Earth's outer core

The Earth's outer core is a fluid layer of molten metal that lies beneath the mantle and surrounds the solid inner core, extending from a depth of approximately 2,890 kilometers to 5,150 kilometers below the surface, with a thickness of about 2,260 kilometers.^[1] It is primarily composed of iron and nickel, alloyed with lighter elements such as sulfur, oxygen, or silicon, which account for roughly 10% of its mass and contribute to its lower density compared to pure iron-nickel mixtures.^[2] The outer core's liquid state results from extreme temperatures ranging from about 4,000°C to 5,700°C and immense pressures, allowing it to flow and convect due to heat escaping from the inner core and residual heat from Earth's formation.^[3] Its density varies from approximately 9.9 g/cm³ at the core-mantle boundary to 12.2 g/cm³ near the inner core boundary, reflecting compression and compositional gradients.^[4] This convective motion in the electrically conductive outer core, driven by Earth's rotation and thermal gradients, generates the planet's geomagnetic field through a process known as the geodynamo.^[5] The magnetic field protects Earth's surface from harmful solar radiation and cosmic rays, influencing atmospheric retention and enabling life as we know it.^[6] Seismic studies, including analysis of P- and S-wave propagation, confirm the outer core's liquidity, as shear (S) waves do not transmit through it, while compressional (P) waves are refracted and slowed.^[7] The core-mantle boundary marks a sharp transition in both chemical composition and physical properties, with the overlying mantle being silicate-rich and solid.^[8] Ongoing research into the outer core's dynamics reveals interactions with the mantle that may influence volcanic activity and plate tectonics at the surface, while laboratory experiments and geophysical models continue to refine estimates of its exact composition and heat flux.^[9] The outer core constitutes about 15% of Earth's volume but over 30% of its mass, underscoring its fundamental role in planetary structure and evolution.^[2]

Discovery and Evidence

Historical Context

The recognition of Earth's outer core as a distinct liquid layer emerged from pioneering seismic investigations in the early 20th century. In 1906, British seismologist Richard Dixon Oldham analyzed earthquake records and identified a shadow zone for S-waves—secondary shear waves that fail to propagate beyond approximately 103° angular distance from the epicenter to the antipode (180°)—suggesting a central region incapable of transmitting shear motion, which he interpreted as evidence for a liquid core beneath the solid mantle.^[10] This marked the first clear indication of the core-mantle boundary (CMB) and highlighted the role of seismic waves in probing Earth's interior structure.^[11] Building on Oldham's work, German seismologist Beno Gutenberg refined these observations in 1913 by compiling extensive travel-time data from global earthquakes and developing mathematical models of wave propagation. He precisely located the CMB at about 2,900 km depth, confirming the existence of a dense central core separated from the mantle by a sharp discontinuity in seismic velocities.^[12] Gutenberg's analysis also reinforced the inference of liquidity, as the absence of direct S-waves through the core aligned with a fluid medium.^[13] A pivotal advancement came in 1936 when Danish seismologist Inge Lehmann discovered the inner core while studying reflected P-waves (primary compressional waves) in earthquake data. By identifying a distinct boundary around 5,100 km depth where P-wave velocities increased abruptly, Lehmann proposed a solid inner core enveloped by the liquid outer core, providing indirect evidence for the outer region's fluidity through contrasting wave behaviors: S-waves absent in the outer core but potentially present in a solid inner one.^[14] This discovery solidified the two-part core model and underscored the outer core's role in differentiating seismic signatures.^[15] Post-1950s progress in earthquake seismology, driven by expanded global networks and improved instrumentation, further established the outer core's fluid properties through detailed body-wave analyses. The deployment of the World Wide Standardized Seismograph Network (WWSSN) in the early 1960s enabled high-resolution recording of thousands of earthquakes, revealing consistent P-wave velocity reductions in the core (around 8 km/s versus 13 km/s in the mantle) and the complete lack of S-wave transmission, unequivocally confirming liquidity without shear rigidity. These observations refined early models and integrated with emerging computational techniques for travel-time inversions.^[16] From the 1960s onward, laboratory experiments complemented seismological evidence by simulating core conditions. Shock-wave compression studies on iron, such as those conducted by McQueen, Marsh, and colleagues, measured the equation of state up to two megabars (200 GPa), demonstrating that liquid iron under high pressure and temperature matches the observed seismic densities and velocities of the outer core. These high-pressure simulations, using explosive-driven techniques, provided critical validation for iron-dominated liquid models and influenced subsequent geophysical interpretations.^[17]

