Geochemistry

(1909) The term geochemistry was first used by the Swiss-German chemist Christian Friedrich Schönbein in 1838: "a comparative geochemistry ought to be launched, before geognosy can become geology, and before the mystery of the genesis of our planets and their inorganic matter may be revealed." However, for the rest of the century the more common term was "chemical geology", and there was little contact between geologists and chemists. The composition of meteorites was investigated and compared to terrestrial rocks as early as 1850. In 1901, Oliver C. Farrington hypothesised that, although there were differences, the relative abundances should still be the same. In the early 20th century, Max von Laue and William L. Bragg showed that X-ray scattering could be used to determine the structures of crystals. In the 1920s and 1930s, Victor Goldschmidt and associates at the University of Oslo applied these methods to many common minerals and formulated a set of rules for how elements are grouped. Goldschmidt published this work in the series Geochemische Verteilungsgesetze der Elemente [Geochemical Laws of the Distribution of Elements]. The research of Manfred Schidlowski from the 1960s to around the year 2002 was concerned with the biochemistry of the Early Earth with a focus on isotope-biogeochemistry and the evidence of the earliest life processes in Precambrian. ==Subfields==

Some subfields of geochemistry are: • Aqueous geochemistry studies the role of various elements in watersheds, including copper, sulfur, mercury, and how elemental fluxes are exchanged through atmospheric-terrestrial-aquatic interactions. • Biogeochemistry is the field of study focusing on the effect of life on the chemistry of the Earth. • Cosmochemistry includes the analysis of the distribution of elements and their isotopes in the cosmos. • Isotope geochemistry involves the determination of the relative and absolute concentrations of the elements and their isotopes in the Earth and on Earth's surface. • Organic geochemistry, the study of the role of processes and compounds that are derived from living or once-living organisms. • Photogeochemistry is the study of light-induced chemical reactions that occur or may occur among natural components of the Earth's surface. • Regional geochemistry includes applications to environmental, hydrological and mineral exploration studies. ==Chemical elements==

The building blocks of materials are the chemical elements. These can be identified by their atomic number Z, which is the number of protons in the nucleus. An element can have more than one value for N, the number of neutrons in the nucleus. The sum of these is the mass number, which is roughly equal to the atomic mass. Atoms with the same atomic number but different neutron numbers are called isotopes. A given isotope is identified by a letter for the element preceded by a superscript for the mass number. For example, two common isotopes of chlorine are 35Cl and 37Cl. There are about 1700 known combinations of Z and N, of which only about 260 are stable. However, most of the unstable isotopes do not occur in nature. In geochemistry, stable isotopes are used to trace chemical pathways and reactions, while radioactive isotopes are primarily used to date samples. ==Differentiation and mixing==

