Introduction
The Earth’s inner core is a roughly spherical central region composed predominantly of a solid iron–nickel alloy with smaller amounts of lighter elements, extending to a radius of about 1,230 km. This radius corresponds to roughly one-fifth of Earth’s total radius and is on the order of 70% of the Moon’s radius, indicating a substantial central volume relative to planetary dimensions.
In the radial sequence from surface to center the principal structural layers are the continental and oceanic crusts, the upper and lower mantles, the liquid outer core and the solid inner core. Key seismic discontinuities that mark these transitions include the Mohorovičić discontinuity (crust–mantle), the core–mantle boundary and the outer-core/inner-core boundary.
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Direct samples of core material are not available, so inferences about composition, state and structure come mainly from the behavior of seismic waves and from geomagnetic observations. Thermal models place the temperature at the inner-core surface at roughly 5,700 K—comparable to the Sun’s photosphere—yet the material remains solid because the enormous pressure near Earth’s center elevates the melting point; this pressure‑dependence of melting is commonly expressed in thermodynamic formulations such as the Simon–Glatzel relation.
Inge Lehmann’s 1936 analysis of New Zealand earthquake seismograms revealed reflections from a distinct internal boundary within Earth’s core, from which she inferred an inner‑core radius of about 1,400 km—close to the modern estimate of ~1,221 km. Subsequent work by Beno Gutenberg and Charles Richter (1938) expanded the observational base and characterized the outer core as roughly 1,950 km thick, with a relatively sharp, approximately 300 km transition zone between outer and inner core, implying an inner‑core radius in the range ~1,230–1,530 km. By the 1940s geophysicists had proposed a solid iron composition for the inner core, and Francis Birch’s mineralogical analysis (1952) argued that the inner core is likely crystalline iron, linking seismic structure to plausible material phases. Terminology has varied: the inner–outer core boundary is sometimes called the “Lehmann discontinuity,” though that label more commonly refers to a different seismic interface; proposals to honor Keith E. Bullen with the compound name “Lehmann‑Bullen discontinuity” have not gained wide acceptance. Independent confirmation of the inner core’s solid behavior arrived in 1971 with measurements demonstrating mechanical rigidity distinct from the surrounding molten outer core. Further geophysical evidence came from studies of Earth’s normal modes—global oscillations excited by large earthquakes—by Adam Dziewonski and James Gilbert, which are consistent with a liquid outer core enclosing a solid inner core. Reports from 2005 of shear‑wave phases traversing the inner core, initially contested because S‑waves are normally excluded from fluids, have since gained broader acceptance and further reinforce the interpretation of a solid inner core within a liquid outer core.
Seismic waves
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Seismology supplies almost all direct constraints on the inner core because elastic waves from earthquakes traverse the deep interior and are recorded globally. Signals from earthquakes occurring well below the surface (commonly >30 km) are particularly valuable: their ray paths sample a relatively homogeneous mantle, producing simpler waveforms at distant stations that are easier to interpret for inner‑core studies.
Two basic wave types convey complementary information. P (compressional) waves are longitudinal and propagate through both solids and liquids; S (shear) waves are transverse and require a rigid solid to travel. P and S phases therefore have distinct phase velocities and attenuation characteristics even within the same material, which helps discriminate material properties along their paths. Seismic ray nomenclature encodes layer interactions: “P” denotes a compressional leg in the mantle/crust, “K” a passage through the liquid outer core, “I” transmission through the solid inner core, lowercase “i” reflection at the inner‑core boundary, and “J” a shear (S‑type) leg within the inner core produced by P→S conversion.
Specific arrivals are diagnostic of inner‑core structure and state. PKiKP phases are P waves that travel mantle→outer core, reflect at the inner‑core boundary, and return through the outer core and mantle; they are most straightforward to interpret when the receiver lies near the epicentral direction so the ray is nearly straight. PKIKP phases follow the same mantle–core–mantle route but transmit through the inner core itself, with the clearest recordings when the receiver is near the source antipode, yielding an approximately central straight‑line path. Although S waves cannot cross the liquid outer core directly, oblique P↔S mode conversions at the inner‑outer core boundary generate converted phases (e.g., PKJKP) in which a P converts to an S on entering the inner core, traverses the inner core as a shear wave, and reconverts to P on exit. The detection of such converted shear arrivals provides direct evidence that the inner core supports shear propagation and therefore behaves mechanically as a solid.
