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Planetary core
Planetary core
Planetary core
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The internal structure of the inner planets.
The internal structure of the outer planets.

A planetary core consists of the innermost layers of a planet.[1] Cores may be entirely liquid, or a mixture of solid and liquid layers as is the case in the Earth.[2] In the Solar System, core sizes range from about 20% (the Moon) to 85% of a planet's radius (Mercury).

Gas giants also have cores, though the composition of these are still a matter of debate and range in possible composition from traditional stony/iron, to ice or to fluid metallic hydrogen.[3][4][5] Gas giant cores are proportionally much smaller than those of terrestrial planets, though they can be considerably larger than the Earth's nevertheless; Jupiter's is 10–30 times heavier than Earth,[5] and exoplanet HD149026 b may have a core 100 times the mass of the Earth.[6]

Planetary cores are challenging to study because they are impossible to reach by drill and there are almost no samples that are definitively from the core. Thus, they are studied via indirect techniques such as seismology, mineral physics, and planetary dynamics.

Discovery

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Earth's core

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Main article: Earth's core

In 1797, Henry Cavendish calculated the average density of the Earth to be 5.48 times the density of water (later refined to 5.53), which led to the accepted belief that the Earth was much denser in its interior.[7] Following the discovery of iron meteorites, Wiechert in 1898 postulated that the Earth had a similar bulk composition to iron meteorites, but the iron had settled to the interior of the Earth, and later represented this by integrating the bulk density of the Earth with the missing iron and nickel as a core.[8] The first detection of Earth's core occurred in 1906 by Richard Dixon Oldham upon discovery of the P-wave shadow zone; the liquid outer core.[9] By 1936 seismologists had determined the size of the overall core as well as the boundary between the fluid outer core and the solid inner core.[10]

Moon's core

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The internal structure of the Moon was characterized in 1974 using seismic data collected by the Apollo missions of moonquakes.[11] The Moon's core has a radius of 300 km.[12] The Moon's iron core has a liquid outer layer that makes up 60% of the volume of the core, with a solid inner core.[13]

Cores of the rocky planets

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The cores of the rocky planets were initially characterized by analyzing data from spacecraft, such as NASA's Mariner 10 that flew by Mercury and Venus to observe their surface characteristics.[14] The cores of other planets cannot be measured using seismometers on their surface, so instead they have to be inferred based on calculations from these fly-by observation. Mass and size can provide a first-order calculation of the components that make up the interior of a planetary body. The structure of rocky planets is constrained by the average density of a planet and its moment of inertia.[15] The moment of inertia for a differentiated planet is less than 0.4, because the density of the planet is concentrated in the center.[16] Mercury has a moment of inertia of 0.346, which is evidence for a core.[17] Conservation of energy calculations as well as magnetic field measurements can also constrain composition, and surface geology of the planets can characterize differentiation of the body since its accretion.[18] Mercury, Venus, and Mars’ cores are about 75%, 50%, and 40% of their radius respectively.[19][20]

Formation

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Accretion

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Planetary systems form from flattened disks of dust and gas that accrete rapidly (within thousands of years) into planetesimals around 10 km in diameter. From here gravity takes over to produce Moon- to Mars-sized planetary embryos over 0.1–1 million years, and these develop into planetary bodies over an additional 10–100 million years.[21]

Jupiter and Saturn most likely formed around previously existing rocky and/or icy bodies, rendering these previous primordial planets into gas-giant cores.[5] This is the planetary core accretion model of planet formation.

Differentiation

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Planetary differentiation is broadly defined as the development from one thing to many things; homogeneous body to several heterogeneous components.[22] The hafnium-182/tungsten-182 isotopic system has a half-life of 9 million years, and is approximated as an extinct system after 45 million years. Hafnium is a lithophile element and tungsten is siderophile element. Thus if metal segregation (between the Earth's core and mantle) occurred in under 45 million years, silicate reservoirs develop positive Hf/W anomalies, and metal reservoirs acquire negative anomalies relative to undifferentiated chondrite material.[21] The observed Hf/W ratios in iron meteorites constrain metal segregation to under 5 million years, the Earth's mantle Hf/W ratio places Earth's core as having segregated within 25 million years.[21] Several factors control segregation of a metal core including the crystallization of perovskite. Crystallization of perovskite in an early magma ocean is an oxidation process and may drive the production and extraction of iron metal from an original silicate melt.

Core merging and impacts

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Impacts between planet-sized bodies in the early Solar System are important aspects in the formation and growth of planets and planetary cores.

Earth–Moon system

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The giant impact hypothesis states that an impact between a theoretical Mars-sized planet Theia and the early Earth formed the modern Earth and Moon.[23] During this impact the majority of the iron from Theia and the Earth became incorporated into the Earth's core.[24]

Mars

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Core merging between the proto-Mars and another differentiated planetoid could have been as fast as 1000 years or as slow as 300,000 years (depending on viscosity of both cores).[25]

Chemistry

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Determining primary composition – Earth

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Using the chondritic reference model and combining known compositions of the crust and mantle, the unknown component, the composition of the inner and outer core, can be determined: 85% Fe, 5% Ni, 0.9% Cr, 0.25% Co, and all other refractory metals at very low concentration.[21] This leaves Earth's core with a 5–10% weight deficit for the outer core,[26] and a 4–5% weight deficit for the inner core;[26] which is attributed to lighter elements that should be cosmically abundant and are iron-soluble; H, O, C, S, P, and Si.[21] Earth's core contains half the Earth's vanadium and chromium, and may contain considerable niobium and tantalum.[26] Earth's core is depleted in germanium and gallium.[26]

Weight deficit components – Earth

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Sulfur is strongly siderophilic and only moderately volatile and depleted in the silicate earth; thus may account for 1.9 weight % of Earth's core.[21] By similar arguments, phosphorus may be present up to 0.2 weight %. Hydrogen and carbon, however, are highly volatile and thus would have been lost during early accretion and therefore can only account for 0.1 to 0.2 weight % respectively.[21] Silicon and oxygen thus make up the remaining mass deficit of Earth's core; though the abundances of each are still a matter of controversy revolving largely around the pressure and oxidation state of Earth's core during its formation.[21] No geochemical evidence exists to include any radioactive elements in Earth's core.[26] Despite this, experimental evidence has found potassium to be strongly siderophilic at the temperatures associated with core formation, thus there is potential for potassium in planetary cores of planets, and therefore potassium-40 as well.[27]

Isotopic composition – Earth

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Hafnium/tungsten (Hf/W) isotopic ratios, when compared with a chondritic reference frame, show a marked enrichment in the silicate earth indicating depletion in Earth's core. Iron meteorites, believed to be resultant from very early core fractionation processes, are also depleted.[21] Niobium/tantalum (Nb/Ta) isotopic ratios, when compared with a chondritic reference frame, show mild depletion in bulk silicate Earth and the moon.[28]

Pallasite meteorites

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Pallasites are thought to form at the core-mantle boundary of an early planetesimal, although a recent hypothesis suggests that they are impact-generated mixtures of core and mantle materials.[29]

Dynamics

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Dynamo

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Dynamo theory is a proposed mechanism to explain how celestial bodies like the Earth generate magnetic fields. The presence or lack of a magnetic field can help constrain the dynamics of a planetary core. Refer to Earth's magnetic field for further details. A dynamo requires a source of thermal and/or compositional buoyancy as a driving force.[28] Thermal buoyancy from a cooling core alone cannot drive the necessary convection as indicated by modelling, thus compositional buoyancy (from changes of phase) is required. On Earth the buoyancy is derived from crystallization of the inner core (which can occur as a result of temperature). Examples of compositional buoyancy include precipitation of iron alloys onto the inner core and liquid immiscibility both, which could influence convection both positively and negatively depending on ambient temperatures and pressures associated with the host-body.[28] Other celestial bodies that exhibit magnetic fields are Mercury, Jupiter, Ganymede, and Saturn.[3]

Core heat source

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A planetary core acts as a heat source for the outer layers of a planet. In the Earth, the heat flux over the core mantle boundary is 12 terawatts.[30] This value is calculated from a variety of factors: secular cooling, differentiation of light elements, Coriolis forces, radioactive decay, and latent heat of crystallization.[30] All planetary bodies have a primordial heat value, or the amount of energy from accretion. Cooling from this initial temperature is called secular cooling, and in the Earth the secular cooling of the core transfers heat into an insulating silicate mantle.[30] As the inner core grows, the latent heat of crystallization adds to the heat flux into the mantle.[30]

