12.1 Earth's internal structure and composition

Earth's interior is a layered system with distinct physical and chemical properties at each depth. Understanding this structure is foundational for geodynamics: it explains why plates move, how Earth generates its magnetic field, and what drives volcanic activity and seismicity.

Geophysicists probe the interior using seismic waves, gravity measurements, magnetic surveys, and heat flow analysis. These methods reveal a planet that is chemically stratified (crust, mantle, core) and mechanically stratified (lithosphere, asthenosphere, mesosphere), with pressure and temperature jointly controlling the behavior of materials at every depth.

Earth's Internal Structure

Major Layers and Their Characteristics

Earth's interior divides into three chemically distinct layers: the crust, mantle, and core.

• Crust: The outermost layer, 5–70 km thick, composed of relatively low-density silicate rocks (granite in continental crust, basalt in oceanic crust).
• Mantle: Extends from the base of the crust to ~2,900 km depth. Composed of hot, dense ultramafic rock, predominantly peridotite.
• Core: Extends from ~2,900 km to Earth's center at ~6,371 km. Composed of iron-nickel alloy. The outer core is liquid, while the inner core is solid. The inner core remains solid despite higher temperatures because the immense pressure (over 300 GPa) raises the melting point of iron above the local temperature.

Lithosphere and Asthenosphere

The chemical layering above doesn't fully capture how the Earth behaves mechanically. For that, you need the rheological layering.

• The lithosphere includes the crust and the uppermost mantle. It behaves as a rigid, brittle shell, typically 70–150 km thick (thinner beneath ocean ridges, thicker beneath old continental interiors).
• The asthenosphere lies beneath the lithosphere, extending to roughly 200–300 km depth. It's solid rock, but at temperatures close to its melting point, so it deforms in a ductile, viscous manner over geological timescales.
• The boundary between them is defined by a change in rheology (mechanical behavior), not composition. This transition is what makes plate tectonics possible: rigid lithospheric plates can slide over the weaker asthenosphere.

Methods for Studying Earth's Interior

Seismic Waves

Seismic waves are the primary tool for imaging Earth's interior. They're generated by earthquakes or controlled-source explosions and change speed, direction, and amplitude as they pass through materials with different elastic properties and densities.

Two key body-wave types:

• P-waves (primary/compressional): Travel through both solids and liquids. They arrive first because they're faster (~6 km/s in the upper mantle, ~13 km/s near the base of the mantle).
• S-waves (secondary/shear): Travel only through solids. The fact that S-waves do not propagate through the outer core is the direct evidence that it's liquid.

How boundaries are detected:

• Seismic waves reflect and refract at interfaces where material properties change abruptly. The Mohorovičić discontinuity (Moho) marks the crust-mantle boundary, identified by a sharp increase in P-wave velocity. The core-mantle boundary (CMB) at ~2,900 km produces strong reflections and the S-wave shadow zone.
• Seismic tomography inverts travel-time data from many earthquakes recorded at many stations to build 3D velocity models of the interior. Regions with anomalously fast velocities typically correspond to colder, denser material (e.g., subducting slabs), while slow anomalies suggest hotter, less dense material (e.g., mantle plumes).

Other Geophysical Methods

• Gravity measurements (from satellites like GRACE/GOCE and surface gravimeters) map lateral density variations. Positive gravity anomalies indicate denser subsurface material; negative anomalies suggest lower-density regions.
• Magnetic surveys detect variations in Earth's magnetic field. At the surface, these reflect the distribution of magnetic minerals in crustal rocks. At a global scale, the field's geometry and secular variation constrain models of fluid flow in the electrically conductive outer core.
• Heat flow measurements quantify the rate of heat loss through Earth's surface (global average ~87 mW/m²). These data constrain the thermal structure of the lithosphere and the vigor of mantle convection beneath it.

Composition and Properties of Earth's Layers

Crust

The crust is composed primarily of silicate minerals, but oceanic and continental crust differ significantly.

• Oceanic crust: 5–10 km thick, mafic in composition (rich in Mg and Fe), dominated by basalt and gabbro. It forms at mid-ocean ridges by partial melting of the underlying mantle. Average density is ~3.0 g/cm³.
• Continental crust: 30–70 km thick, felsic in composition (rich in Si and Al), with an average composition close to granodiorite. It forms through a complex history of magmatic accretion, metamorphism, and differentiation over billions of years. Average density is ~2.7 g/cm³.

