemical heterogeneity. The inner core is a solid sphere with a radius of ~1,220 km, spanning depths from 5,150 km to the Earth's center at 6,371 km, also dominated by iron-nickel alloys but solidified under extreme pressure despite temperatures around 5,700–6,000 K at the inner core boundary. It exhibits seismic , with faster wave speeds along polar directions than equatorial ones, attributed to aligned iron crystals; this structure includes an innermost inner core layer (~300–650 km radius) with distinct crystal orientations and slower velocities compared to the outer inner core. The inner core grows slowly at ~0.5–1 mm per year as outer core material freezes onto its surface, releasing that influences core dynamics. A 2025 study suggests that carbon played a key role in enabling the inner core's formation by lowering the of iron alloys, allowing solidification under extreme conditions. Recent seismic studies have revealed structural complexities, such as a low-velocity toroidal (ring-shaped) region in the outer core at low latitudes, with ~2% slower seismic speeds, possibly linked to compositional variations. As of 2025, further analyses indicate the inner core's surface is changing shape over decades, suggesting it is less rigidly solid than previously assumed, with potential implications for geodynamo stability. Observations from repeating earthquakes indicate dynamic changes in the core over decades, including potential deformations of the inner core's and surface undulations at the core-mantle boundary driven by outer core , which exert topographic torques on the overlying mantle. These findings, based on waveform analyses from 1991–2023, show shifts in seismic signal , suggesting the inner core's and form have varied, with backtracking motions and hemispheric asymmetries. Such complexities highlight the core's evolving role in Earth's geodynamo and , contrasting its dense metallic nature with the overlying silicate mantle.

Methods of Investigation

Seismology

is the primary method for probing the Earth's internal structure, relying on the propagation of seismic waves generated by earthquakes to infer discontinuities and material properties at depth. These waves travel through the planet at varying speeds depending on the and elasticity of the rocks they encounter, allowing scientists to map radial variations in Earth's layers. By analyzing arrival times and amplitudes recorded at global seismograph networks, researchers identify boundaries where wave speeds abruptly change, revealing compositional and phase transitions. The two main types of body waves are primary (P) waves, which are compressional and propagate by alternating compression and dilation, and secondary (S) waves, which are shear waves that cause particle motion. P waves travel faster, with velocities ranging from 5 to 13 km/s, while S waves are slower at 3 to 7 km/s, enabling their distinction in recordings. At internal discontinuities, such as layer boundaries, these waves refract due to velocity contrasts or reflect back toward the surface, creating detectable signals that outline structural interfaces. For instance, the sudden increase in P-wave velocity from about 7.5 km/s in the to 8.1 km/s deeper marks a key transition. A hallmark of seismology's revelations came from observing shadow zones—regions where certain waves are absent or weakened on the opposite side of the from an . The S-wave shadow zone begins at approximately 103° angular distance and extends to 180°, as S waves cannot propagate through the liquid outer core, which lacks . The P-wave shadow zone, from 103° to about 143°, results from at the core-mantle boundary, where P-wave velocities drop sharply from 13.7 km/s in the to 8.1 km/s in the outer core, bending waves away from direct paths. These zones, first noted in early 20th-century data, provided conclusive evidence for a liquid outer core. Key discoveries of internal boundaries trace back to pioneering analyses of earthquake records. In 1909, Croatian seismologist Andrija Mohorovičić identified the crust-mantle boundary, now called the or Moho, at about 30-50 km depth beneath continents, where P-wave velocity jumps from 6-7 km/s to 8 km/s. German seismologist Beno Gutenberg determined the core-mantle boundary in 1913, placing it at roughly 2,900 km depth based on patterns and travel-time delays. Danish seismologist Inge Lehmann proposed the inner core boundary in 1936, interpreting reflected P waves (PKiKP phase) as evidence of a solid inner core starting at about 5,150 km depth, where velocities increase again. These findings established the basic layered model of Earth. Travel-time curves, plotting wave arrival times against epicentral distance, form the foundation for constructing one-dimensional (1D) velocity models of Earth's radial structure. These curves exhibit kinks at discontinuities, corresponding to refraction paths that graze boundaries. The Preliminary Reference Earth Model (PREM), developed in 1981 by Adam M. Dziewonski and Don L. Anderson, synthesizes global seismic data into a standard 1D profile, specifying P- and S-wave velocities, density, and attenuation as functions of depth. PREM, for example, models the upper mantle's velocity gradient and the core's low velocities, serving as a benchmark for interpreting new observations. Recent 2025 studies using seismic tomography have provided evidence of morphological changes in the inner core over decades and ancient island-like structures in the deep mantle. Advancements in computational seismology have enabled three-dimensional (3D) imaging through , which inverts vast datasets of wave delays to map lateral heterogeneities. This technique reveals variations in anomalies, such as faster regions in subducting slabs or slower ones in hotspots. Notably, two massive antipodal structures in the lowermost mantle, termed (LLSVPs), appear as broad zones (up to thousands of kilometers across) with shear-wave reductions of 1-3% beneath and the Pacific, potentially indicating compositionally distinct, dense material. Discovered in the 1980s via global , LLSVPs occupy about 8% of the mantle volume and influence deep mantle dynamics. These methods have confirmed boundaries like the Moho, core-mantle interface, and inner core surface, providing the framework for Earth's structural layers.