Stability and instability

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Small planetary cores may experience catastrophic energy release associated with phase changes within their cores. Ramsey (1950) found that the total energy released by such a phase change would be on the order of 1029 joules; equivalent to the total energy release due to earthquakes through geologic time. Such an event could explain the asteroid belt. Such phase changes would only occur at specific mass to volume ratios, and an example of such a phase change would be the rapid formation or dissolution of a solid core component.[31]

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Inner rocky planets

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All of the rocky inner planets, as well as the moon, have an iron-dominant core. Venus and Mars have an additional major element in the core. Venus’ core is believed to be iron-nickel, similarly to Earth. Mars, on the other hand, is believed to have an iron-sulfur core and is separated into an outer liquid layer around an inner solid core.[20] As the orbital radius of a rocky planet increases, the size of the core relative to the total radius of the planet decreases.[15] This is believed to be because differentiation of the core is directly related to a body's initial heat, so Mercury's core is relatively large and active.[15] Venus and Mars, as well as the moon, do not have magnetic fields. This could be due to a lack of a convecting liquid layer interacting with a solid inner core, as Venus’ core is not layered.[19] Although Mars does have a liquid and solid layer, they do not appear to be interacting in the same way that Earth's liquid and solid components interact to produce a dynamo.[20]

Outer gas and ice giants

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Current understanding of the outer planets in the solar system, the ice and gas giants, theorizes small cores of rock surrounded by a layer of ice, and in Jupiter and Saturn models suggest a large region of liquid metallic hydrogen and helium.[19] The properties of these metallic hydrogen layers is a major area of contention because it is difficult to produce in laboratory settings, due to the high pressures needed.[32] Jupiter and Saturn appear to release a lot more energy than they should be radiating just from the sun, which is attributed to heat released by the hydrogen and helium layer. Uranus does not appear to have a significant heat source, but Neptune has a heat source that is attributed to a “hot” formation.[19]

Observed types

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The following summarizes known information about the planetary cores of given non-stellar bodies.

Within the Solar System

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Mercury

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Mercury has an observed magnetic field, which is believed to be generated within its metallic core.[28] Mercury's core occupies 85% of the planet's radius, making it the largest core relative to the size of the planet in the Solar System; this indicates that much of Mercury's surface may have been lost early in the Solar System's history.[33] Mercury has a solid silicate crust and mantle overlying a solid metallic outer core layer, followed by a deeper liquid core layer, and then a possible solid inner core making a third layer.[33] The composition of the iron-rich core remains uncertain, but it likely contains nickel, silicon and perhaps sulfur and carbon, plus trace amounts of other elements.[34]

Venus

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The composition of Venus' core varies significantly depending on the model used to calculate it, thus constraints are required.[35]

Element Chondritic Model Equilibrium Condensation Model Pyrolitic Model
Iron 88.6% 94.4% 78.7%
Nickel 5.5% 5.6% 6.6%
Cobalt 0.26% Unknown Unknown
Sulfur 5.1% 0% 4.9%
Oxygen 0% Unknown 9.8%

Moon

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The existence of a lunar core is still debated; however, if it does have a core it would have formed synchronously with the Earth's own core at 45 million years post-start of the Solar System based on hafnium-tungsten evidence[36] and the giant impact hypothesis. Such a core may have hosted a geomagnetic dynamo early on in its history.[28]

Earth

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Main article: Structure of the Earth § Core

The Earth has an observed magnetic field generated within its metallic core.[28] The Earth has a 5–10% mass deficit for the entire core and a density deficit from 4–5% for the inner core.[26] The Fe/Ni value of the core is well constrained by chondritic meteorites.[26] Sulfur, carbon, and phosphorus only account for ~2.5% of the light element component/mass deficit.[26] No geochemical evidence exists for including any radioactive elements in the core.[26] However, experimental evidence has found that potassium is strongly siderophile when dealing with temperatures associated with core-accretion, and thus potassium-40 could have provided an important source of heat contributing to the early Earth's dynamo, though to a lesser extent than on sulfur rich Mars.[27] The core contains half the Earth's vanadium and chromium, and may contain considerable niobium and tantalum.[26] The core is depleted in germanium and gallium.[26] Core mantle differentiation occurred within the first 30 million years of Earth's history.[26] Inner core crystallization timing is still largely unresolved.[26]

Mars

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Mars possibly hosted a core-generated magnetic field in the past.[28] The dynamo ceased within 0.5 billion years of the planet's formation.[2] Hf/W isotopes derived from the martian meteorite Zagami, indicate rapid accretion and core differentiation of Mars; i.e. under 10 million years.[23] Potassium-40 could have been a major source of heat powering the early Martian dynamo.[27]

Core merging between proto-Mars and another differentiated planetoid could have been as fast as 1000 years or as slow as 300,000 years (depending on the viscosity of both cores and mantles).[25] Impact-heating of the Martian core would have resulted in stratification of the core and kill the Martian dynamo for a duration between 150 and 200 million years.[25] Modelling done by Williams, et al. 2004 suggests that in order for Mars to have had a functional dynamo, the Martian core was initially hotter by 150 K than the mantle (agreeing with the differentiation history of the planet, as well as the impact hypothesis), and with a liquid core potassium-40 would have had opportunity to partition into the core providing an additional source of heat. The model further concludes that the core of mars is entirely liquid, as the latent heat of crystallization would have driven a longer-lasting (greater than one billion years) dynamo.[2] If the core of Mars is liquid, the lower bound for sulfur would be five weight %.[2]

Ganymede

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Ganymede has an observed magnetic field generated within its metallic core.[28]

Jupiter

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Jupiter has an observed magnetic field generated within its core, indicating some metallic substance is present.[3] Its magnetic field is the strongest in the Solar System after the Sun's.

Jupiter has a rock and/or ice core 10–30 times the mass of the Earth, and this core is likely soluble in the gas envelope above, and so primordial in composition. Since the core still exists, the outer envelope must have originally accreted onto a previously existing planetary core.[5] Thermal contraction/evolution models support the presence of metallic hydrogen within the core in large abundances (greater than Saturn).[3]

Saturn

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Saturn has an observed magnetic field generated within its metallic core.[3] Metallic hydrogen is present within the core (in lower abundances than Jupiter).[3] Saturn has a rock and or ice core 10–30 times the mass of the Earth, and this core is likely soluble in the gas envelope above, and therefore it is primordial in composition. Since the core still exists, the envelope must have originally accreted onto previously existing planetary cores.[5] Thermal contraction/evolution models support the presence of metallic hydrogen within the core in large abundances (but still less than Jupiter).[3]

Remnant planetary cores

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Missions to bodies in the asteroid belt will provide more insight to planetary core formation. It was previously understood that collisions in the solar system fully merged, but recent work on planetary bodies argues that remnants of collisions have their outer layers stripped, leaving behind a body that would eventually become a planetary core.[37] The Psyche mission, titled “Journey to a Metal World,” is aiming to studying a body that could possibly be a remnant planetary core.[38]

Extrasolar

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As the field of exoplanets grows as new techniques allow for the discovery of both diverse exoplanets, the cores of exoplanets are being modeled. These depend on initial compositions of the exoplanets, which is inferred using the absorption spectra of individual exoplanets in combination with the emission spectra of their star.

Chthonian planets

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A chthonian planet results when a gas giant has its outer atmosphere stripped away by its parent star, likely due to the planet's inward migration. All that remains from the encounter is the original core.