The chemical composition of the Earth and other bodies is determined by two opposing processes: differentiation and mixing. In the Earth's mantle, differentiation occurs at mid-ocean ridges through partial melting, with more refractory materials remaining at the base of the lithosphere while the remainder rises to form basalt. After an oceanic plate descends into the mantle, convection eventually mixes the two parts together. Erosion differentiates granite, separating it into clay on the ocean floor, sandstone on the edge of the continent, and dissolved minerals in ocean waters. Metamorphism and anatexis (partial melting of crustal rocks) can mix these elements together again. In the ocean, biological organisms can cause chemical differentiation, while dissolution of the organisms and their wastes can mix the materials again. This can occur when the solid is heterogeneous or a solid solution, and part of the melt is separated from the solid. The process is known as equilibrium or batch melting if the solid and melt remain in equilibrium until the moment that the melt is removed, and fractional or Rayleigh melting if it is removed continuously. Isotopic fractionation can have mass-dependent and mass-independent forms. Molecules with heavier isotopes have lower ground state energies and are therefore more stable. As a result, chemical reactions show a small isotope dependence, with heavier isotopes preferring species or compounds with a higher oxidation state; and in phase changes, heavier isotopes tend to concentrate in the heavier phases. Mass-dependent fractionation is largest in light elements because the difference in masses is a larger fraction of the total mass. Ratios between isotopes are generally compared to a standard. For example, sulfur has four stable isotopes, of which the two most common are 32S and 34S. The ratio of their concentrations, , is reported as :\delta{}^{34}\mathrm{S} = 1000\left(\frac{R}{R_\mathrm{s}}-1\right), where is the same ratio for a standard. Because the differences are small, the ratio is multiplied by 1000 to make it parts per thousand (referred to as parts per mil). This is represented by the symbol . Equilibrium Equilibrium fractionation occurs between chemicals or phases that are in equilibrium with each other. In equilibrium fractionation between phases, heavier phases prefer the heavier isotopes. For two phases A and B, the effect can be represented by the factor : a_\mathrm{A-B} = \frac{R_\mathrm{A}}{R_\mathrm{B}}. In the liquid-vapor phase transition for water, at 20 degrees Celsius is 1.0098 for 18O and 1.084 for 2H. In general, fractionation is greater at lower temperatures. At 0 °C, the factors are 1.0117 and 1.111. Kinetic When there is no equilibrium between phases or chemical compounds, kinetic fractionation can occur. For example, at interfaces between liquid water and air, the forward reaction is enhanced if the humidity of the air is less than 100% or the water vapor is moved by a wind. Kinetic fractionation generally is enhanced compared to equilibrium fractionation and depends on factors such as reaction rate, reaction pathway and bond energy. Since lighter isotopes generally have weaker bonds, they tend to react faster and enrich the reaction products. Biological fractionation is a form of kinetic fractionation since reactions tend to be in one direction. Biological organisms prefer lighter isotopes because there is a lower energy cost in breaking energy bonds. In addition to the previously mentioned factors, the environment and species of the organism can have a large effect on the fractionation. == Cycles ==

Through a variety of physical and chemical processes, chemical elements change in concentration and move around in what are called geochemical cycles. An understanding of these changes requires both detailed observation and theoretical models. Each chemical compound, element or isotope has a concentration that is a function of position and time, but it is impractical to model the full variability. Instead, in an approach borrowed from chemical engineering, In a type of model called a box model, a reservoir is represented by a box with inputs and outputs. Geochemical models generally involve feedback. In the simplest case of a linear cycle, either the input or the output from a reservoir is proportional to the concentration. For example, salt is removed from the ocean by formation of evaporites, and given a constant rate of evaporation in evaporite basins, the rate of removal of salt should be proportional to its concentration. For a given component , if the input to a reservoir is a constant and the output is for some constant , then the mass balance equation is {{NumBlk2|:|\frac{d C}{d t} = a - kC.|1}} This expresses the fact that any change in mass must be balanced by changes in the input or output. On a time scale of , the system approaches a steady state in which . The residence time is defined as :\tau_\mathrm{res} = C^\text{steady}/I = C^\text{steady}/O, where and are the input and output rates. In the above example, the steady-state input and output rates are both equal to , so . If the input and output rates are nonlinear functions of , they may still be closely balanced over time scales much greater than the residence time; otherwise, there will be large fluctuations in . In that case, the system is always close to a steady-state and the lowest order expansion of the mass balance equation will lead to a linear equation like Equation (). In most systems, one or both of the input and output depend on , resulting in feedback that tends to maintain the steady-state. If an external forcing perturbs the system, it will return to the steady-state on a time scale of . ==Abundance of elements==

