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Composition of Earth  

H. Palme

Early models of the composition of the Earth relied heavily on meteorites. In all these models Earth had different layers, each layer corresponded to a different type of meteorite or meteorite component. Later, more realistic models based on analyses of samples from Earth began with Ringwood’s pyrolite composition in the 1960s. Further improvement came with the analyses of rare MgO rich peridotites from a variety of occurrences all over the Earth, as xenoliths enclosed in melts from the upper mantle or as ultramafic massifs, tectonically emplaced on the Earth’s surface. Chemical systematics of these rocks allow the determination of the major element composition of the primitive upper mantle (PUM), the upper mantle after core formation and before extraction of basalts ultimately leading to the formation of the crust. Trace element analyses of upper mantle rocks confirmed their primitive nature. Geochemical and geophysical evidence argue for a bulk Earth mantle of uniform composition, identical to the PUM, also designated as “bulk silicate Earth” (BSE). The formation of a metal core was accompanied by the removal of siderophile and chalcophile elements into the core. Detailed modeling suggests that core formation was an ongoing process parallel to the accretion of Earth. The composition of the core is model dependent and thus uncertain and makes reliable estimates for siderophile and chalcophile element concentrations of bulk Earth difficult. Improved stable isotope analyses show isotopic similarities with noncarbonaceous chondrites (NCC), while the chemical composition of the mantle of the Earth indicates similarities with carbonaceous chondrites (CC). In detail, however, it can be shown that no single known meteorite group, nor any mixture of meteorite groups can match the chemical and isotopic composition of Earth. This conclusion is extremely important for any formation model of the Earth.

Article

Element Partitioning (Mineral-Melt, Metal-/Sulfide-Silicate) in Planetary Sciences  

Brandon Mahan

Element partitioning—at its most basic—is the distribution of an element of interest between two constituent phases as a function of some process. Major constituent elements generally affect the thermodynamic environment (chemical equilibrium) and therefore trace element partitioning is often considered, as trace elements are present in minute quantities and their equilibrium exchange reactions do not impart significant changes to the larger system. Trace elements are responsive to thermodynamic conditions, and thus they act as passive tracers of chemical reactions without appreciably influencing the bulk reactions themselves. In planetary sciences, the phase pairs typically considered are mineral-melt, metal-silicate, and sulfide-silicate, owing largely to the ubiquity of their coexistence in planetary materials across scales and context, from the micrometer-sized components of meteorites up to the size of planets (thousands of kilometers). It is common to speak of trace elements in terms of their tendency toward forming metallic, sulfidic, or oxide phases, and the terms “siderophile,” “chalcophile,” and “lithophile” (respectively) are used to define these tendencies under what is known as the Goldschmidt Classification scheme. The metric of an element’s tendency to concentrate into one phase relative to another is expressed as the ratio of its concentration (as a weight or molar fraction) in one phase over another, where convention dictates the reference frame as solid over liquid, and metal or sulfide over silicate; this mathematical term is the element’s partition coefficient, or distribution coefficient, between the two respective phases, D M Phase B Phase A (where M is the element of interest, most often reported as molar fraction), or simply D M . In general, trace elements obey Henry’s Law, where the element’s activity and concentration are linearly proportional. Practically speaking, this means that the element is sufficiently dilute in the system such that its atoms interact negligibly with one another compared to their interactions with major element phases, and thus the trace element’s partition coefficient in most settings is not appreciably affected by its concentration. The radius and charge of an element’s ionized species (its ionic radius and valence state)—in relation to either the major element ion for which it is substituting or the lattice site vacancy or interstitial space it is filling—generally determine the likelihood of trace element substitution or vacancy/interstitial fill (along with the net charge of the lattice space). The key energy consideration that underlies an element’s partitioning is the Gibbs free energy of reaction between the phases involved. Gibbs free energy is the change in internal energy associated with a chemical reaction (at a given temperature and pressure) that can be used to do work, and is denoted as Δ G rxn . Reactions with negative Δ G rxn values are spontaneous, and the magnitude of this negative value for a given phase, for example, a metal oxide, denotes the relative affinity of the metal toward forming oxides. That is to say, an element with a highly negative Δ G rxn for its oxide species at relevant pressure-temperature conditions will tend to be found in oxide and silicate minerals, that is, it will be lithophile (and vice versa for siderophile elements). Trace element partitioning systematics in mineral-melt and metal-/sulfide-silicate systems have boundless applications in planetary science. A growing collective understanding of the partition coefficients of elements has been built on decades of physical chemistry, deterministic theory, petrology, experimental petrology, and natural observations. Leveraging this immense intellectual, technical, and methodological foundation, modern trace element partitioning research is particularly aimed at constraining the evolution of plate tectonics on Earth (conditions and timing of onset), understanding the formation history of planetary materials such as chondrite meteorites and their constituents (e.g., chondrules), and de-convolving the multiply operating processes at play during the accretion and differentiation of Earth and other terrestrial planets.