Planets derived from stellar cores and diamond planets

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Carbon planets, previously stars, are formed alongside the formation of a millisecond pulsar. The first such planet discovered was 18 times the density of water, and five times the size of Earth. Thus the planet cannot be gaseous, and must be composed of heavier elements that are also cosmically abundant like carbon and oxygen; making it likely crystalline like a diamond.[39]

PSR J1719-1438 is a 5.7 millisecond pulsar found to have a companion with a mass similar to Jupiter but a density of 23 g/cm3, suggesting that the companion is an ultralow mass carbon white dwarf, likely the core of an ancient star.[40]

Hot ice planets

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Exoplanets with moderate densities (more dense than Jovian planets, but less dense than terrestrial planets) suggests that such planets like GJ1214b and GJ436 are composed of primarily water. Internal pressures of such water-worlds would result in exotic phases of water forming on the surface and within their cores.[41]

References

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Revisions and contributorsEdit on WikipediaRead on Wikipedia

Planetary core

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The planetary core is the central, densest region of a differentiated , primarily composed of heavy metallic elements such as iron and that sink toward the center during the planet's formation due to gravitational settling in a molten state. This process, known as , requires sufficient heat—typically exceeding 1,300 K—to allow denser materials to separate from lighter silicates, resulting in a layered internal structure that persists as the planet cools. Cores vary significantly in size and makeup across planetary types, influencing a 's overall , , and potential for . In terrestrial planets like , Mercury, , and Mars, the core typically consists of an iron-nickel alloy, often with trace lighter elements such as or . 's core, for instance, accounts for about one-third of the planet's mass and features a solid inner core surrounded by a liquid outer core, where the inner portion is primarily iron alloyed with in exotic forms under extreme pressures exceeding 840,000 atmospheres and temperatures around 4,200°C. Mercury exemplifies an extreme case, with its core comprising roughly 85% of the planet's volume and dominated by iron, while smaller bodies like the have cores estimated at 20% of their volume, containing iron, , and . Gas giant planets, such as and Saturn, possess deeper, more massive cores of rock, metal, and ice—estimated at 10–25 masses—surrounded by vast envelopes of and that constitute about 75% and 25% of their atmospheres by , respectively. These cores form the seed for the planet's growth in protoplanetary disks, accreting lighter gases once reaching sufficient to trigger runaway accumulation. In contrast to terrestrial cores, those of gas giants are enveloped by layers, and their exact compositions remain under study through missions like Juno, which probe elemental distributions including potential water content. Planetary cores play a crucial role in generating dynamo effects that produce , which protect atmospheres from stellar winds and enable conditions for , as seen in Earth's core-driven . Ongoing research, including 's Psyche mission targeting a metallic core remnant, continues to refine models of core formation and evolution across the solar system and beyond.

Introduction

Definition and general properties

A planetary core is the innermost layer of a differentiated planet, formed through the gravitational segregation of denser materials toward the center during planetary formation. It primarily consists of heavy elements such as iron and nickel, which sink below lighter materials to form this central region, often surrounded by a mantle and, in terrestrial , a crust. In terrestrial () planets, cores exhibit high densities, typically ranging from 10 to 13 g/cm³ due to their iron-rich composition under extreme . These cores often feature a solid inner core surrounded by a liquid outer core, with sizes varying widely: for example, 's core has a radius of approximately 3,480 km (about 55% of the planet's radius) and accounts for roughly 32% of its mass, while Mercury's disproportionately large core occupies about 85% of its volume. The state of the core—solid or liquid—depends on , , and composition, with the liquid outer core enabling convective motions in many cases. In contrast, the cores of (Jovian) planets are more diffuse and less well-defined, often described as "fuzzy" due to gradual transitions in composition rather than sharp boundaries. These cores consist of heavier elements like rock, metals, and ices, with estimated masses of 10–20 masses for and Saturn, potentially extending to several tens of percent of the planet's radius in fuzzy core models, as revealed by recent observations from NASA's Juno mission (as of 2025). Their densities are lower overall than terrestrial cores, blending into surrounding layers of and . The presence of a dense core affects a planet's mass distribution and , as quantified by the normalized I/MR2I / MR^2, where II is the , MM the , and RR the radius. For a uniform-density , this factor is 0.4, but Earth-like bodies with concentrated cores yield values around 0.33, indicating central excess that shapes oblateness and gravitational harmonics.

Significance in planetary evolution

Planetary cores drive the process of differentiation, where denser metallic materials sink to form a central core, resulting in layered planetary structures consisting of a core, mantle, and crust. This separation occurs primarily during early accretion when high temperatures from impacts and facilitate metal-silicate . Core formation thus establishes the foundational architecture of terrestrial , with initial temperatures and core segregation playing pivotal roles in the timing and extent of this layering. within the core and overlying mantle, sustained by residual heat from differentiation, generates buoyancy-driven flows that promote mantle upwelling and , shaping surface and volatile release over billions of years. Cores enable the generation of dynamo-driven through convective motions in their outer layers, which shield planetary atmospheres from erosion by stellar winds and cosmic radiation. On , this geodynamo, active for at least 4.2 billion years, has preserved the atmosphere and , creating stable conditions essential for the emergence and persistence of . Such protection is crucial for long-term , as the absence of a can lead to significant atmospheric loss, altering climate and surface environments. Core-mantle interactions exert torque and dissipative coupling that influence planetary rotation rates and obliquity, the determining seasonal variations. These couplings, arising from viscous and gradients at the core-mantle boundary, can dampen or excite rotational variations, stabilizing or altering the planet's spin axis over geological timescales. For instance, inertial and dissipative effects from core-mantle contribute to secular changes in obliquity for both and Mars, affecting orbital dynamics and surface climate patterns. As repositories of primordial heat from accretion and core formation, planetary cores release thermal energy gradually through conduction and convection, sustaining internal dynamics and geochemical processes. Cores also sequester siderophile elements like , , and , which, through partitioning during differentiation, deplete the mantle and influence long-term cycles of , , and isotopic in the silicate portions. This sequestration affects mantle heterogeneity and the recycling of elements via and , driving planetary geochemical differentiation. Core solidification, as a late evolutionary stage, can halt and activity, leading to the shutdown of and exposing the planet to atmospheric stripping. On Mars, core cooling and partial solidification contributed to the cessation of its around 3.7 billion years ago. This process underscores how core evolution dictates the trajectory of and surface modification.

History of Discovery

Earth's core

The discovery of Earth's core began in the early 20th century through analyses of seismic wave propagation from earthquakes. In 1906, Richard Dixon Oldham identified a "shadow zone" in seismograms where shear (S) waves were absent beyond about 100–105 degrees from the epicenter, suggesting a central region that did not transmit these waves, indicative of a liquid-like interior. This was refined in by Beno Gutenberg, who used travel-time data from multiple earthquakes to pinpoint the core-mantle boundary (CMB) at approximately 2900 km depth, marking the abrupt increase in seismic velocity and density at the interface. Gutenberg's work established the core as a distinct layer comprising about half of Earth's radius, with P (compressional) waves refracting sharply at the CMB while S waves ceased entirely, creating the shadow zone due to the outer core's inability to support . Further evidence emerged in 1936 when Inge Lehmann, analyzing reflected P waves (PKiKP phases) in the shadow zone, detected a discontinuity at around 5100 km depth, proposing a solid inner core surrounded by the liquid outer core. These reflections indicated higher P-wave velocities in the inner core compared to the outer, consistent with a from liquid to solid. Early models interpreted the outer core as liquid iron-rich material, explaining the absence of S waves (effectively zero shear velocity) and the of P waves at reduced speeds within it. The inner core was modeled as solid, with increased rigidity allowing P-wave propagation and reflection, though its exact size and boundary were initially imprecise due to sparse data. Prior to 2000, these models faced historical debates, particularly on the core's overall state and dimensions. Early 20th-century arguments, led by figures like Emil Wiechert, favored a fully solid core to match density estimates, but the lack of S-wave transmission prompted contention, resolved in 1926 by who confirmed the outer core's liquidity through refined travel-time curves. Variations in estimated CMB depth (ranging 2700–3000 km) and inner core radius persisted until better global seismograph networks in the mid-20th century narrowed them. Confirmations came from Earth's free oscillations excited by large earthquakes, such as the 1960 event; spectral analysis in the 1970s revealed mode splittings consistent with a solid inner core's elasticity, supporting Lehmann's boundary at ~5100 km. Parallel laboratory experiments on iron alloys under high pressure and temperature, including shock-wave studies in the 1950s–1980s, demonstrated melting curves where iron solidifies at inner core conditions (above ~300 GPa and 5000 K), aligning seismic velocities with a solid iron-nickel phase. By the late 1990s, consensus solidified around the dual-structure model, with the outer core spanning ~2200 km thick and the inner core ~1200 km in radius.