Seismic Indicators

The absence of S-waves (shear waves) through the outer core provides key evidence for its liquid state, as S-waves cannot propagate through fluids. This results in an S-wave shadow zone where no direct S-waves are detected beyond an angular distance of approximately 103° from the earthquake epicenter, extending to the antipode.^[18] The shadow zone arises because S-waves are absorbed or refracted at the core-mantle boundary (CMB), confirming the outer core's fluidity and distinguishing it from the solid mantle.^[19] P-waves (compressional waves), which can travel through both solids and liquids, exhibit a sharp velocity reduction at the CMB, dropping from about 13.7 km/s in the lowermost mantle to approximately 8 km/s in the outer core.^[20] This discontinuity causes P-waves to refract downward into the core, creating a P-wave shadow zone between 103° and 142° angular distance, where direct arrivals are absent due to the slower velocity and bending paths.^[19] At the inner core boundary (ICB), P-wave velocities increase abruptly to around 11 km/s as waves enter the solid inner core, observable in travel time curves that show refraction patterns and a velocity jump.^[21] Seismic waves traversing the liquid outer core experience significant attenuation, primarily from scattering off small-scale heterogeneities and viscous dissipation in the fluid.^[22] This damping reduces wave amplitudes and generates coda waves, providing insights into the core's low viscosity and turbulent flow.^[23] Modern observations leverage global seismograph networks like the Incorporated Research Institutions for Seismology (IRIS) Global Seismographic Network, which records core-penetrating phases such as PKP waves.^[24] Array processing techniques, including beamforming and interferometry, enhance signal detection by suppressing noise and precisely mapping boundary locations through slowness analysis of wave arrivals.^[25] These methods, applied to data from temporary and permanent arrays, have refined models of core boundaries with sub-degree resolution.^[26]

Location and Boundaries

Depth and Extent

The Earth's outer core is situated between the core-mantle boundary (CMB) at a depth of 2,891 km below the surface and the inner core boundary (ICB) at 5,150 km depth, according to the Preliminary Reference Earth Model (PREM).[]^[27] This positioning places the outer core within the planet's deeper interior, encompassing a thickness of approximately 2,260 km.[]^[27] For context, Earth's mean radius measures 6,371 km, rendering the outer core's volume about 16% of the total planetary volume.[] In comparison, the solid inner core has a radius of roughly 1,220 km, highlighting the outer core's dominance in the central region's scale.[]^[27] These boundary depths represent global averages derived from seismic data, but lateral heterogeneities introduce variations, such as CMB topography with amplitudes up to ±20 km.[]^[28]