Solar System The composition of the Solar System is similar to that of many other stars, and aside from small anomalies it can be assumed to have formed from a solar nebula that had a uniform composition, and the composition of the Sun's photosphere is similar to that of the rest of the Solar System. The composition of the photosphere is determined by fitting the absorption lines in its spectrum to models of the Sun's atmosphere. By far the largest two elements by fraction of total mass are hydrogen (74.9%) and helium (23.8%), with all the remaining elements contributing just 1.3%. There is a general trend of exponential decrease in abundance with increasing atomic number, although elements with even atomic number are more common than their odd-numbered neighbors (the Oddo–Harkins rule). Compared to the overall trend, lithium, boron and beryllium are depleted and iron is anomalously enriched. The pattern of elemental abundance is mainly due to two factors. The hydrogen, helium, and some of the lithium were formed in about 20 minutes after the Big Bang, while the rest were created in the interiors of stars. Most of our direct information on the composition of the giant planets is from spectroscopy. Since the 1930s, Jupiter was known to contain hydrogen, methane and ammonium. In the 1960s, interferometry greatly increased the resolution and sensitivity of spectral analysis, allowing the identification of a much greater collection of molecules including ethane, acetylene, water and carbon monoxide. However, Earth-based spectroscopy becomes increasingly difficult with more remote planets, since the reflected light of the Sun is much dimmer; and spectroscopic analysis of light from the planets can only be used to detect vibrations of molecules, which are in the infrared frequency range. This constrains the abundances of the elements H, C and N. and the final mission of the Cassini probe in 2017 was to enter the atmosphere of Saturn. In the atmosphere of Jupiter, He was found to be depleted by a factor of 2 compared to solar composition and Ne by a factor of 10, a surprising result since the other noble gases and the elements C, N and S were enhanced by factors of 2 to 4 (oxygen was also depleted but this was attributed to the unusually dry region that Galileo sampled). Outside the core, Jupiter has a mantle of liquid metallic hydrogen and an atmosphere of molecular hydrogen and helium. Metallic hydrogen does not mix well with helium, and in Saturn, it may form a separate layer below the metallic hydrogen. Direct information on Mars, Venus and Mercury largely comes from spacecraft missions. Using gamma-ray spectrometers, the composition of the crust of Mars has been measured by the Mars Odyssey orbiter, the crust of Venus by some of the Venera missions to Venus, Additional information on Mars comes from meteorites that have landed on Earth (the Shergottites, Nakhlites, and Chassignites, collectively known as SNC meteorites). Abundances are also constrained by the masses of the planets, while the internal distribution of elements is constrained by their moments of inertia. The compositions of the planets and the Moon are chondritic, meaning that within each group the ratios between elements are the same as in carbonaceous chondrites. The estimates of planetary compositions depend on the model used. In the equilibrium condensation model, each planet was formed from a feeding zone in which the compositions of solids were determined by the temperature in that zone. Thus, Mercury formed at 1400 K, where iron remained in a pure metallic form and there was little magnesium or silicon in solid form; Venus at 900 K, so all the magnesium and silicon condensed; Earth at 600 K, so it contains FeS and silicates; and Mars at 450 K, so FeO was incorporated into magnesium silicates. The greatest problem with this theory is that volatiles would not condense, so the planets would have no atmospheres and Earth no atmosphere. In chondritic mixing models, the compositions of chondrites are used to estimate planetary compositions. For example, one model mixes two components, one with the composition of C1 chondrites and one with just the refractory components of C1 chondrites. In another model, the abundances of the five fractionation groups are estimated using an index element for each group. For the most refractory group, uranium is used; iron for the second; the ratios of potassium and thallium to uranium for the next two; and the molar ratio FeO/(FeO+MgO) for the last. Using thermal and seismic models along with heat flow and density, Fe can be constrained to within 10 percent on Earth, Venus, and Mercury. U can be constrained within about 30% on Earth, but its abundance on other planets is based on "educated guesses". One difficulty with this model is that there may be significant errors in its prediction of volatile abundances because some volatiles are only partially condensed. ==Earth's crust==