The solid inner core exerts a first-order control on planetary-scale geophysics through both dynamical and gravitational pathways. The geomagnetic field is generated by convective motion and induced electric currents in the liquid outer core; the presence of a solid inner core and the heat it supplies alter those convective and electromagnetic patterns, so inner-core structure and thermal state strongly influence geodynamo behavior. Although the inner core is composed mainly of iron, its temperature exceeds iron’s Curie point, so it does not contribute permanent magnetization; its influence on the magnetic field is therefore indirect and dynamic, mediated by flow modification and heat flux.
Independent gravitational and seismic observables provide complementary constraints on inner-core properties. Global measurements of Earth’s mass distribution, gravity anomalies and moment of inertia are sensitive to the density and geometry of deep internal layers, and thus limit inner-core density, size and radial structure. Large earthquakes excite whole-Earth normal modes whose frequencies and spatial patterns depend sensitively on the density, radial extent and shape of the inner layers; analysis of these free oscillations therefore yields independent, global information about inner-core structure. Robust inference about the inner core therefore integrates magnetic-field behavior (geodynamo dynamics and heat flow), gravimetric and inertia measurements (density and dimensions), and seismic normal-mode spectra, each responding in different but complementary ways to composition, physical state, geometry and thermal output.
Seismic velocities within Earth’s inner core exhibit a smooth radial diminution for both shear (S) and compressional (P) phases, and an abrupt change at the transition to the outer core. Shear‑wave speed decreases gradually from about 3.7 km s−1 at the geocenter to roughly 3.5 km s−1 at the inner‑core surface, while compressional‑wave speed falls modestly from ~11.4 km s−1 at the center to ~11.1 km s−1 near the inner‑core boundary. The P‑wave reduction with radius implies a systematic, though small, decline in compressional rigidity toward the inner‑core surface.
Across the inner‑core/outer‑core interface the character of seismic propagation changes sharply: P‑wave velocity drops abruptly from the inner‑core value near 11.1 km s−1 to about 10.4 km s−1 in the outer core, marking a clear discontinuity in elastic properties. Inner‑core S‑wave speeds are notably lower than typical values in the lower crust (~4.5 km s−1) and are much smaller than shear speeds observed in the deep mantle above the core (~7.3 km s−1), underscoring the distinct mechanical regime of the inner core relative to overlying layers.
Size and shape
The Earth’s inner core has a mean radius of about 1,221 km (diameter ≈ 2,442 km), which is roughly 19% of the planet’s mean radius and about 70% of the Moon’s radius. Its volume is ≈ 7.6 × 10^9 km^3 (7.6 × 10^18 m^3), representing approximately 1/146 (≈0.69%) of Earth’s total volume.
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Morphologically, the inner core is best represented as an oblate ellipsoid of revolution—that is, rotationally symmetric with a slight equatorial bulge. Its flattening f is small, estimated between 1/400 and 1/416, which corresponds to a polar radius only about 3 km shorter than the equatorial radius. By comparison, the whole Earth is more oblate (flattening ≈ 1/300), where the polar radius is about 21 km shorter than the equatorial radius; thus the inner core is measurably more nearly spherical than the planet as a whole.
At the inner core—the central solid sphere enclosed by the liquid outer core—internal pressure attains values on the order of 330–360 GPa (≈3.3–3.6 × 10^6 atm), modestly exceeding the pressure at the inner‑core boundary.
The gravitational acceleration at the inner‑core surface is calculated to be about 4.3 m·s^−2, less than half the standard surface gravity of 9.8 m·s^−2; this substantially reduced effective gravity reflects the interior gravitational field geometry, in which only the mass enclosed within a given radius contributes to the local gravitational acceleration and the field decreases toward the planet’s center.
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The solid inner core displays a smooth, monotonic decrease in density with radius, from about 13.0 kg·L−1 at the geometric center to approximately 12.8 kg·L−1 at the inner-core boundary (kg·L−1 ≡ g·cm−3 ≡ t·m−3). Superimposed on this gentle intrainner-core gradient is a sharp density discontinuity at the inner-core/outer-core boundary: the immediately overlying liquid outer core has a markedly lower density of ~12.1 kg·L−1.