Cores in other Solar System bodies

The detection of planetary cores beyond has historically relied on indirect methods due to the challenges of deploying seismometers on most bodies, unlike the direct seismic networks available for . Orbital measurements of fields, magnetic signatures, and rotational parameters have been primary tools, but they often yield ambiguous results requiring models to infer core properties, with limitations in resolution compared to seismic wave propagation that can directly probe density contrasts and phase boundaries. For the Moon, initial evidence of a small core emerged from analysis of seismic data collected by the Apollo Passive Seismic Experiment in the 1970s, which detected low-velocity zones suggestive of a partially molten interior structure. Reinterpretation of these Apollo 12, 14, 15, and 16 recordings indicated a core with a radius of approximately 200-300 km, later refined to ~300-400 km through improved modeling of reflected waves. As of November 2025, further analysis confirmed a solid inner core radius of approximately 258 km within a liquid outer core of about 362 km radius. On Mars, the core was first inferred in the late from estimates, but definitive detection came from the lander's seismic and magnetotelluric observations between 2018 and 2022, which captured core-reflected phases like PcP and . These data revealed a core consisting of a solid inner core with a radius of ~610 km surrounded by a liquid outer core extending to a total radius of ~1,830 km, larger than prior geophysical models predicted, highlighting the value of single-station for remote bodies. Venus's core models originated from the Magellan spacecraft's mapping in the , which measured tidal Love number k₂ ≈ 0.295 from Doppler tracking data, indicating a liquid outer core capable of deforming under tidal forces but lacking an active to generate a global . This inference relied on combining anomalies with rotational dynamics, as no seismic data exist for . Mercury's disproportionately large core, comprising ~85% of the planet's radius, was confirmed by the mission's topographic and measurements from 2008 to 2015, which showed high-density anomalies and a low consistent with a massive iron-rich interior partially molten at the boundary. Unlike seismic methods, these orbital data exploited Mercury's weak tidal response to constrain core size without direct wave sampling. Early indications of cores in gas giants arose from and 11 flybys in the and and 2 encounters in the late and , which detected strong magnetic fields implying action in conductive interiors, likely metallic hydrogen overlying rocky/icy cores. For , the Juno orbiter's gravity data since 2016 confirmed a diluted, fuzzy core extending approximately 30–50% of the planet's radius, blending gradually into the surrounding envelope rather than forming a sharp boundary.

Formation Mechanisms

Planetesimal accretion and differentiation

The formation of planetary cores begins during the accretion phase of planet formation, where dust grains in the aggregate into larger bodies. These particles, initially micrometer-sized, grow through collisions and sticking mechanisms influenced by turbulence and gas drag, eventually forming kilometer-scale typically 1–10 km in diameter. This formation occurs via processes such as the streaming instability, which concentrates solids in pressure bumps within the disk, enabling rapid over timescales of about 100–1000 years per local region. then collide and merge, growing into planetary embryos (moon- to mars-sized bodies) through runaway and oligarchic accretion, a process spanning 1–10 million years after the disk's formation. Core formation proceeds via differentiation, where denser metallic components separate from lighter under the influence of . During accretion, frequent impacts between planetesimals generate significant heating, raising temperatures above 2000 K in the interiors of larger embryos, sufficient to iron-nickel alloys while partially or fully surrounding . The molten metal, being denser (ρ ≈ 7–8 g/cm³ compared to silicate melt at ρ ≈ 3 g/cm³), sinks toward the center, driven by instabilities. This segregation releases energy, which further contributes to heating and sustains . A key mechanism for this metal descent is the Rayleigh-Taylor instability, occurring at the interface between the overlying lighter layer and the underlying denser molten metal pond formed by percolating droplets. In this model, small metal blebs ( <30°) percolate through a porous, partially molten matrix at rates of 1–10 m/year, accumulating into a layer that becomes unstable when its thickness exceeds ~1 km, triggering large-scale sinking plumes on timescales of hours to days. Experimental validations under high-pressure conditions confirm that this instability operates effectively at mantle pressures up to 10 GPa, facilitating efficient core assembly without requiring complete mantle melting. An emerging alternative mechanism, proposed in 2025, suggests that in some cases, particularly for sulfur-rich environments, molten sulfide rather than metal may percolate through solid rock to initiate core formation, lowering the melting point and enabling differentiation without extensive silicate melting. This model, supported by 3D imaging and geochemical analysis of meteorites, implies earlier core formation timelines and could revise radiogenic isotope dating, such as Hf-W, for planets like Mars. Core formation timescales vary but generally complete within 10–100 million years following the dispersal of the protoplanetary gas disk, constrained by short-lived radiometric chronometers like the Hf-W system, which records metal-silicate equilibration during early accretion. The energy released during differentiation arises primarily from the gravitational potential difference between undifferentiated and separated states. For a simplified uniform sphere, the total gravitational potential energy is E=−35GM2RE = -\frac{3}{5} \frac{G M^2}{R}, so the release upon core-mantle separation approximates ΔE≈35G(Mc2Rc−Mm2Rm)\Delta E \approx \frac{3}{5} G \left( \frac{M_c^2}{R_c} - \frac{M_m^2}{R_m} \right), where subscripts denote core (c) and mantle (m) masses and radii; this can contribute up to 10–20% of a planet's initial thermal budget. This process differs across planetary types: in terrestrial planets, rapid accretion and high impact energies enable near-complete core formation concurrently with growth, often finishing by ~10 million years. In contrast, gas giant cores form more slowly via planetesimal accretion to ~10 Earth masses before runaway gas capture, extending differentiation over tens of millions of years due to the insulating envelope suppressing heat loss.

Giant impacts and core merging

Giant impacts, occurring primarily in the final 10-100 million years of terrestrial planet formation, play a crucial role in determining the ultimate size and composition of planetary cores by merging the iron-rich cores of colliding bodies. These collisions involve protoplanets with masses comparable to , adding 1-2% to the target's total mass while facilitating the blending of metallic cores. For instance, in the Earth-Moon system, the impact of —a Mars-sized body—resulted in the rapid merger of Theia's core with proto-'s, contributing to a hybrid core structure enriched in Theia's material. This process is supported by dynamical models showing that the cores of miscible iron alloys blend efficiently upon collision, often as a sinking metallic diapir through the mantle, completing in less than 1 million years for impactors larger than 100 km in radius. Recent analyses of iron meteorites (October 2025) indicate protracted core formation in outer Solar System planetesimals, involving multi-stage differentiation disrupted by early high-energy impacts around 1–2.5 million years after calcium-aluminum-rich inclusions (CAIs). These events extracted sulfur-rich protocores initially, followed by sulfur-poor core segregation, shaping compositions in carbonaceous chondrite-related bodies and highlighting impact disruptions in smaller progenitors. The dynamics of core merging during these events typically involve complete thermal mixing for large impacts, enhancing the target's core heat budget and potentially driving transient dynamos, though incomplete chemical equilibration can occur if the diapir remains stratified. High-resolution smoothed particle hydrodynamics simulations demonstrate that planetary cores survive such collisions intact, even for impactor-to-target mass ratios up to 0.1 (or 10:1 inversely), with the impactor's core sinking and integrating without significant disruption. In some cases, however, the merging process introduces thermal stratification that temporarily inhibits convection in the core. These outcomes depend on impact parameters like velocity (6-10 km/s) and angle (30-60°), but the core's integrity is preserved across a wide range of giant impact scenarios. Planetary migration models, such as the Grand Tack hypothesis, further illustrate how giant impacts shape inner Solar System cores by altering the impactor flux. In this scenario, Jupiter's inward-then-outward migration scatters planetesimals, reducing material available for inner planet growth and promoting late-stage collisions that refine core compositions. Evidence from hafnium-tungsten (Hf-W) isotope dating corroborates this late accretion phase, indicating that Earth received an additional 0.5-1% of its mass after primary core formation, primarily through such impacts, as inferred from the mantle's tungsten isotopic anomaly. Specific examples highlight the diverse effects on core properties. The Earth-Moon-forming impact not only merged cores but also homogenized isotopic signatures, with Theia's contribution explaining similarities in Earth and lunar tungsten isotopes. On Mars, the Borealis basin-forming giant impact is proposed to have induced core asymmetry through uneven heating and iron addition, generating a deep thermal anomaly and potentially contributing to hemispheric differences in crustal thickness and mantle evolution. These cases underscore how giant impacts finalize core merging, influencing long-term planetary differentiation and dynamics.