The more common rock constituents are nearly all oxides; chlorides, sulfides and fluorides are the only important exceptions to this and their total amount in any rock is usually much less than 1%. By 1911, F. W. Clarke had calculated that a little more than 47% of the Earth's crust consists of oxygen. It occurs principally in combination as oxides, of which the chief are silica, alumina, iron oxides, and various carbonates (calcium carbonate, magnesium carbonate, sodium carbonate, and potassium carbonate). The silica functions principally as an acid, forming silicates, and all the commonest minerals of igneous rocks are of this nature. From a computation based on 1672 analyses of numerous kinds of rocks Clarke arrived at the following as the average percentage composition of the Earth's crust: SiO2=59.71, Al2O3=15.41, Fe2O3=2.63, FeO=3.52, MgO=4.36, CaO=4.90, Na2O=3.55, K2O=2.80, H2O=1.52, TiO2=0.60, P2O5=0.22, (total 99.22%). All the other constituents occur only in very small quantities, usually much less than 1%. The other determining factor, namely the physical conditions attending consolidation, plays, on the whole, a smaller part, yet is by no means negligible. Certain minerals are practically confined to deep-seated intrusive rocks, e.g., microcline, muscovite, diallage. Leucite is very rare in plutonic masses; many minerals have special peculiarities in microscopic character according to whether they crystallized in-depth or near the surface, e.g., hypersthene, orthoclase, quartz. There are some curious instances of rocks having the same chemical composition, but consisting of entirely different minerals, e.g., the hornblendite of Gran, in Norway, which contains only hornblende, has the same composition as some of the camptonites of the same locality that contain feldspar and hornblende of a different variety. In this connection, we may repeat what has been said above about the corrosion of porphyritic minerals in igneous rocks. In rhyolites and trachytes, early crystals of hornblende and biotite may be found in great numbers partially converted into augite and magnetite. Hornblende and biotite were stable under the pressures and other conditions below the surface, but unstable at higher levels. In the ground-mass of these rocks, augite is almost universally present. But the plutonic representatives of the same magma, granite, and syenite contain biotite and hornblende far more commonly than augite. ==Trace metals in the ocean==

Trace metals readily form complexes with major ions in the ocean, including hydroxide, carbonate, and chloride and their chemical speciation changes depending on whether the environment is oxidized or reduced. Benjamin (2002) defines complexes of metals with more than one type of ligand, other than water, as mixed-ligand-complexes. In some cases, a ligand contains more than one donor atom, forming very strong complexes, also called chelates (the ligand is the chelator). One of the most common chelators is EDTA (ethylenediaminetetraacetic acid), which can replace six molecules of water and form strong bonds with metals that have a plus two charge. With stronger complexation, lower activity of the free metal ion is observed. One consequence of the lower reactivity of complexed metals compared to the same concentration of free metal is that the chelation tends to stabilize metals in the aqueous solution instead of in solids. Within aquatic environments, cadmium(II) can either be in the form CdCl+(aq) in oxic waters or CdS(s) in a reduced environment. Thus, higher concentrations of Cd in marine sediments may indicate low redox potential conditions in the past. For copper(II), a prevalent form is CuCl+(aq) within oxic environments and CuS(s) and Cu2S within reduced environments. The reduced seawater environment leads to two possible oxidation states of copper, Cu(I) and Cu(II). Trace metals with nutrient-type distributions are strongly associated with the internal cycles of particulate organic matter, especially the assimilation by plankton. The lowest dissolved concentrations of these metals are at the surface of the ocean, where they are assimilated by plankton. As dissolution and decomposition occur at greater depths, concentrations of these trace metals increase. Residence times of these metals, such as zinc, are several thousand to one hundred thousand years. Finally, an example of a scavenged-type trace metal is aluminium, which has strong interactions with particles as well as a short residence time in the ocean. The residence times of scavenged-type trace metals are around 100 to 1000 years. The concentrations of these metals are highest around bottom sediments, hydrothermal vents, and rivers. For aluminium, atmospheric dust provides the greatest source of external inputs into the ocean. Iron and copper show hybrid distributions in the ocean. They are influenced by recycling and intense scavenging. Iron is a limiting nutrient in vast areas of the oceans and is found in high abundance along with manganese near hydrothermal vents. Here, many iron precipitates are found, mostly in the forms of iron sulfides and oxidized iron oxyhydroxide compounds. Concentrations of iron near hydrothermal vents can be up to one million times the concentrations found in the open ocean. Using electrochemical techniques, it is possible to show that bioactive trace metals (zinc, cobalt, cadmium, iron, and copper) are bound by organic ligands in surface seawater. These ligand complexes serve to lower the bioavailability of trace metals within the ocean. For example, copper, which may be toxic to open ocean phytoplankton and bacteria, can form organic complexes. The formation of these complexes reduces the concentrations of bioavailable inorganic complexes of copper that could be toxic to sea life at high concentrations. Unlike copper, zinc toxicity in marine phytoplankton is low and there is no advantage to increasing the organic binding of Zn2+. In high-nutrient, low-chlorophyll regions, iron is the limiting nutrient, with the dominant species being strong organic complexes of Fe(III). ==See also==