This combination of a continuous radial gradation within the solid interior and an abrupt drop to the fluid outer core signals both compositional/structural variation with depth in the solid phase and a clear phase-related density contrast at the boundary. Compared with near-surface materials (the mean density of the outermost ~100 km is only ~3.4 kg·L−1), core materials are several times denser, reflecting their distinct composition and pressure–temperature state.
The inferred density distribution yields an inner-core mass on the order of 1×10^23 kg, roughly 1/60 (≈1.7%) of Earth’s total mass. These density characteristics—gradual intrainner-core variation plus a pronounced boundary contrast—are fundamental constraints on models of Earth’s internal structure, thermal evolution, and core dynamics.
Temperature
Estimates of the inner core temperature are derived by comparing the melting point of iron-bearing alloys at the pressure of the inner‑core boundary—about 330 GPa—with the actual thermal state inferred for that depth. First‑principles calculations and melting‑curve extrapolations (e.g., D. Alfè et al., 2002) placed the temperature near 5,400–5,700 K (≈5,100–5,400 °C). Later laboratory measurements (S. Anzellini et al., 2013) reported a substantially higher iron melting temperature at core pressures, about 6,230 ± 500 K (≈5,957 ± 500 °C).
Iron can remain solid at these very high temperatures because the melting point increases strongly with pressure: at ~330 GPa the melting curve is shifted to temperatures high enough to stabilize a solid inner core. This pressure dependence of the melting temperature is described quantitatively by the Clausius–Clapeyron relation, which relates the slope of the solid–liquid phase boundary to the enthalpy and volume changes on melting and thereby predicts how the melting temperature rises with increasing pressure.
Magnetic field
Bruce Buffett (2010) used observations of Earth’s nutation together with the tidal forcing of the electrically conducting outer core by the Moon and Sun to infer the characteristic magnetic field within the core. Tidal forces drive motions in the fluid outer core analogous to ocean tides; as this conducting fluid moves through the ambient magnetic field it induces electric currents, and the resistive loss of those currents (Ohmic or Joule dissipation) provides a damping torque on the tidal motion. By relating the magnitude of the nutation anomaly to the required electromagnetic damping, Buffett derived an average core field of roughly 2.5 milliteslas (≈25 gauss), about forty times larger than the strongest field measured at Earth’s surface. He further argued that the solid inner core is likely subject to a comparable internal field. Because the estimate depends primarily on dynamical electromagnetic coupling rather than on detailed models of thermal evolution or exact core composition, it constitutes a relatively model‑insensitive, indirect geophysical constraint on the amplitude of the core magnetic field.
Viscosity
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Seismic records indicate that waves traverse the inner core in a manner characteristic of a solid, but current seismological constraints do not definitively exclude a material that is fluid-like on very long timescales. This ambiguity in rheological interpretation permits models in which the inner core behaves as an effectively rigid body during short‑period observations while permitting imperceptibly slow deformation over geological time.
One class of models invokes extremely slow, long‑timescale convection within the inner core—conceptually analogous to mantle convection—to account for the directional dependence of seismic wave speeds (seismic anisotropy) observed in multiple studies. Such slow flow could reorient or recrystallize inner‑core material and thereby produce the anisotropic fabric inferred from seismic data.
Quantitative estimates illustrate this intermediate behaviour: Buffett (2009) estimated the inner core’s dynamic viscosity at about 10^18 Pa·s. To place that number in context, it is orders of magnitude larger than common viscous substances (roughly 10^24 times the viscosity of water when expressed in relative terms commonly cited in geophysics, and more than 10^9 times that of pitch), implying a medium that responds elastically on observational timescales but is capable of extremely slow, cumulative flow over millions to billions of years. Such a high but finite viscosity therefore reconciles seismic evidence for solidity with theoretical mechanisms for generating the observed anisotropy.