Chemical Composition

Determination methods

The determination of planetary core compositions relies on a suite of indirect geophysical, geochemical, and computational techniques, as direct sampling is impossible for most bodies. These methods infer core properties from surface observations, laboratory simulations, and theoretical models, providing constraints on elements like iron, nickel, and light alloys in the core. Seismic methods are primary for Earth, analyzing wave propagation through the interior to derive core-mantle boundary depth, density, and elastic properties. By measuring P- and S-wave travel times, attenuation (which indicates material damping), and anisotropy (variations in wave speed by direction), researchers model core wave speeds and composition; for instance, the Preliminary Reference Earth Model (PREM) integrates global seismological data to specify radial variations in density and velocities, revealing an outer core with shear-wave velocity near 0 km/s due to its liquid state. For other planets, such as the Moon or Mars, limited seismic data from lander missions (e.g., Apollo seismometers or InSight) provide analogous constraints on core size and state, though with lower resolution. Gravimetric and magnetic data offer complementary insights into core mass distribution and dynamo activity. Spacecraft measurements of gravitational harmonics, such as the J2 oblateness coefficient, combined with factors (typically 0.3–0.4 for rocky planets), constrain core radius and density; for example, Mercury's low suggests a large, iron-rich core comprising over 70% of its radius. Paleomagnetic records from crustal rocks preserve ancient field directions and intensities, indicating core convection and composition via models; Martian meteorites, for instance, show evidence of a past reversing consistent with a liquid iron core. Geochemical proxies examine siderophile (iron-loving) element abundances in mantle-derived rocks to infer core sequestration. Partitioning coefficients for elements like nickel (Ni), cobalt (Co), and gold (Au) between metal and silicate phases during differentiation reveal core-mantle separation efficiencies; Earth's mantle depletions in these elements (e.g., Ni at ~0.2× chondritic) imply ~90% of planetary Ni resides in the core under equilibrium conditions at high pressures. These ratios, calibrated against meteoritic compositions, extend to other bodies like Mars, where mantle siderophile patterns suggest core formation at lower pressures than Earth. High-pressure experiments replicate core conditions using diamond anvil cells (DACs) to measure phase stability and partitioning. Laser-heated DACs achieve pressures exceeding 100 GPa and temperatures of 3000–6000 K, simulating core-mantle boundaries; these tests show that light elements like or alloy with iron-nickel at these extremes, lowering core to match seismic profiles. Numerical modeling, particularly quantum mechanical calculations, computes equations of state for candidate core materials. simulations of Fe-Ni alloys at core pressures (up to 360 GPa) and temperatures predict thermodynamic stability, elasticity, and sound velocities; for example, such models indicate that 5–10 wt% stabilizes the hexagonal close-packed phase in . These computations benchmark experimental data and extend inferences to extrasolar planets lacking direct observations. Despite these advances, determining core compositions remains indirect, especially for non-Earth bodies, where methods rely on Earth analogs or sparse mission data, introducing uncertainties in scaling pressure-temperature effects and light element roles.

Earth's core composition

Earth's core is primarily composed of iron, with making up approximately 5-10% of its mass, and the remainder consisting of lighter elements that account for the observed deficit relative to pure iron. Seismological models indicate that the outer core has a of about 9.9 to 12.2 g/cm³, while the inner core ranges from 12.8 to 13.1 g/cm³, compared to pure iron's of roughly 13.0 to 13.5 g/cm³ under core conditions. This deficit, estimated at 4-10% by volume or 6-10% by weight, necessitates the incorporation of light elements to reduce the overall while maintaining the core's seismic properties. The leading candidates for these light elements include (S), oxygen (O), (Si), (H), and carbon (C), collectively comprising 5-10% of the core's mass. Sulfur is inferred from geochemical models showing that up to 90% of Earth's total resides in the core, supported by sulfur isotope ratios in mantle-derived rocks that indicate depletion during core formation. Oxygen's presence is suggested by the mantle's relative depletion in siderophile elements, implying co-segregation with iron during differentiation. Silicon and are viable based on partitioning experiments under high-pressure conditions, while carbon emerges as a key component from recent atomic-scale simulations. Isotopic studies provide further constraints on core composition. Iron isotope fractionation between the core and mantle, with the core enriched in lighter isotopes (δ⁵⁶Fe ≈ -0.01 to -0.1‰ relative to the mantle), arises from equilibrium partitioning during metal-silicate separation under high pressure and temperature. The hafnium-tungsten (Hf-W) chronometer reveals that core formation occurred rapidly, within about 30 million years after solar system formation, influencing the distribution of light elements through high-temperature equilibration. Compositional differences exist between the inner and outer core. The outer core likely hosts higher concentrations of sulfur (up to 2-6 wt%), as evidenced by its lower density and seismic velocities consistent with liquid Fe-Ni-S alloys, whereas the inner core may be enriched in carbon. A 2025 study using density functional theory simulations demonstrates that approximately 3.8 wt% carbon is required to initiate inner core solidification by lowering the nucleation barrier, enabling crystallization under realistic thermal gradients without excessive supercooling. Silicon and sulfur, in contrast, hinder this process by stabilizing the liquid state. Recent 2025 analyses of isotopes in mantle rocks have uncovered traces of a primordial "," representing remnants of proto-Earth material that survived the planet's violent accretion and core formation. These ancient reservoirs, detected as isotopic imbalances (ε⁴⁰K deficits of approximately 65 ppm), suggest preserved pre-impact mantle heterogeneity with implications for early including core formation.

Evidence from meteorites

Iron meteorites are widely regarded as direct samples of the metallic cores from differentiated planetesimals that formed early in the Solar System. These meteorites, comprising primarily , represent fragments exposed through subsequent collisions and disruption of their parent bodies. Chemical groupings such as IIAB and IVA reflect distinct core compositions from these planetesimals, with nickel contents typically ranging from 5% to 20%, indicating variations in the metallic reservoirs during differentiation. Pallasites, another class of stony-iron meteorites, provide evidence of core-mantle interfaces in differentiated bodies, featuring intimate mixtures of crystals embedded in a Fe-Ni metal matrix. The metal phase in consists of approximately 90% Fe-Ni alloy, with the representing mantle material, suggesting these meteorites originated near the boundary between a metallic core and a mantle in their parent . This structure mimics the transitional zones expected in planetary interiors, offering insights into the physical separation of core and mantle materials during early differentiation. Patterns of siderophile elements—those with affinity for metal phases—further link meteorites to core formation processes, showing depletions in planetary mantles that correspond to enrichments in cores. For instance, highly siderophile elements like , , and are strongly partitioned into metallic cores, as evidenced by their elevated abundances in s compared to mantle-derived samples. These patterns align with the late addition of such elements to mantles via impacts, balancing the core enrichment observed in meteoritic analogs. The Widmanstätten patterns, characteristic interlocking structures of kamacite and lamellae in iron meteorites, record the slow cooling histories of these ancient cores. These patterns form through diffusion-controlled exsolution during subsolidus cooling, with bandwidth measurements indicating rates of approximately 1–10 K per million years for most groups, consistent with burial depths of several kilometers in parent bodies. Recent isotope studies from 2024 and 2025 have connected meteoritic compositions to heterogeneity in the , revealing nucleosynthetic variations that influenced the incorporation of light elements into cores. For example, analyses of and aluminum isotopes in meteorites suggest inherited presolar signatures that contributed to disk-scale differences, affecting light element budgets like carbon and in metallic cores. Similarly, iron-nickel isotope systematics in iron meteorites indicate distinct reservoirs in the inner Solar System, informing models of light element partitioning during core formation. Despite their value, meteorites serve as proxies primarily from small planetesimals rather than large , limiting direct comparisons to full-scale planetary cores. No provides an exact match to Earth's core, as these samples derive from disrupted asteroids with different sizes, accretion environments, and evolutionary paths.