Interfaces with Adjacent Layers

The core-mantle boundary (CMB), located at a depth of approximately 2,890 km, represents a profound physical and chemical discontinuity between the silicate-rich lower mantle and the metallic outer core.^[29] This interface exhibits a sharp density increase from about 5.6 g/cm³ in the lowermost mantle to 9.9 g/cm³ in the outer core, resulting in a total jump of roughly 4.3 g/cm³.^[30] The primary contrast arises from the chemical separation of oxide-dominated mantle silicates and iron-nickel alloys in the core, though seismic evidence points to additional heterogeneities, such as iron-enriched silicate layers or partial melts within the D″ layer of the lowermost mantle (extending 200–300 km above the CMB).^[31] These features may form through the accumulation of dense, iron-rich MgSiO₃ melts at CMB pressures (around 135 GPa), creating localized zones with density contrasts of up to 1.6% relative to surrounding bridgmanite solids.^[31] Seismic studies reveal enhanced reflectivity and azimuthal anisotropy at the CMB, often linked to ultra-low velocity zones (ULVZs)—patchy structures 5–50 km thick where P- and S-wave velocities drop by 10–30% compared to the bulk lower mantle.^[32] ULVZs, detected globally but concentrated near large low-shear-velocity provinces, likely stem from partial silicate melting (melt fractions of 10–30%) or compositional anomalies enriched in iron or bridgmanite, contributing to the boundary's irregular topography (up to ±15 km).^[32] Such seismic signatures indicate a complex, potentially stratified interface that modulates heat transfer, with turbulent outer core flows generating pressure variations that deform the CMB and couple with overlying mantle dynamics.^[33] At the inner core boundary (ICB), approximately 5,150 km from the surface, the transition from liquid outer core to solid inner core occurs via a gradual solidification front characterized by a mushy zone several kilometers thick.^[34] This zone facilitates dendritic growth of iron crystals as outer core fluid freezes onto the inner core, driven by latent heat release and compositional buoyancy from light elements like sulfur or oxygen rejected during solidification.^[34] The temperature at the ICB is estimated at roughly 5,000–6,000 K, corresponding to the melting point of iron-nickel alloys at inner core boundary pressures of approximately 330–360 GPa.^[35][36] Seismic reflectivity at the ICB shows weak precursors and anisotropy aligned with inner core fabrics, reflecting the textural heterogeneity of the mushy layer and a density jump of 0.8 ± 0.2 g/cm³.^[37] Dynamically, turbulence in the outer core boundary layers at both the CMB and ICB influences global circulation, with CMB interactions promoting localized upwellings in mantle convection through variable heat flux (3–15 TW) and topographic torques that transfer angular momentum to the mantle.^[33] At the ICB, the mushy zone regulates inner core growth rates (0.3–1 mm/year) and geomagnetic field stability by buffering compositional convection in the outer core.^[34]

Physical Properties

Composition Overview

The Earth's outer core is primarily composed of a liquid alloy dominated by iron and nickel, forming the bulk of its metallic structure. Iron constitutes approximately 85-90% by weight, while nickel accounts for about 5-10%, consistent with analyses of iron meteorites and geophysical models of core formation.^[38][39] This Fe-Ni alloy provides the high-density foundation observed in seismic data, with the liquid state enabling fluid dynamics essential to planetary processes. To reconcile the outer core's density with that of pure iron under equivalent pressure and temperature conditions, a total of 5-15% light elements by weight is required, creating an estimated density deficit of 5-10%.^[38][2] These lighter components alloy with the iron-nickel matrix, lowering overall density without significantly altering the core's metallic character. The exact nature of these light elements remains under investigation, but their presence is crucial for matching seismological observations. In addition to the major constituents, the outer core incorporates trace amounts of siderophile elements, such as cobalt (around 0.2-0.3% by weight) and platinum-group metals (at parts-per-billion to parts-per-million levels), which partitioned into the metallic phase during Earth's differentiation.^[40] These elements, metal-loving by affinity, were sequestered from the silicate mantle into the growing core, reflecting the planet's accretionary history. The composition is generally assumed to be homogeneous throughout the outer core due to vigorous convective mixing, which promotes thorough blending of the alloy over geological timescales.^[40] However, subtle radial variations may exist near the boundaries with the mantle and inner core, influenced by pressure gradients and thermal stratification.

Density, Temperature, and Pressure

The density of the outer core increases with depth due to compression, ranging from approximately 9.9 g/cm³ at the core-mantle boundary (CMB) to 12.2 g/cm³ at the inner core boundary (ICB), as established by the Preliminary Reference Earth Model (PREM).^[27] This profile reflects the material's response to increasing pressure, with the observed densities exhibiting a 5-10% deficit relative to pure iron-nickel alloys under equivalent conditions, attributable to the presence of light elements. Temperatures in the outer core span 4,000-6,000 K, following an adiabatic gradient that maintains near-isentropic conditions conducive to convection. At the CMB, temperatures are estimated at around 4,000 K, rising to approximately 5,700 K at the ICB.^[41] Pressures within the outer core escalate from 136 GPa at the CMB to 330 GPa at the ICB, driving the compression that shapes the density profile. The equation of state for liquid iron in the outer core is commonly described using the Birch-Murnaghan formulation, which relates density to pressure through finite strain theory. This parameterization aligns seismic observations with laboratory-derived parameters for iron alloys under core conditions.^[42] The viscosity of the outer core material is low, on the order of $10^{-2}$10−2 Pa·s, which facilitates rapid fluid motion essential for dynamo processes.^[43]