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Composition
Direct sampling of the inner core is not possible, so its chemical makeup is inferred from Solar System element abundances, models of planetary differentiation, and the bulk chemistry of the Earth constrained by seismic observations. These lines of evidence point to an iron-dominated alloy with significant nickel—reflecting the relative abundance of Fe and Ni and their siderophile behavior during core formation. Thermodynamic and seismic constraints, however, rule out pure iron: computed densities for pure Fe at inner‑core pressures and temperatures exceed the seismically determined density by roughly 3%, which requires the presence of lighter alloying elements. Silicon, oxygen and sulfur are leading candidates to account for this deficit; quantitative models from about 2007 permit on the order of 10% Ni plus a few percent (≈2–3%) of unidentified light components to match bulk density and seismic properties. First‑principles calculations (e.g., D. Alfè and collaborators) indicate the outer core is relatively oxygen‑rich (≈8–13 wt% O) and that oxygen is preferentially excluded from the solid phase during inner‑core crystallization, explaining a lower oxygen content in the solid inner core. Laboratory experiments and comparisons with seismic wave speeds further indicate the inner core is largely composed of ε‑iron (hexagonal close‑packed structure); this HCP lattice can incorporate modest amounts of nickel and trace light elements while remaining consistent with observed seismic velocities.
Early conceptions of the inner core treated it as compositionally and structurally uniform, reflecting the view that its growth proceeded in a steady, homogeneous manner and therefore produced consistent physical and chemical properties with depth. From this uniform-formation perspective some researchers advanced an extreme end-member hypothesis that the inner core might be a single, continuous crystal of iron rather than a mosaic of smaller crystallites. That monocrystalline model rests on the premise that solidification occurred without disruption, yielding an unbroken iron lattice throughout the inner-core volume.
Axis-aligned anisotropy
In 1983 Poupinet et al. reported a systematic travel‑time anomaly for PKIKP phases—compressional waves that cross the solid inner core—finding arrivals following near north–south trajectories through the inner core to be roughly 2 seconds earlier than those confined to equatorial paths. After accounting for non‑core geometrical effects (notably the Earth’s polar flattening, ≈0.33% for the whole Earth and ≈0.25% for the inner core) and for known crustal and upper‑mantle heterogeneities, the remaining differential requires an intrinsic anisotropy of compressional wave speed in the inner core. Quantitatively, this residual indicates P waves propagate about 1% faster along roughly polar directions than in equatorial directions.
This anisotropic signature is observed over a wide range of seismic wavelengths, implying it is a robust, scale‑persistent property of inner‑core structure rather than a narrow‑band artifact. Independent approaches—analyses of much larger seismic catalogs and examinations of whole‑Earth normal modes—have reinforced the travel‑time evidence and lent additional confidence to the existence of inner‑core anisotropy. Reported estimates of the anisotropy magnitude have varied across studies, with some upper‑bound values reaching as high as ~4.8% between polar and equatorial directions. A more recent reassessment by Frost and Romanowicz (2017) tightened these bounds, concluding that the inner‑core P‑wave speed difference lies in the range of approximately 0.5–1.5%, thereby narrowing the plausible magnitude of axis‑aligned anisotropy.
Non-axial anisotropy
Seismic observations of P-wave speeds in the Earth’s inner core reveal anisotropy—velocity depends on propagation direction. Some studies report regions where P waves travel fastest in directions that deviate substantially from the geographic north–south (rotation) axis, including roughly east–west orientations and intermediate oblique angles. These findings imply spatially heterogeneous anisotropic fabrics, with locally varying fast directions rather than a single, coherent alignment.
By contrast, Frost and Romanowicz argue that available seismic data are best interpreted as indicating a fast axis closely aligned with the planet’s rotation axis (approximately geographic N–S), within the limits of current resolution. The disagreement between these views therefore centers on whether anisotropy is dominantly polar-aligned at the global scale or includes significant local departures from that alignment.
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The contrasting interpretations reflect limitations of seismic sampling, methodological differences, and interpretative uncertainty: qualifiers such as “in at least some regions” and “as can be determined” underscore that data coverage and resolution constrain inferred orientations. The choice between a polar-aligned fast axis and regionally oblique/perpendicular fast directions carries distinct geodynamical implications. A rotation-axis-aligned fast direction favors coherent alignment of anisotropic material (e.g., crystal preferred orientation) parallel to the rotation axis, whereas locally oblique or perpendicular fast directions point to heterogeneous deformation, variable crystal alignment, or regionally distinct growth and flow regimes within the inner core.