Physical Structure and Dynamics

Internal structure models

Models of planetary cores typically describe a layered internal structure, consisting of a solid inner core surrounded by a liquid outer core, with potential transitional zones at the boundaries. For , the (PREM) delineates the inner core as a solid iron-rich sphere with a radius of approximately 1,220 km, where compressional wave velocity (Vp) jumps to about 11 km/s and shear wave velocity (Vs) to 3.5 km/s at the inner core boundary (ICB), marking the transition from the liquid outer core. Above the outer core lies , and at the core-mantle boundary (CMB), a possible F-layer—a thin, stably stratified region in the lowermost outer core—may exist, spanning 150–300 km thick and influencing propagation. Planetary models vary significantly by body type. Mercury features a disproportionately large core occupying about 85% of its radius, overlain by a thin silicate mantle roughly 400–600 km thick, consistent with Messenger mission gravity data indicating a partially molten outer core. In contrast, Jupiter's core is not sharply defined but exhibits a gradual dilution from a central rocky-metallic region of 7–25 Earth masses into surrounding metallic hydrogen layers, as inferred from Juno spacecraft measurements of gravitational harmonics. Recent 2025 simulations incorporating carbon as a alloying element in have refined models of seismic , showing how carbon enhances lattice preferred orientation and explains observed variations in wave speeds up to 3–4% faster along polar paths. Similarly, Juno data analyses in 2025 confirm Jupiter's blended core structure, with heavy elements dispersing outward over thousands of kilometers rather than forming a compact boundary, challenging prior giant impact formation scenarios. The radius of Earth's ICB is derived from contrasts in seismic wave speeds and Poisson's ratio (σ\sigma), where the liquid outer core exhibits σ≈0.5\sigma \approx 0.5 due to negligible shear modulus. The formula for Poisson's ratio is: σ=Vp2−2Vs22(Vp2−Vs2)\sigma = \frac{V_p^2 - 2 V_s^2}{2(V_p^2 - V_s^2)} At the ICB, the discontinuity in VsV_s from 0 km/s (outer core) to 3.5 km/s (inner core), combined with VpV_p increasing from ~8 km/s to 11 km/s, constrains the boundary radius through travel-time inversions in models like PREM. Uncertainties persist regarding transitional regions, such as potential mushy zones at the ICB where creates a of crystals and over 4–8% of the volume, or iron snowing in the outer core, where solid iron crystals form and settle from the upper outer core, forming piles up to 200 km thick. These features complicate precise boundary definitions and may influence core dynamics.

Heat sources and convection

The primary heat sources in planetary cores include residual heat from accretion during formation, possible minor radiogenic decay of in the outer core, though primary radiogenic heating occurs in from , , and potassium, released during inner core solidification, and compositional arising from the exclusion of lighter elements like or oxygen during crystallization. Accretional heat originates from the released as planetesimals collided and sank toward the center, providing an initial thermal endowment that persists as stored internal heat even after differentiation. Radiogenic contributions, while dominant in , are estimated to supply a smaller but sustained portion in the core, with potentially playing a key role due to its siderophile partitioning into metallic phases. Latent heat from inner core growth represents a major ongoing source, as the exothermic solidification of the liquid outer core releases approximately 3 terawatts (TW) in Earth's case, powering by adding to the fluid. Compositional buoyancy further drives motion, as the growing solid inner core rejects lighter elements into the outer core, creating contrasts that promote upward flow of less dense material. These sources collectively sustain a at the core-mantle boundary estimated at around 15 TW for , which is crucial for maintaining vigorous fluid motion. Convection in planetary liquid outer cores is primarily double-diffusive, involving coupled and composition gradients where thermal diffusion outpaces chemical diffusion, leading to and layered motions. This process is driven by adverse gradients: cooling at the core-mantle boundary promotes downward plumes, while inner core growth supplies both and lighter-element enrichment at the inner boundary, enhancing upward compositional plumes. The resulting convective power can be approximated by for buoyancy-driven work, P=ρgαΔTF,P = \rho g \alpha \Delta T F, where ρ\rho is fluid density, gg is gravitational acceleration, α\alpha is the thermal expansivity, ΔT\Delta T is the temperature contrast driving convection, and FF is the heat flux; this power balance links heat transport to the vigor of fluid motion and, indirectly, to dynamo activity in active cores. Over planetary lifetimes, core heat budgets evolve through secular cooling at rates of about 100 K per billion years (Gyr) for Earth, which drives inner core expansion at roughly 1 mm per year by lowering the freezing point of the liquid. This gradual solidification modulates convection by increasing compositional contrasts while reducing the liquid volume available for thermal stirring. In smaller bodies like Mars, faster cooling due to higher surface-to-volume ratios and lower radiogenic inventory leads to weaker convection, as the core loses heat more rapidly and transitions to conduction-dominated regimes sooner.

Magnetic dynamo generation

The magnetic dynamo in planetary cores generates self-sustaining through the inductive action of convective flows in electrically conducting fluids, primarily liquid iron or iron alloys. This process, known as , relies on the motion of the conducting fluid stretching and twisting existing lines, thereby amplifying the field until it balances dissipative effects. The theory posits that without such motion, magnetic fields would decay due to ohmic diffusion, but provides the necessary energy input to maintain the field over geological timescales. Key requirements for dynamo generation include vigorous convection exceeding a critical threshold, quantified by the Rayleigh number (Ra) surpassing approximately 10^4, which ensures instability in the fluid layer; high electrical conductivity of the core material, typically greater than 10^5 S/m for liquid iron to enable sufficient magnetic Reynolds numbers (Rm > 10-100) for field amplification; and rapid planetary rotation to impose the Coriolis force, which organizes flows into columnar structures aligned with the rotation axis and breaks symmetry for field generation. These conditions are met in planets with molten metallic cores where thermal or compositional buoyancy drives convection against stable stratification. In Earth's geodynamo, convective motions in the liquid outer core produce a strong zonal toroidal magnetic field of approximately 1-5 mT within the core, while the poloidal component emerges at the surface as a field of 0.3-0.6 gauss (30-60 μT). This configuration arises from the ω-effect, where shears poloidal fields into toroidal ones, and the α-effect, where helical regenerates poloidal fields from toroidal ones. The governing physics is captured by the induction equation, ∂B∂t=∇×(u×B)+η∇2B,\frac{\partial \mathbf{B}}{\partial t} = \nabla \times (\mathbf{u} \times \mathbf{B}) + \eta \nabla^2 \mathbf{B}, where B\mathbf{B} is the , u\mathbf{u} is the , and η\eta is the magnetic diffusivity; this equation demonstrates how (first term) can overcome (second term), circumventing anti-dynamo theorems like Cowling's, which prohibit axisymmetric fields in spherical geometries without additional asymmetries from . Dynamo-generated fields exhibit variations across planetary types: gas giants like and Saturn produce multipolar, non-dipolar fields due to deep-seated in metallic hydrogen layers with slower effective rotation relative to core size; in contrast, dynamos in Mars and are extinct, attributed to core solidification that halted by removing light elements and reducing fluidity. Numerical simulations have advanced understanding, with models incorporating light elements like carbon, which lower electrical conductivity compared to pure iron, thereby influencing onset and field morphology by altering Rm and requiring adjusted convective vigor for sustainability.

Stability and recent observations

The stability of planetary cores is maintained primarily through a delicate thermal balance that regulates the rate of inner core solidification, releasing and compositional buoyancy to sustain in the outer core and prevent complete freezing. This ensures the longevity of the geodynamo, as excessive cooling could lead to dynamo decay, as observed in ancient Martian cores that ceased generating a approximately 4 billion years ago due to insufficient convective vigor following planetary cooling. Instabilities, such as convective overshoot at the core-mantle boundary (CMB), can disrupt this equilibrium by allowing hot plumes to penetrate or cold material to sink, potentially altering and core dynamics over geological timescales. Recent seismic observations of , spanning the past two decades, reveal dynamic changes that challenge notions of its long-term solidity and uniformity. Studies from 2025 indicate that the inner core's surface is less rigid than previously assumed, exhibiting structural transformations likely driven by interactions with the fluid outer core, including variations in shear wave velocities suggestive of a superionic iron-carbon state. Over the last 20 years, seismic data have documented alterations in the inner core's shape, with evidence of hemispheric asymmetry where the displays distinct compared to the western, potentially linked to differential growth rates. Additionally, the inner core's rotation has slowed relative to since around 2010, with some analyses suggesting a temporary reversal or backtracking at rates of about 0.05–0.15° per year, part of a ~70-year oscillatory cycle. At the CMB, recent evidence points to instabilities influenced by subducted oceanic slabs, which accumulate as ultra-low velocity zones (ULVZs) and may form a "secret realm" of heterogeneous material affecting core-mantle coupling. These structures, detected globally through seismic imaging, retain water and other volatiles, potentially triggering reactions that produce iron hydrides and alter boundary dynamics, with implications for and . Such couplings could link core evolution to surface phenomena, including subtle variations in Earth's length-of-day, as inner core slowdowns propagate torque changes through the outer core to the mantle. For gas giants, 2025 data from NASA's Juno mission have refined models of Jupiter's core stability, showing a gradual blending of heavy elements into the surrounding envelope rather than a sharp boundary. This dilute, fuzzy core structure, extending over thousands of kilometers, suggests ongoing erosion and mixing driven by deep , challenging earlier compact-core paradigms and indicating long-term stability through dilution rather than solidification. These observations highlight broader trends in core evolution, where thermal imbalances can lead to weakening, as inferred from paleomagnetic records of extinct fields on smaller bodies like Mars.