Other Geophysical Techniques

Satellite gravimetry missions, such as the Gravity Recovery and Climate Experiment (GRACE) and the Gravity Field and Steady-State Ocean Circulation Explorer (GOCE), have provided high-resolution maps of Earth's gravity field, revealing density variations that inform the internal structure. These missions detect subtle gravity anomalies associated with mantle plumes, which appear as positive or negative signals indicating low-density material, and subducting slabs, marked by linear negative anomalies tracing descending high-density into the mantle. For instance, GOCE data have imaged the as a broad low-gravity anomaly extending to the core-mantle boundary, supporting plume models, while GRACE observations delineate subduction zones like the with resolutions down to 100-200 km. Geomagnetic studies utilize paleomagnetic preserved in rocks to reconstruct the history of Earth's , which generates the through convective motions in the liquid outer . These , derived from the alignment of magnetic minerals during rock formation, document periodic field s, with evidence of at least 183 reversals in the last 83 million years, averaging every 300,000 years but with irregular intervals. The transitional fields during reversals show complex, non-dipole configurations, providing insights into dynamo stability and without direct access to the interior. Current observations indicate the field has weakened by about 9% since 1840, potentially signaling an impending , though the process typically spans thousands of years. Measurements of heat flow at Earth's surface, typically ranging from 40 to 100 mW/ with a global average of approximately 87 mW/, quantify the outward flux of thermal energy and constrain models of core cooling. Continental sites often record lower values around 60-80 mW/ due to thicker insulating crust, while oceanic measurements near mid-ocean ridges reach up to 100 mW/ from hydrothermal activity. This heat budget, totaling about 47 terawatts, arises primarily from radiogenic decay in and from core solidification, implying a core cooling rate of roughly 100-200 per billion years to sustain the geodynamo. High-pressure laboratory experiments, particularly using diamond anvil cells (DACs), replicate extreme conditions of the mantle and core to study material properties under pressures up to 400 GPa and temperatures exceeding 4000 K. In DACs, samples are compressed between opposing diamond tips, often heated by lasers, enabling observations of phase transitions, such as the perovskite-to-post-perovskite shift at the core-mantle boundary or the melting behavior of iron alloys in the core. These experiments confirm that the outer core is primarily liquid iron with light elements like or oxygen, reducing by 5-10% compared to pure iron, and reveal seismic wave speeds matching geophysical models. For the , DAC simulations demonstrate variations with depth, influencing patterns. Geoneutrino detection, as achieved by the Borexino experiment, probes Earth's composition by measuring antineutrinos from (²³⁸U) and (²³²Th) decay chains, which produce about 20-25 terawatts of radiogenic heat. Borexino, located underground in , has detected 53 geoneutrinos since 2007, with fluxes indicating mantle abundances of approximately 20 ppb and 80 ppb , consistent with chondritic models but suggesting a depleted core with minimal radiogenic contribution. These low-energy neutrinos (below 50 MeV) pass through unimpeded, providing direct evidence for heat sources driving internal dynamics without relying on seismic interpretations.