Causes of anisotropy
Laboratory and theoretical studies indicate that hexagonal close-packed (ε) iron exhibits strong intrinsic anisotropy: individual crystals have a single “fast” axis for P-wave propagation and two orthogonal slower directions. If these fast axes are preferentially aligned on a regional scale, seismic velocities measured through the inner core will display systematic directional dependence. A concerted north–south alignment of fast axes, for example, readily explains the observed enhancement of P‑wave speeds along that meridional direction.
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Several physical processes have been proposed to produce such a lattice‑preferred orientation. Mechanical reorientation by slow viscous flow (creep) within the inner core will tend to rotate and align crystals in the direction of material transport. Yoshida et al. (1996) formulated a specific variant of this idea in which enhanced freezing at equatorial latitudes drives an equator‑to‑pole internal flow that both restores isostatic equilibrium of the inner‑core surface and aligns crystals along the poleward flow. An alternative but related thermal mechanism is slow convective motion: depending on the inner‑core rheology and its past size, thermal convection could have reoriented crystals. This possibility has been debated—Yukutake (1998) argued convective motions are unlikely, whereas Buffett (2009), using viscosity estimates and considerations of an initially smaller inner core, concluded convection could plausibly have occurred during some intervals.
Non‑deformation mechanisms have also been advanced. Bergman (1997) proposed that growth textures acquired during solidification can produce anisotropy if crystals preferentially grow with particular crystallographic axes aligned with the radial heat flux; outward radial cooling during freezing would then bias crystal orientations. Karato (1998) suggested a further contribution from magnetically driven stresses: secular changes in the geomagnetic field could impose long‑term stresses that deform the inner core and progressively reorient crystals.
These hypotheses are not mutually exclusive; differential freezing–driven creep, thermally driven flow, growth‑aligned solidification, and magnetically induced deformation can each contribute to lattice‑preferred orientation under different conditions. Discriminating among them requires tighter quantitative constraints on inner‑core viscosity, the inner core’s temporal growth history, spatial patterns of solidification and heat flux, and the evolution of the geomagnetic field.
Seismic investigations over the past two decades have suggested that Earth’s solid inner core may itself be heterogeneous, with an innermost region whose properties differ from the surrounding inner-core material. Early evidence for such an “innermost inner core” (IMIC) was presented by Ishii and Dziewoński (2002), and subsequent work has continued to refine models of internal structure while leaving key parameters—most notably the IMIC radius—in dispute.
Proposed geometries vary. Wang and Song (2018) advanced a three-layer scheme consisting of a central “inner inner core” (IIC) of roughly 500 km radius, an outer inner-core (OIC) layer on the order of 600 km thick, and an outermost isotropic shell approximately 100 km thick; in their reconstruction the fast P-wave symmetry axis is aligned with the rotation axis in the OIC but rotated relative to it within the IIC. A 2023 seismic study described an alternative picture of a roughly 650-km-diameter innermost ball with strong anisotropic character transitioning to a weakly anisotropic outer shell, and interpreted this pattern as potentially preserving a record of a major ancient global event.
Mechanistically, recent analyses attribute the observed directional variation of P-wave speeds to slight differences in atomic packing or crystal-preferred orientation within the deepest core, producing seismic anisotropy. Quantitatively, the 2023 work reports P-wave speeds about 4% slower at roughly 50° from the rotation axis inside the innermost region compared with other directions, indicating a systematic angular dependence of compressional-wave velocity.
Despite these model-specific descriptions, published estimates span a broad numeric range (IMIC radii proposed between ~300 km and ~750 km across studies from 2002–2019, with specific proposals at ~500 km and ~650 km in 2018–2023 work), and interpretive debate remains active. Several authors argue that the seismic signatures invoked to justify discrete concentric sublayers can instead be produced by a continuous radial gradation of material properties; thus, the central question is whether the deepest core contains sharp internal boundaries or a smooth depth-dependent evolution of texture and anisotropy.