Cores in the Solar System

Terrestrial planets and Moon

The cores of the terrestrial planets and the represent the densest, metal-rich interiors of these rocky bodies, primarily composed of iron and with varying amounts of lighter elements such as . These cores are inferred from geophysical measurements including fields, seismic , and magnetic observations, revealing a range of sizes, states (solid, liquid, or partially molten), and dynamo activities that reflect differences in planetary size, , and composition. Unlike the gaseous cores of outer planets, these are metallic and occupy significant fractions of the planetary volume, influencing surface and magnetic histories. Mercury possesses an unusually large core that extends to approximately 85% of the planet's radius, with a radius of about 2,074 km, making it the most core-dominated terrestrial body. This oversized core, comprising up to 80% of Mercury's mass, is surrounded by a thin mantle only about 300-400 km thick, and it likely contains significant to account for its of around 6.5-7 g/cm³ in models incorporating light elements. Data from NASA's mission indicate a solid inner core nearly as large as Earth's, surrounded by a liquid outer core, which supports the generation of Mercury's weak through dynamo action. Venus's core is estimated to occupy roughly 50% of the planet's radius, with a radius of about 3,000-3,200 km based on interior structure models constrained by gravity and rotation data. Likely fully liquid due to the planet's high internal temperatures and lack of a solid inner core, this iron-rich core does not currently sustain a , as evidenced by Venus's absence of a global despite its similar size to . Inferences from such as Pioneer Venus and suggest a core around 10-12 g/cm³, potentially alloyed with light elements, but without seismic confirmation. Earth's core comprises about 55% of the planet's , with an outer layer extending from approximately 2,900 km depth to an inner solid core boundary at around 5,150 km from the center, totaling a core of 3,480 km. This structure, determined from analyses and Earth's , features a outer core of molten iron-nickel alloy that drives a geodynamo, generating the planet's protective . The solid inner core, growing through at its boundary, has a of about 1,220 km and a exceeding 12 g/cm³. Mars harbors a core that occupies approximately 50% of its radius, with a total radius of 1,810 ± 150 km as measured by seismic data from NASA's mission. This core is primarily liquid, lacking a confirmed solid inner core until recent 2025 analyses suggesting a small solid inner core of about 600 km radius, and it does not currently support a dynamo, consistent with Mars's weak remnant crustal magnetism. The core's density is estimated at 5.5-6.5 g/cm³, enriched in sulfur to explain the planet's overall lower density compared to . The Moon's core is notably small, extending to only about 20% of the lunar , with a of roughly 330-380 km, comprising just 1.5-2% of the Moon's mass. Seismic data from lunar experiments and measurements indicate a partially molten structure, possibly with a solid inner core and liquid outer layer, but too weak to generate a present-day . Paleomagnetic analyses of Apollo mission samples reveal evidence of an ancient , implying a transient active from ~4.2 Ga to between ~3.5 and 1 Ga that has since ceased. Across the terrestrial planets and , core radius fractions generally decrease with increasing heliocentric distance—from Mercury's ~85% to the 's ~20%—reflecting accretion from increasingly volatile-rich materials in the inner Solar System. This trend is accompanied by progressive enrichment in and other light elements in the cores outward, which lowers core densities and influences dynamo cessation in smaller bodies like Mars and the .

Icy moons and gas giants

Among the icy moons of the outer Solar System, Ganymede stands out for its substantial core, which constitutes approximately 50% of the moon's radius and consists primarily of rock and iron, overlain by a rocky mantle and thick ice layers. This differentiated structure was inferred from gravity measurements and the detection of an intrinsic by the Galileo spacecraft's , indicating an active driven by in a partially liquid metallic core. The process likely involves fluid motion in the core, generating a field that interacts with Jupiter's to produce auroral emissions observed on Ganymede's surface. The gas giants exhibit cores that are more diluted and extended compared to those of terrestrial planets, blending gradually into surrounding layers of . For , interior models derived from NASA's Juno mission data reveal a fuzzy core with a mass of 10-20 masses, composed of rock and that mixes with the overlying hydrogen envelope, resulting in a gradual increase in rather than a sharp boundary. This "fuzzy" model, based on precise measurements, suggests the core formed through or giant impacts during the planet's early history, extending up to 50% of 's radius. possesses a similarly dilute core of 10-20 masses, enriched in rock and but diffused into the metallic hydrogen layer, as constrained by Cassini mission gravity data and ring-seismology observations that reveal wave patterns indicative of internal gradients. The planet's weaker compared to 's implies reduced in the core region, possibly due to compositional stratification that inhibits efficiency. Uranus and Neptune harbor icy-rocky cores estimated at 10-15 Earth masses, surrounded by mantles rich in water, ammonia, and methane ices that transition into hydrogen-helium envelopes. Their magnetic fields are notably tilted relative to the rotation axis, a feature potentially resulting from giant impacts that displaced core material and altered the dynamo geometry during the planets' formation. flyby observations provided the primary data on these fields, confirming non-axisymmetric multipolar structures generated in partially convective interiors, while Ulysses contributed polar magnetosphere insights for as a comparative baseline. A key trend across these bodies is the increasing fraction of ice in core compositions moving outward from to , reflecting formation in colder regions where volatiles condensed more readily onto rocky nuclei. These cores contribute to internal heat budgets through residual formation energy and differentiation, sustaining that powers dynamos and, in turn, auroral activity via interactions with particles. Cassini and Voyager missions have been instrumental in mapping these dynamics, highlighting how diluted cores enable prolonged thermal evolution in volatile-rich environments.

Remnant and transitional cores

Remnant cores represent the exposed or partially preserved metallic interiors of ancient that underwent differentiation but were later disrupted by collisions, providing direct samples of early planetary core formation. Iron meteorites, primarily composed of iron-nickel alloys, are widely regarded as fragments from the shattered cores of these protoplanetary bodies, which formed within the first few million years of the Solar System's history. These meteorites exhibit Widmanstätten patterns indicative of slow cooling in a core environment, supporting their origin as remnants from in the terrestrial planet-forming region. A prominent example is the , hypothesized to be a largely exposed core of a differentiated planetesimal stripped of its mantle and crust through violent impacts; NASA's Psyche mission, launched in 2023, is en route to orbit the in 2029 to test this hypothesis via in-situ measurements of its composition and magnetic properties. Transitional core-mantle fragments, such as and mesosiderites, offer insights into the interfaces between metallic cores and overlying layers in disrupted bodies. consist of crystals embedded in a nickel-iron matrix, interpreted as samples from the core-mantle boundary of a differentiated , where metal intruded into mantle material during or collision-induced mixing. Mesosiderites, in contrast, are breccias of metallic iron-nickel with basaltic and gabbroic s, suggesting they formed from high-velocity impacts that fragmented and reaccreted core and crustal materials, with minimal mantle involvement. These stony-iron meteorites highlight the dynamic processes of core-mantle differentiation followed by catastrophic disruption in the early Solar . In the Solar System, remnant cores persist in smaller bodies like asteroids, where seismic and gravitational data infer their presence. For instance, the Dawn mission's observations of Vesta revealed evidence consistent with partial differentiation and retention of a dense iron-rich interior despite subsequent impacts. Recent 2025 analyses of Vesta's from Doppler tracking and imaging refine this to a very small metallic core with a radius of ~45 km (possibly up to 150 km or absent), suggesting incomplete differentiation or a fragmented origin from a larger reshaped by early collisions, indicating limited convective activity in its early history. Fossilized remnants of ancient s, which powered core-generated , are preserved in magnetized lunar rocks and Martian crustal strips; these paleomagnetic signatures in lunar samples suggest a dynamo active from ~4.2 Ga to between ~3.5 and 1 Ga, while Martian stripes record a dynamo active over 4 billion years ago before cessation. Beyond intact remnants, studies emphasize the role of collisions in creating "patchwork" planetesimals, where repeated impacts disrupted early cores, scattering metallic fragments that later reaggregated into heterogeneous bodies; this links core formation to widespread early Solar System violence, as evidenced by isotopic and structural analyses of meteoritic materials. Observational techniques, including to detect metallic absorption features on surfaces and to probe subsurface densities, enable identification of these remnants; for example, surveys of M-class asteroids reveal high metal content consistent with exposed cores. Fossilized cores extend to extrasolar contexts through debris disks, where polluted atmospheres and circumstellar disks contain from disrupted planetary remnants accreted onto the stellar core. These disks, often composed of rocky and metallic debris from inner planets destroyed during post-main-sequence expansion, exhibit infrared excesses and metal lines (e.g., iron, ) indicative of core-like compositions vaporized and falling inward. Recent observations of s like LSPM J0207+3331 show ongoing accretion of such planetary debris, providing a window into the end states of core in other systems.