This density contrast is why continental crust "floats" higher on the mantle (isostasy) and is not easily subducted.

Mantle

The mantle makes up ~84% of Earth's volume. Its dominant rock type is peridotite, rich in olivine and pyroxene.

• The upper mantle (down to ~410 km) is relatively cooler and more rigid. Xenoliths brought up by volcanic eruptions provide direct samples of this region.
• The transition zone (410–660 km) is defined by pressure-induced mineral phase changes: olivine transforms to wadsleyite at ~410 km and then to ringwoodite at ~520 km. These transitions produce sharp increases in seismic velocity and density.
• The lower mantle (660–2,900 km) is hotter and composed of high-pressure phases, primarily bridgmanite (Mg-silicate perovskite structure) and ferropericlase. Near the CMB, bridgmanite may further transform to post-perovskite, which influences the dynamics of the lowermost mantle (D" layer).

Mantle convection, driven by heat from the core and internal radioactive decay (primarily U, Th, K), is the engine of plate tectonics. Convective flow transports heat from the interior to the surface and drives the motion of lithospheric plates. Partial melting of the mantle at ridges and above subduction zones generates the magmas that feed volcanism and build new crust.

Core

The core is composed primarily of iron-nickel alloy, with a small fraction of lighter elements (likely S, O, Si, or some combination). Evidence for its composition comes from seismic velocities, Earth's bulk density (~5.5 g/cm³ vs. crustal rocks at ~2.7–3.0 g/cm³), and comparison with iron meteorites.

• Outer core (~2,900–5,150 km): Liquid. Its density ranges from ~9.9 to ~12.2 g/cm³. Convective motions of this electrically conductive fluid generate Earth's magnetic field through a self-sustaining geodynamo.
• Inner core (~5,150–6,371 km): Solid, with a density of ~13 g/cm³. It's slowly growing as the Earth cools and the outer core crystallizes at the inner core boundary (ICB). This crystallization releases latent heat and light elements into the outer core, helping to drive the convection that sustains the geodynamo.

Pressure and Temperature's Influence on Earth's Interior

Pressure Effects

Pressure increases with depth due to the weight of overlying material. At the center of the Earth, pressure reaches approximately 360 GPa (about 3.6 million atmospheres).

• Increasing pressure raises the density and seismic velocity of materials, even without a change in composition.
• Pressure drives mineral phase transitions that define key boundaries in the mantle. For example, olivine ($\alpha$-phase) transforms to wadsleyite ($\beta$-phase) near 410 km depth, and ringwoodite ($\gamma$-phase) transforms to bridgmanite + ferropericlase near 660 km. Each transition involves a denser crystal structure accommodating the same bulk chemistry.
• In the lower mantle, high-pressure phases like bridgmanite and post-perovskite have distinct elastic properties that influence seismic wave propagation and mantle dynamics.

Temperature Effects

Temperature increases with depth, but the geothermal gradient is not constant.

• In the lithosphere, the gradient is steep (~25–30 °C/km near the surface).
• In the convecting mantle, the gradient is much gentler (close to the adiabatic gradient, ~0.3–0.5 °C/km), because convection efficiently homogenizes temperature.
• Sharp thermal boundary layers exist at the base of the lithosphere and at the CMB, where temperature jumps by an estimated 1,000–1,500 K over a relatively thin zone.

High temperatures reduce viscosity and promote ductile flow, enabling mantle convection. They also control where partial melting occurs: when the local temperature exceeds the solidus of mantle rock (which itself depends on pressure and composition), melt is generated.

Combined Pressure and Temperature Effects

Pressure and temperature act together to determine the phase, rheology, and density of materials at any given depth.

• The inner core is solid not because it's cold, but because pressure raises the melting point of iron above the local temperature. The outer core is liquid because pressure is lower there, and the temperature exceeds the melting point.
• Mantle convection patterns depend on the pressure- and temperature-dependent viscosity of mantle rock. Viscosity variations of several orders of magnitude exist between the upper and lower mantle.
• Unique mineral assemblages stable only under extreme P-T conditions (e.g., post-perovskite near the CMB) influence seismic anisotropy, heat transport, and the dynamics of thermal boundary layers.

Understanding how pressure and temperature jointly control material behavior is essential for modeling mantle convection, the geodynamo, and the long-term thermal evolution of the planet.