Seismic studies of the inner core have long probed lateral heterogeneity by testing for directional dependence of wave speeds (anisotropy) and its geographic variability. A widely cited 1997 interpretation by Tanaka and Hamaguchi inferred a north–south fast axis and a pronounced east–west hemispheric contrast, with stronger anisotropy beneath an “eastern” hemisphere centered near 110°E (approximately under Borneo) than beneath a “western” hemisphere near 70°W (approximately under Colombia). Geophysically, such observations were taken to indicate a preferred crystallographic alignment of the iron alloy in the solid inner core—faster seismic propagation along polar directions than along the equator—and spatial changes in that lattice-preferred orientation with longitude.
Mechanistic hypotheses have been proposed to account for the putative hemispheric asymmetry. Alboussère et al. suggested that asymmetric melting and recrystallization at inner‑core depths (preferential melting in the eastern hemisphere with compensating recrystallization in the west) could generate contrasting microstructures and hence contrasting anisotropy. Extending this idea, Finlay argued that such hemispheric-scale material processing might couple to larger‑scale geodynamo behavior and thereby offer a potential explanation for observed hemispheric differences in the geomagnetic field.
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More recent reanalysis complicates the picture. Frost and Romanowicz (2017) reported only weak north–south anisotropy—seismic speeds faster by roughly 0.5–1.5% along the axis relative to equatorial directions—and found no robust east–west contrast at the longitudes previously emphasized. Together, these divergent results underscore that estimates of inner‑core anisotropy and its lateral variability remain sensitive to data selection and modeling choices. The proposed causal chain from localized inner‑core microphysics beneath ~110°E and ~70°W to hemispheric magnetic asymmetries therefore remains a plausible but unproven hypothesis, requiring additional seismic constraints, mineral‑physics experiments, and geomagnetic analyses to be tested.
High‑resolution seismological studies indicate that the inner‑core boundary (ICB) may be laterally heterogeneous at very fine spatial scales, with reported variations down to about 1 km. Such small‑scale heterogeneity is unexpected because independent constraints from Earth’s magnetic field imply that lateral temperature contrasts along the ICB are extremely small; geomagnetic observations therefore strongly limit thermally driven variability at these depths.
If kilometer‑scale contrasts at the ICB are real despite negligible temperature gradients, they must reflect non‑thermal causes—for example, compositional differences, variations in microstructure or crystal orientation, small‑scale phase or solidification textures, or seismic scattering from discrete features. Resolving this apparent contradiction demands an integrated program combining high‑resolution seismic imaging, geomagnetic modeling, laboratory mineral‑physics on core materials, and numerical core‑dynamics simulations to test whether and how ~1 km non‑thermal heterogeneities can form and be maintained at the ICB.
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Growth
The inner core is a solid iron-rich sphere immersed in a convecting, electrically conducting liquid outer core; interactions between their motions underpin the geodynamo that sustains Earth’s magnetic field. Over geological time the planet’s deep interior cools (on the order of ~100 °C per billion years), causing liquid at the inner-core boundary to freeze and the solid inner core to expand gradually.
Crystallization preferentially incorporates iron into the solid phase, leaving the adjacent liquid relatively enriched in lighter components (notably oxygen). This compositional differentiation lowers the density of the boundary-layer fluid relative to the surrounding outer core, producing a buoyancy anomaly that drives upward flow. Such compositionally driven buoyancy is a primary source of convective motion in the outer core and thus provides the mechanical and electrical circulation required by the geodynamo.
Thermodynamic and geophysical calculations (e.g., Alfé and colleagues) support that inner-core crystallization can generate an oxygen-enriched, low-density layer capable of becoming convectively unstable and sustaining outer-core circulation. Because the growing solid inner core alters the boundary conditions and flow geometry of the outer core, its presence and evolution influence the morphology and temporal stability of the geomagnetic field and may help fix certain magnetic features—an effect that remains under active investigation.