Light Elements in Composition

Candidate Elements and Proportions

The Earth's outer core consists primarily of iron and nickel, with the remaining non-Fe-Ni fraction comprising light elements that account for a density deficit relative to pure Fe-Ni alloys. These light elements constitute a combined mass fraction of 2-16 wt% to reconcile seismic observations with cosmochemical models of planetary composition.^[38] Among the leading candidate light elements, sulfur (S) is estimated at up to 6 wt%, based on geochemical partitioning and seismic velocity constraints that limit higher abundances to avoid excessive sound speed reductions. Oxygen (O) features in models up to 8 wt%, particularly in scenarios requiring compatibility with high-pressure metal-silicate equilibria. Silicon (Si) ranges from 2-7 wt%, supported by experimental partitioning data indicating its solubility in liquid iron under core conditions. Hydrogen (H), though present in smaller amounts of 0.5-1 wt%, is increasingly favored as a contributor due to its efficiency in matching density and velocity profiles without over-depressing seismic velocities. Carbon (C) is also a candidate, with models suggesting up to 4-6 wt% in Fe-C alloys to explain inner core nucleation and density deficits.^{[39][44][45][46][47]} Partitioning experiments under inner core boundary (ICB) conditions demonstrate that sulfur and oxygen preferentially partition into the liquid outer core, depressing the melting point of iron alloys by 300-1000 K compared to pure iron, which influences the phase relations and compositional gradient at the ICB. For instance, additions of 2-12 wt% S or O lower the liquidus temperature, promoting a stable liquid outer core while enabling inner core solidification.^[39][48] Post-2020 models increasingly support hybrid compositions incorporating multiple light elements, such as S-O-Si combinations at 2-6 wt% total to satisfy geophysical constraints, or H-Si mixtures totaling around 5 wt% (e.g., 4.1 wt% Si and 0.47 wt% H) derived from first-principles simulations and seismic data as of 2024. Fe-C models from 2025 suggest 4 wt% C for homogeneous inner core nucleation. Emerging neutrino oscillation tomography studies suggest potential for detecting lower overall light element contents, such as 1 wt% H in FeNiH alloys, by probing electron density profiles through the core. Recent thermodynamic models as of 2025 indicate light elements may be stratified, increasing rapidly in the top 200–400 km of the outer core under equilibrium conditions.^{[46][49][50][47]}