Article

Formation, Composition, and Evolution of the Earth’s Core  

Francis Nimmo

The Earth’s core formed by multiple collisions with differentiated protoplanets. The Hf-W (hafnium-tungsten) isotopic system reveals that these collisions took place over a timescale of tens of megayears (Myr), in agreement with accretion simulations. The degree to which the iron and silicates re-equilibrated during each collision is uncertain and affects the apparent core age derived from tungsten isotopic measurements. Seismological data reveal that the core contains light elements in addition to Fe-Ni, and the outer core is more enriched in such elements than the inner core. Because O is excluded efficiently from solid iron, O is almost certainly an important constituent of the outer core. The identity of other elements is less certain, despite intensive measurements of their effects on seismic velocities, densities, and partitioning behavior at appropriate pressures and temperatures. Si and O are very likely present, with perhaps some S; C and H are less likely. Si and Mg may have exsolved over time, potentially helping to drive the geodynamo and producing a low-density layer at the top of the core. Radioactive elements (U, Th, K) are unlikely to be present in important concentrations. The cooling of the core is controlled by the mantle’s ability to extract heat. The geodynamo has existed for at least 3.5 gigayears (Gyr), placing a lower bound on the heat flow out of the core. Because the thermal conductivity of the core is uncertain by a factor of ~3, the lower bound on this heat flow is similarly uncertain. Once the inner core started to crystallize, additional sources of energy were available to power the geodynamo. Inner core crystallization likely started in the time range 0.5 to 2.0 Gyr Before Present (BP); paleomagnetic arguments have been advanced for inner core growth starting at several different epochs within this time range.

Article

Science and Exploration of the Moon: Overview  

Bradley L. Jolliff

Earth’s moon, hereafter referred to as “the Moon,” has been an object of intense study since before the time of the Apollo and Luna missions to the lunar surface and associated sample returns. As a differentiated rocky body and as Earth’s companion in the solar system, much study has been given to aspects such as the Moon’s surface characteristics, composition, interior, geologic history, origin, and what it records about the early history of the Earth-Moon system and the evolution of differentiated rocky bodies in the solar system. Much of the Apollo and post-Apollo knowledge came from surface geologic exploration, remote sensing, and extensive studies of the lunar samples. After a hiatus of nearly two decades following the end of Apollo and Luna missions, a new era of lunar exploration began with a series of orbital missions, including missions designed to prepare the way for longer duration human use and further exploration of the Moon. Participation in these missions has become international. The more recent missions have provided global context and have investigated composition, mineralogy, topography, gravity, tectonics, thermal evolution of the interior, thermal and radiation environments at the surface, exosphere composition and phenomena, and characteristics of the poles with their permanently shaded cold-trap environments. New samples were recognized as a class of achondrite meteorites, shown through geochemical and mineralogical similarities to have originated on the Moon. New sample-based studies with ever-improving analytical techniques and approaches have also led to significant discoveries such as the determination of volatile contents, including intrinsic H contents of lunar minerals and glasses. The Moon preserves a record of the impact history of the solar system, and new developments in timing of events, sample based and model based, are leading to a new reckoning of planetary chronology and the events that occurred in the early solar system. The new data provide the grist to test models of formation of the Moon and its early differentiation, and its thermal and volcanic evolution. Thought to have been born of a giant impact into early Earth, new data are providing key constraints on timing and process. The new data are also being used to test hypotheses and work out details such as for the magma ocean concept, the possible existence of an early magnetic field generated by a core dynamo, the effects of intense asteroidal and cometary bombardment during the first 500 million–600 million years, sequestration of volatile compounds at the poles, volcanism through time, including new information about the youngest volcanism on the Moon, and the formation and degradation processes of impact craters, so well preserved on the Moon. The Moon is a natural laboratory and cornerstone for understanding many processes operating in the space environment of the Earth and Moon, now and in the past, and of the geologic processes that have affected the planets through time. The Moon is a destination for further human exploration and activity, including use of valuable resources in space. It behooves humanity to learn as much about Earth’s nearest neighbor in space as possible.

Article

The Atmosphere of Uranus  

Leigh N. Fletcher

Uranus provides a unique laboratory to test current understanding of planetary atmospheres under extreme conditions. Multi-spectral observations from Voyager, ground-based observatories, and space telescopes have revealed a delicately banded atmosphere punctuated by storms, waves, and dark vortices, evolving slowly under the seasonal influence of Uranus’s extreme axial tilt. Condensables like methane and hydrogen sulphide play a crucial role in shaping circulation, clouds, and storm phenomena via latent heat release through condensation, strong equator-to-pole gradients suggestive of equatorial upwelling and polar subsidence, and the formation of stabilizing layers that may decouple different circulation and convective regimes as a function of depth. Phase transitions in the watery depths may also decouple Uranus’s atmosphere from motions within the interior. Weak vertical mixing and low atmospheric temperatures associated with Uranus’s negligible internal heat means that stratospheric methane photochemistry occurs in a unique high-pressure regime, decoupled from the influx of external oxygen. The low homopause also allows for the formation of an extensive ionosphere. Finally, the atmosphere provides a window on the bulk composition of Uranus—the ice-to-rock ratio, supersolar elemental and isotopic enrichments inferred from remote sensing, and future in situ measurements—providing key insights into its formation and subsequent migration. As a cold, hydrogen-dominated, intermediate-sized, slowly rotating, and chemically enriched world, Uranus could be the closest and best example of atmospheric processes on a class of worlds that may dominate the census of planets beyond our own solar system. Future missions to the Uranian system must carry a suite of instrumentation capable of advancing knowledge of the time-variable circulation, composition, meteorology, chemistry, and clouds on this enigmatic “ice giant.”