Dynamics
The Earth’s solid inner core is mechanically decoupled from the overlying mantle and crust, permitting it to rotate at a rate that can differ modestly from the bulk planet (either faster or slower). Seismology has been central to detecting this differential motion: the inner core is seismically anisotropic, and multi-decade comparisons of travel-time patterns—most robustly using “seismic doublets” (repeated earthquakes from the same source recorded years apart at the same stations)—allow inference of slow relative rotation. Early analyses in the 1990s suggested a super-rotation on the order of one degree per year; subsequent reanalysis of doublets in 2005 revised that estimate downward to roughly 0.3–0.5° yr−1. More recent work indicates a change in behaviour around 2009, with the inner core no longer spinning faster than the surface and now likely rotating more slowly; this evolution appears oscillatory with a multi‑decadal period of about seven decades and is not expected to have major direct effects on surface processes. External tidal torques from the Sun and Moon, which primarily act on the solid Earth and drive secular changes in day length, produce differential forcing that can decouple the rotation of the crust and mantle from that of the fluid outer core and inner core. Inside the core region, the situation is further complicated by the coupling of fluid motions, induced electrical currents, and magnetic fields; theoretical work has predicted small daily nutations of the inner-core axis and a rich spectrum of dynamical behaviour. Under particular models of Earth history, resonant interactions between tidal forcing and outer-core flow may have occurred episodically (proposed at about 3.0, 1.8 and 0.3 billion years ago, each lasting some 200–300 million years), during which enhanced fluid motion could generate sufficient dissipation to affect the thermal balance and transiently impede inner‑core growth. Together these observations and models portray inner‑core rotation as a slowly evolving, magnetohydrodynamically coupled phenomenon subject to ongoing refinement.
The solid inner core is believed to have grown by crystallization as Earth’s initially liquid core cooled, but the timing of nucleation remains unresolved. Two broad approaches constrain its age: thermodynamic cooling models (T), which compute when crystallization becomes thermodynamically favorable and are often run either with (R) or without (N) contributions from radiogenic heat, and paleomagnetic analyses (P), which infer core state from the long-term behavior of Earth’s magnetic field.
Published estimates illustrate the methodological divergence. Thermodynamic calculations yield ages sensitive to assumptions about heat sources and core composition: Labrosse (2001, T(N)) inferred ~1.0 ± 0.5 Ga, while inclusion of radiogenic heating produced older values (Labrosse 2003, T(R) ≈ 2 Ga). More recent thermodynamic studies favor later nucleation in some parameterizations (Driscoll & Bercovici 2014, T ≈ 0.65 Ga; Labrosse 2015 and Ohta et al. 2016, T < 0.7 Ga), whereas broader uncertainty bounds allow much earlier onset (Konôpková et al. 2016, upper bound < 4.2 Ga).
Paleomagnetic constraints also vary: Smirnov et al. (2011, P) proposed a 2–3.5 Ga interval, Biggin et al. (2015, P) estimated 1–1.5 Ga, and Bono et al. (2019, P) inferred as young as ≈0.5 Ga. Reported ages across studies range from about 0.5 Ga to upper bounds near 4.2 Ga, but many syntheses adopt a working interval of roughly 0.5–2.0 Ga. The wide spread reflects differences in paleomagnetic interpretation, datasets, and especially the strong sensitivity of thermodynamic models to assumptions about radiogenic heating, core composition, and the planet’s cooling history, leaving inner-core nucleation an active area of geophysical research.
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Thermodynamic evidence for inner‑core age
Thermal-evolution models constrain the age of Earth’s solid inner core by combining two observationally anchored requirements: (1) a minimum outward heat flux across the core–mantle boundary (CMB) sufficient to drive convection in the liquid outer core and thus sustain the geodynamo, and (2) empirical estimates of present-day CMB heat flux that are linked quantitatively to measured surface heat flow and observed rates of mantle convection. Imposing the geodynamo threshold on cooling histories ties models of inner‑core nucleation to directly measurable geophysical quantities.
Early models that assumed negligible radioactivity in the core (Labrosse et al., ~2001–2003) produced a relatively young inner core, ca. 1.0 ± 0.5 Ga, with the authors noting that modest radiogenic heating would shift this age upward by a few 10^8 years. Subsequent revisions of core transport properties strongly altered these conclusions. Theoretical calculations (Pozzo et al., 2012) and laboratory confirmation (Gomi et al., 2013) indicated electrical—and by implication thermal—conductivities of core materials several times larger than earlier values, implying thermal conductivities approaching ~90 W m−1 K−1. Incorporating such high conductivities into thermal histories concentrates heat transport by conduction, reduces the available convective buoyancy, and forces modeled inner‑core nucleation to much more recent times (typically <700 Ma).