Geophysical and Geochemical Constraints

Seismic data impose fundamental constraints on the composition of Earth's outer core by linking observed P- and S-wave velocities to density profiles. The Adams-Williamson equation, derived under hydrostatic equilibrium and assuming chemical homogeneity, describes the radial density variation as $\frac{d\rho}{dr} = -\frac{\rho g}{\phi}$drdρ​=−ϕρg​, where $\rho$ρ is density, $g$g is gravitational acceleration, and $\phi = K_S / \rho$ϕ=KS​/ρ is the adiabatic incompressibility, with $K_S$KS​ the adiabatic bulk modulus.^[51] For the liquid outer core, this equation integrates Poisson's ratio near 0.5, reflecting near-incompressibility typical of fluids, which allows inversion of seismic velocities to infer density and thus compositional deficits.^[52] The Preliminary Reference Earth Model (PREM), constructed via least-squares inversion of travel-time data while obeying the Adams-Williamson equation in the outer core, yields a density about 8-10% lower than pure liquid iron at core pressures, requiring light elements to explain the deficit.^[53] Geochemical analyses further limit outer core composition through mantle depletions in siderophile elements relative to chondritic meteorites, which represent primitive solar system material. The mantle's nickel content is depleted by a factor of ~20 compared to chondrites, indicating preferential partitioning of Ni into the core during metal-silicate differentiation under reducing conditions.^[54] Similar depletions in cobalt and other siderophiles, such as iridium and platinum, quantified via mass balance between the bulk silicate Earth and core, imply that the outer core sequesters ~85-90% of Earth's siderophile budget, consistent with equilibrium partitioning models at high pressures and temperatures.^[55] Cosmochemical modeling of CI chondrites reinforces these constraints, showing that core formation must account for the observed mantle abundances to maintain bulk Earth chondritic ratios for refractory elements.^[56] High-pressure laboratory experiments using diamond anvil cells provide direct tests of how light elements affect iron alloy properties at outer core conditions. Melting curve determinations for Fe-S alloys reveal eutectic behavior, where the liquidus temperature depresses by 1,000-2,000 K relative to pure iron across 50-200 GPa, due to sulfur's role in lowering the melting point and stabilizing the liquid phase.^[57] These experiments, combining laser heating with in situ X-ray diffraction, demonstrate that such depressions persist to inner core boundary pressures, constraining the thermal profile and potential light element inventories needed to sustain outer core liquidity.^[58] Electromagnetic constraints arise from the outer core's role in generating the geomagnetic field, requiring high electrical conductivity consistent with metallic iron alloys. Inversions of geomagnetic secular variation and dynamo simulations, aligned with PREM's density and velocity structure, estimate outer core conductivity at approximately $10^5$105 S/m, which demands a predominantly metallic composition with light elements that preserve electron mobility without insulating phases.^[59] This value, derived from time-dependent magnetic field models, rules out non-metallic enrichments and supports alloys where light elements like sulfur or silicon enhance rather than disrupt conductive bonding.^[39] Advancements in the 2020s, including ab initio molecular dynamics simulations, have tightened bounds on light element content by computing equation-of-state properties for Fe-light element mixtures at core conditions. These calculations, incorporating density functional theory, indicate a total light element fraction of less than 10 wt% to match PREM's seismic profile while accounting for thermodynamic stability.^[59] Complementarily, neutrino geochemistry uses oscillations of geoneutrinos from U and Th decay to probe core-mantle partitioning, yielding upper limits on light elements that align with <10% total abundance and favor oxygen-poor compositions to avoid excessive neutrino scattering.^[60] Such methods provide model-independent verification, emphasizing the outer core's Fe-dominated nature with minor admixtures like S, O, Si, H, or C.^[49]

Formation and Evolutionary Implications

Role in Planetary Accretion

The formation of Earth's outer core is intrinsically linked to the planet's accretion from planetesimals and protoplanets during the early Solar System, approximately 4.5 billion years ago. In models such as the Grand Tack hypothesis, Jupiter's inward then outward migration disrupted the protoplanetary disk, facilitating the accretion of terrestrial planets through a series of giant impacts that delivered metallic cores from impactors. These collisions enabled metal-silicate equilibration at high pressures (tens of gigapascals), partitioning iron-nickel alloys into the growing core while leaving silicates in the mantle. The Nice model complements this by describing subsequent instabilities among the outer giant planets, which scattered planetesimals inward and influenced the late stages of inner planet growth, including Earth's core assembly.^[61] Isotopic signatures provide key evidence for the rapidity of core formation during accretion. Tungsten-182 (¹⁸²W) anomalies in the mantle, arising from the short-lived ¹⁸²Hf-¹⁸²W decay system (half-life ~8.9 million years), indicate that core-mantle differentiation occurred within about 30 million years of the Solar System's formation, as tungsten was preferentially sequestered into the metallic core while hafnium remained in the silicate mantle. Positive ¹⁸²W excesses (up to +15 ppm) preserved in Archean rocks and sediments reflect incomplete early mixing and limited convection, consistent with a core that formed via rapid bombardment and sinking of metal during the planet's embryonic stages. These anomalies underscore how accretionary impacts minimized the time for isotopic equilibration, shaping the outer core's initial composition.^[62] Light elements in the outer core serve as tracers of volatile delivery during late accretion phases. Sulfur (S), for instance, likely originated from carbonaceous chondrite-like materials, with CI chondrites contributing 10-15% of Earth's building blocks before core formation and CM-like bodies adding ~0.5% post-core, delivering volatiles such as S, selenium (Se), and tellurium (Te) that exhibit a characteristic "hockey stick" depletion pattern relative to refractory elements. This pattern matches the bulk Earth abundances (e.g., S at 5300-7900 µg/g), with partitioning into the core yielding higher concentrations (e.g., S at 16000-24000 µg/g), reflecting incomplete siderophile depletion during metal-silicate separation under accretionary conditions. Such delivery highlights the outer core's role in retaining volatiles that evaded earlier high-temperature loss. The energy budget of accretion profoundly influenced the outer core's liquidity and differentiation. Giant impacts converted gravitational potential energy and kinetic energy into internal heat, with total budgets reaching ~10³¹ J per event, elevating mantle entropy by 6-10 kJ/K/kg and causing widespread melting that facilitated metal percolation and core growth. Gravitational energy release during metal sinking further contributed comparable heat, driving the separation of the liquid outer core from any inner solidification while preventing full crystallization. This incomplete process, sustained by impact heating, ensured the outer core's molten state, essential for its dynamical properties.^[63]