Extrasolar Planetary Cores

Theoretical models

Theoretical models for extrasolar planetary cores extend the core accretion paradigm, which posits that solid cores form from in protoplanetary disks before accreting gaseous envelopes if they reach a threshold. In this framework, core properties such as mass and composition are influenced by disk ZZ, orbital migration, and dynamical instabilities during formation and evolution. For instance, the expected core mass scales approximately as Mcore≈(Z0.02)3/2M⊕M_{\rm core} \approx \left( \frac{Z}{0.02} \right)^{3/2} M_\oplus, where Z=0.02Z = 0.02 represents solar , reflecting higher solid material availability in metal-rich disks that enables larger cores. This relation arises from models integrating accretion rates with disk properties, extended to migration scenarios where inward orbital drift concentrates solids and accelerates core growth. Chthonian planets represent a class of stripped rocky cores hypothesized to form when close-in or exoplanets lose their hydrogen-helium envelopes through photoevaporation driven by stellar radiation. Theoretical models predict such remnants with sizes of approximately 2-10 Earth radii and high densities around 5-8 g/cm³ due to exposed iron-rich surfaces after loss of lighter volatiles. These objects are expected in the , where intense irradiation prevents gas retention, leaving behind bare cores that may retain trace atmospheres of refractory elements. Diamond planets, or carbon-dominated worlds, are modeled as cores under extreme pressures where carbon compresses into layers overlying iron-nickel alloys, particularly in systems with carbon-to-oxygen ratios exceeding solar values. The concept arises in core accretion scenarios around carbon-enhanced , where high-pressure (beyond 100 GPa and temperatures above 2000 K) stabilizes phases comprising potentially 20-50% of the planetary radius. Early models suggested 55 Cancri e as a candidate with up to one-third by mass, but JWST observations as of 2024 indicate a rocky with a secondary atmosphere rich in CO or CO₂ from a ocean, challenging the . Hypothetical dense cores may form from debris of tidally disrupted planets around white dwarfs, where dynamical instabilities scatter planetary fragments, allowing reassembly of metal-polluted material into iron-dominated bodies. In such scenarios, remnants can have compositions mirroring the progenitor planet's heavy elements, including up to 90% iron by mass in low-mass objects, with radii under 0.5 radii but masses exceeding 2 M⊕M_\oplus, stable against further tidal forces at separations beyond 0.1 AU. Hot ice planets feature rocky cores enveloped by high-pressure ices in water-rich super-Earths, where interfaces separate the core from exotic phases like superionic or ice VII under pressures of 10-100 GPa and temperatures of 1000-3000 . Theoretical interior models indicate these ices form during core accretion in disks with high ice-to-rock ratios, migrating inward to retain hot mantles that conduct heat and possibly sustain activity, with core masses around 1-5 M⊕M_\oplus comprising silicates and metals. The supercritical boundary facilitates material transport, potentially eroding the core over gigayears while maintaining structural integrity. Dynamical instabilities further shape these cores, including erosion in hot Jupiters where convective mixing dissolves rocky interiors into envelopes, reducing core masses by 10-30% over 1-5 Gyr via double-diffusive processes at the core-mantle boundary. In close binary systems, core mergers during giant impacts or disk instabilities can form hybrid compositions, blending iron-rich and volatile-depleted materials to yield inflated envelopes in resulting hot Jupiters, with merger rates enhanced by orbital perturbations.

Inferred from observations

Indirect evidence for the presence and composition of extrasolar planetary cores is primarily derived from transit photometry, measurements, and atmospheric spectroscopy, which constrain planetary masses, radii, and compositions without direct imaging. These methods reveal density profiles that suggest rocky or metallic cores in various types, though interpretations often rely on mass-radius relationships calibrated against theoretical interior models. For instance, transit timing variations (TTVs) in multi-planet systems provide dynamical masses that, when combined with radii from transits, yield bulk densities indicative of core-dominated structures. In the Kepler-36 system, TTVs measured from precise transit timings of the inner rocky planet Kepler-36b and outer Kepler-36c imply masses of approximately 4.45 masses for b and 8.08 masses for c, leading to densities of about 7.5 g/cm³ and 3.7 g/cm³, respectively. These values suggest that Kepler-36b possesses a substantial rocky core comprising silicates and iron, while Kepler-36c likely has a similar core enveloped by a hydrogen-helium layer, highlighting how TTVs can differentiate core masses in close-orbiting systems. Radial velocity observations, which measure minimum planetary masses through stellar wobble, combined with transit-derived radii, enable calculations that infer core sizes via mass-radius relations. Super-Earths with densities exceeding 5 g/cm³, such as b ( ≈5.8 g/cm³ from a mass of 4.7 masses and radius of 1.6 radii), are interpreted as having iron-rich cores making up 50-70% of their mass, surrounded by mantles and thin volatile layers. This high places b firmly in the rocky regime, contrasting with lower-density mini-Neptunes and supporting core formation mechanisms similar to 's. Atmospheric spectroscopy from transmission observations during transits can reveal depleted volatiles or high-metallicity compositions, suggesting atmospheric stripping that exposes underlying cores. For , early spectra showed a flat transmission spectrum, interpreted as indicative of a high-mean-molecular-weight atmosphere possibly from a stripped hydrogen-helium , leaving a chthonian-like core of rock and with depleted light volatiles. Recent analyses reinforce this as a potential remnant core scenario, though hazy aerosols complicate unambiguous confirmation. As of 2025, (JWST) observations have provided new insights into hot s, revealing stripped s and exposed cores in several systems. For example, JWST near-infrared spectra of the hot sub-Neptune TOI-421 b indicate a low-mean-molecular-weight atmosphere with minimal , consistent with partial envelope stripping that exposes a rocky core, bridging the radius valley between super-Earths and sub-Neptunes. Similarly, ultra-hot Neptune LTT 9779 b shows spectral features suggesting core exposure through hydrodynamic escape, with enhanced in the remaining atmosphere. White dwarf pollution spectra offer another line of evidence for planetary core fragments, as these remnants accrete debris from disrupted planets, revealing core compositions via atmospheric metal lines. Recent 2025 spectroscopic surveys of polluted white dwarfs detect spikes in iron (Fe) and magnesium (Mg) abundances, consistent with accretion from rocky cores similar to those in super-Earths, with Fe/Mg ratios matching differentiated planetary interiors. For instance, the hydrogen-rich white dwarf LSPM J0207+3331 exhibits extreme pollution with 13 heavy elements, including elevated Fe and Mg, pointing to ongoing accretion from core fragments of former planets. A notable example of an inferred eroded core is TOI-849b, a Neptune-sized with a of approximately 5.2 g/cm³ derived from mass (≈39 masses) and transit radius (3.45 radii), suggesting it is the stripped remnant of a whose hydrogen-helium envelope was photoevaporated by its host star. This object occupies the "hot Neptune desert," providing a snapshot of core survival post-envelope loss. Inferring core properties remains challenging due to degeneracies in mass-radius-composition , where similar densities can arise from varying core sizes and envelope thicknesses, requiring multi-wavelength observations to resolve. No direct of cores has been achieved, limiting constraints to indirect methods that often yield bimodal solutions (rocky vs. gaseous).

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