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Conversely, direct conductivity measurements on solid iron at inner‑core conditions (Konôpková et al., 2016) yielded substantially lower thermal conductivities (≈18–44 W m−1 K−1). When these lower values are used, conductive cooling is less efficient, allowing prolonged convective heat loss and permitting an inner core substantially older—models constrained by these conductivities can accommodate an inner‑core age as old as ~4.2 Ga, consistent with paleomagnetic indications of a long‑lived geodynamo.
An alternative resolution of model tensions was proposed by Driscoll and Bercovici (2014), who invoked ~3 TW of radiogenic heating from 40K in the core to avoid a “mantle thermal catastrophe” and reconcile a long‑lived dynamo with high conductivity; this hypothesis requires mantle–core partitioning that supplies very large potassium abundances to the core, a requirement not supported by most experimental partitioning data and thus remains contentious.
Overall, thermodynamic constraints on inner‑core age are highly sensitive to three coupled factors: the minimum CMB heat flux required to sustain core convection and the geodynamo, the thermal (and electrical) conductivity of core materials under extreme pressure–temperature conditions (reported values span ≈18–44 to ≈90 W m−1 K−1), and any radiogenic heat contribution from the core (e.g., a hypothesized ~3 TW from 40K). Variations and uncertainties in these parameters yield permissible inner‑core ages that range from less than ≈700 Ma to of order 4.2 Ga, with direct consequences for the timing, duration, and driving mechanisms of Earth’s magnetic field.
Paleomagnetic evidence
Paleomagnetism reconstructs past geomagnetic field behavior from remanent magnetization preserved in dated rocks; because the presence and growth of a solid inner core alters core convection patterns and electromagnetic coupling, changes in field strength, variability and geometry preserved in the rock record provide indirect constraints on the timing and dynamical consequences of inner‑core nucleation. Analyses of Neoarchean versus Proterozoic rocks by Smirnov et al. (2011) found a closer approach to an ideal axial dipole in the Neoarchean than in later Proterozoic intervals, a result they interpret as a shift in dynamo activity from deeper regions toward currents nearer the core‑mantle boundary—consistent with inner‑core growth sometime between ~3.5 and 2.0 Ga. Biggin et al. (2015), using a larger, carefully screened Precambrian compilation, identified a marked increase in both field intensity and its variance around 1.0–1.5 Ga; they proposed this signature as possible evidence for inner‑core nucleation at that time and inferred a relatively low outer‑core thermal conductivity compatible with simple thermal‑evolution models.
Numerical models that couple core thermal evolution and dynamo action provide a temporal framework linking these paleomagnetic signals to specific dynamo regimes. Driscoll (2016), building on thermal histories from Driscoll and Bercovici (2014), produced an “evolving dynamo” that predicts distinct regimes through the past two billion years: an early strong, multipolar dynamo before ~1.7 Ga; a strong, predominantly dipolar state from ~1.0–1.7 Ga; a weaker, non‑axial dipole mode from ~0.6–1.0 Ga; and, after inner‑core nucleation, a renewed strong dipolar field in the past ~0.6 Ga. Empirical observations from the Ediacaran (~565 Ma) reported by Bono et al. (2019)—anomalously low intensities, two discrete directions, and independently inferred high reversal rates—align with model expectations for a transitional dynamo and have been interpreted (by Bono and commentators, and summarized by Driscoll) as recording the onset of inner‑core formation near ~0.5 Ga; subsequent Cambrian data have been reported as consistent with this relatively young nucleation scenario.
Taken together, these independent lines—Neoarchean versus Proterozoic geometry shifts, the 1.0–1.5 Ga intensity/variance increase, the thermal‑history driven evolving dynamo, and the late Neoproterozoic–early Phanerozoic paleomagnetic anomalies—converge on overlapping but not identical windows for inner‑core nucleation ranging from the Proterozoic into the late Neoproterozoic/early Phanerozoic. The competing hypotheses carry directly testable implications: the inferred nucleation age constrains permissible outer‑core thermal conductivities, determines whether the long‑term field should be dominantly dipolar or multipolar (and axial versus non‑axial) at given times, and shapes models of Earth’s thermal and magnetic evolution. Continued expansion of high‑quality paleomagnetic datasets, together with coupled thermal‑dynamo modeling, remains essential to narrowing the nucleation age and its geodynamic consequences.