Core Differentiation Processes

The formation of Earth's outer core involved the separation of denser metallic materials from lighter silicates during the planet's early differentiation, a process driven by metal-silicate partitioning. In this mechanism, iron-nickel alloys, being significantly denser, sank toward the center while leaving behind a silicate-rich mantle, facilitated by gravitational instabilities such as Rayleigh-Taylor instability that promoted the descent of metallic blobs through the proto-mantle.^[64][65] This partitioning occurred rapidly following Earth's accretion around 4.5 billion years ago, establishing the core-mantle boundary.^[66] At the inner core boundary (ICB), partial solidification of the outer core proceeds via crystal fractionation, where iron crystals slowly grow outward from the solid inner core into the overlying liquid. This growth occurs at a rate of approximately 0.5–1 mm per year, moderated by the release of latent heat during crystallization, which inhibits rapid freezing and maintains the outer core in a predominantly liquid state.^[67][68] The presence of light elements, such as sulfur, oxygen, or silicon, further suppresses full solidification by depressing the freezing point of the iron alloy by 500–1,000 K relative to pure iron, ensuring compositional buoyancy that sustains the liquid layer.^[69] Thermal evolution models of Earth's core indicate cooling rates of 100–200 K per gigayear, with the outer core acting as a thermal buffer that regulates the pace of inner core expansion. These models account for heat loss primarily through conduction and convection, balancing the energy budget to prevent excessive solidification while allowing gradual growth of the inner core over billions of years.^[70][71]

Dynamical Role

Convection and Heat Flow

The convection within Earth's outer core is characterized by vigorous fluid motions driven by buoyancy forces from both thermal and compositional gradients, resulting in advection-dominated heat transfer that efficiently transports heat from the core interior to the core-mantle boundary (CMB). This process delivers an estimated heat flux of approximately 10 TW across the CMB to the overlying mantle, with the primary mechanism being compositional convection fueled by the ongoing growth of the solid inner core, which releases lighter elements into the surrounding liquid.^[72] Such advection ensures that conductive heat transfer plays a negligible role, as the turbulent flows maintain a well-mixed state throughout much of the outer core volume. A key feature of this convection is its double-diffusive nature, where a stabilizing temperature gradient (due to adiabatic heating during ascent) coexists with a destabilizing compositional gradient from the enrichment of light elements near the inner core boundary. This interplay promotes oscillatory instabilities and layered flows, operating at Rayleigh numbers exceeding $10^{20}$1020, which signifies highly turbulent, chaotic motion far exceeding the critical threshold for onset of convection.^[73][74] The resulting dynamics enhance vertical mixing while preserving some radial stratification, contributing to the overall efficiency of heat and momentum transport in the rotating system. At the CMB, thin thermal boundary layers form, typically 100-200 km thick, where conductive heat loss to the cooler mantle concentrates temperature gradients and influences the pattern of convective upwellings. These layers act as a thermal buffer, potentially seeding large-scale mantle plumes through localized hot spots, and their thickness modulates the global heat flux by controlling the rate of boundary layer instability and turnover.^[72] The energy powering these convective processes stems from a combination of sources that maintain the necessary buoyancy contrasts: primarily secular cooling of the core (~60% of the total power through gradual thermal contraction) and latent heat released during inner core solidification (~40% via the exothermic phase change and associated compositional buoyancy), with uncertain minor contributions from radiogenic heating in the core (0-10%).^[72] Advancements in modeling during the 2020s have leveraged geomagnetic secular variation data for tomographic imaging of outer core flows, enabling reconstructions of three-dimensional velocity patterns that reveal westward drift and localized accelerations at rates of several kilometers per year near the CMB.^[75] These models highlight the turbulent, geostrophic nature of the flows and their sensitivity to boundary conditions, providing insights into the core's dynamical evolution over decadal timescales.