Formation and Dynamics

Planetary Differentiation

The Earth accreted from the solar nebula approximately 4.54 billion years ago, beginning as a heterogeneous body of dust, gas, and fragments that gradually coalesced into a proto- through gravitational instability and collisions. This process formed a hot, partially molten with a composition reflecting the inner solar system's materials, including silicates, metals, and volatiles. A pivotal event in early Earth history was the giant impact hypothesis, which posits that around 4.5 billion years ago, a Mars-sized protoplanet named Theia collided with the proto-Earth at an oblique angle, ejecting debris that coalesced to form the Moon while melting much of the Earth's surface into a global magma ocean. This impact not only stripped volatile elements and angular momentum but also homogenized the proto-Earth's interior, setting the stage for subsequent chemical separation. Planetary differentiation followed, driven by density contrasts in the molten proto-Earth: denser iron-nickel alloys sank toward the center via gravitational , while lighter materials rose to form the outer layers, a facilitated by Rayleigh-Taylor instabilities that caused denser metallic droplets to penetrate the surrounding matrix. Core formation completed within about 30 million years of solar system inception, as indicated by hafnium-tungsten isotopic ratios in mantle-derived rocks, while the magma ocean overlying solidified progressively over hundreds of millions to billions of years through fractional , beginning from the bottom up with minerals like accumulating at depth. Key evidence for this differentiation includes the depletion of siderophile (iron-loving) elements such as , , and platinum-group metals in the relative to their abundances in chondritic meteorites, which represent undifferentiated solar nebula remnants and suggest these elements were sequestered into during metal-silicate separation under high-pressure conditions. This pattern, observed in mantle xenoliths and basalts, aligns with equilibrium partitioning experiments and indicates late-stage additions of chondrite-like material ("late veneer") that slightly replenished mantle siderophiles after primary core formation.

Convection and Geodynamo

Mantle convection refers to the slow, heat-driven circulation of material within Earth's mantle, which plays a crucial role in driving geological processes at the surface. This convection is primarily powered by three sources of heat: secular cooling of the planet, radiogenic decay of isotopes such as uranium, thorium, and potassium, and latent heat released during inner core solidification. These mechanisms create temperature gradients that cause less dense, hotter material to rise and cooler, denser material to sink, forming large-scale circulation cells. Geologists debate between two main models for this process: whole-mantle convection, where material mixes throughout the entire mantle from the core-mantle boundary to the lithosphere, and two-layer (or layered) convection, which posits limited exchange between an upper and lower mantle separated by chemical or rheological barriers like the 660 km discontinuity. Evidence from seismic tomography and geochemical signatures supports aspects of both, but numerical simulations suggest whole-mantle convection can preserve long-term heterogeneities while allowing broad mixing. The surface manifestation of mantle convection is plate tectonics, where rigid lithospheric plates move over the underlying due to convective forces. Key driving mechanisms include slab pull, in which descending oceanic slabs at zones gravitationally pull the plate, and ridge push, where elevated mid-ocean ridges exert a gravitational force promoting plate motion away from spreading centers. These forces, rooted in , account for the majority of plate velocities, with slab pull often dominant at zones. thus links deep dynamics to surface features like volcanoes, earthquakes, and mountain building, with upwellings beneath ridges and downwellings at trenches forming the primary flow pattern. In the outer core, convection of liquid iron generates through the geodynamo process, where fluid motions organize into helical flows that amplify and sustain a geocentric axial field. This is driven by thermal buoyancy from core cooling and compositional differences arising from inner core growth, with the from organizing the flow into columnar structures aligned with the rotation axis. At the surface, the field strength varies between approximately 0.3 and 0.6 gauss, providing protection against . The field periodically reverses polarity, with an average interval of about 300,000 years, as evidenced by paleomagnetic records in oceanic basalts and continental sediments. Growth of the solid inner core, which has been ongoing for roughly 1 billion years, influences geodynamo stability by releasing and lighter elements into the outer core, enhancing compositional and providing flux to sustain the against ohmic dissipation. This growth stabilizes the dominance but can modulate field intensity, with models showing that a smaller inner core in the past led to weaker or more unstable fields. Recent studies from 2023 to 2025 highlight how turbulence in the outer core, driven by convective vigor, contributes to observed slowdowns in inner core rotation relative to , with seismic analyses confirming a deceleration since around 2010 as part of a ~70-year oscillation cycle. These insights, derived from modeling of data, suggest outer core flows may periodically couple more strongly with the inner core, altering its super-rotation and potentially affecting long-term behavior. As of 2025, seismic evidence also indicates that the inner core's outer boundary has undergone shape changes over recent decades, suggesting it is less solid and more dynamic than previously assumed, which may further influence outer core .

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