Generation of the Geomagnetic Field

The geomagnetic field is generated by the geodynamo process operating within Earth's liquid outer core, where convective motions of the electrically conductive molten iron alloy induce electric currents that, in turn, amplify and sustain the magnetic field through electromagnetic induction. This self-sustaining dynamo mechanism relies on the advection of magnetic field lines by swirling fluid flows outweighing ohmic diffusion, characterized by a magnetic Reynolds number $Rm \approx 10^3 - 10^4$Rm≈103−104.^[76] The turbulent convection patterns in the outer core provide the necessary organized motion to stretch, twist, and fold field lines, enabling the regeneration of the poloidal field from the toroidal field and vice versa via the $\alpha$α- and $\omega$ω-effects in mean-field dynamo theory.^[77] At Earth's surface, the geomagnetic field exhibits a predominantly dipolar structure, with the axial dipole component accounting for approximately 80-90% of the total field intensity, which averages around 50 $\mu$μT. Within the core, the field comprises intertwined poloidal (source-free, curl-free outside the core) and toroidal (azimuthally directed) components, generated and maintained by the differential rotation and helical convection of the fluid.^[78] This configuration periodically undergoes full reversals of polarity, occurring on average every 200,000 to 300,000 years over the past 20 million years, as recorded in paleomagnetic data from oceanic and continental rocks. The reversals involve a temporary weakening and multipolar instability before reestablishing a dominant dipole, without interrupting the overall dynamo operation. Recent paleomagnetic studies as of 2024 indicate the geodynamo initiated around 3.7 billion years ago, approximately 0.8 billion years after planetary formation, with evidence for possible earlier onset at 4.2 Ga.^[79] Sustaining the geodynamo demands significant power, estimated at 1-3 TW, primarily to counteract ohmic dissipation—the resistive heating from induced currents in the conductive fluid—which represents the main energy sink. This power is supplied by the buoyancy-driven convection in the outer core, ensuring the field's long-term stability over billions of years. Numerical simulations have been pivotal in validating this process; for instance, the Glatzmaier-Roberts model, a three-dimensional, time-dependent magnetohydrodynamic simulation incorporating compositional and thermal convection, reproduces key observed features such as dipole dominance, westward drift, and secular variation in field intensity and morphology over timescales of tens of thousands of years.^[80] These models also demonstrate spontaneous field reversals, aligning with paleomagnetic records and highlighting the dynamo's sensitivity to core parameters like rotation rate and boundary conditions. The incorporation of light elements, such as oxygen, silicon, or sulfur, into the outer core alloy reduces its density while influencing electrical conductivity, which is crucial for dynamo efficiency. Higher conductivity from these elements lowers the critical Rm threshold for dynamo onset and enhances field amplification, facilitating the geodynamo's initiation in the early Earth roughly 1-2 Ga after planetary formation, as inferred from paleointensity data from ancient rocks. This compositional effect underscores how light elements not only aid convection through buoyancy contrasts but also optimize the electromagnetic coupling necessary for sustaining the field against diffusive losses.