Some subfields of geochemistry are:[7]​.

The building blocks of materials are 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 numbers is the mass number, which is approximately equal to the atomic mass. Atoms with the same atomic number but different numbers of neutrons are called isotopes. A given isotope is identified by the element symbol preceded by a superscript mass number. For example, two common isotopes of chlorine are Cl and Cl. There are about 1700 known combinations of Z and N, of which only 260 are stable. However, most unstable isotopes do not occur in nature. In geochemistry, stable isotopes are used to trace chemical pathways and reactions, while isotopes are primarily used to date samples.[4]​.

The chemical behavior of an atom—its affinity for other elements and the type of bonds it forms—is determined by the arrangement of electrons in orbitals, particularly the outermost electrons (valence "Valence (chemistry)"). These arrangements are reflected in the position of the elements in the periodic table.[4]​ Depending on the position, the elements are divided into broad groups of alkali metals, alkaline earth metals, transition metals, semimetals (also known as metalloids), halogens, noble gases, lanthanides and actinides.[4]​.

Another useful classification scheme for geochemistry is the Goldschmidt classification, which places elements into four main groups:

Within each group, some elements are refractory, remaining stable at high temperatures, while others are volatile&action=edit&redlink=1 "Volatiity (chemistry) (not yet written)"), evaporating more easily, so heating can separate them.[1]​[4]​.

Contenido

La composición química de la Tierra y de otros cuerpos está determinada por dos procesos opuestos: la diferenciación y la mezcla. En el manto "Manto (geología)") de la Tierra, la diferenciación se produce en las dorsales mediooceánicas a través de la fusión parcial, con más materiales refractarios que permanecen en la base de la litosfera mientras que el resto se eleva para formar basalto. Después de que una placa oceánica descienda en el manto, la convección finalmente mezclará las dos partes. La erosión diferenciará el granito separándolo en arcilla, en el fondo del océano, arenisca, en el borde del continente, y minerales, disueltos en las aguas del océano. El metamorfismo y la anatexia (fusión parcial de las rocas de la corteza) pueden volver a mezclar esos elementos. En el océano, los organismos biológicos pueden causar una diferenciación química, mientras que la disolución de los organismos y sus desechos puede mezclar los materiales nuevamente.[1]​.

Division

An important source of differentiation is fractionation, an unequal distribution of elements and isotopes. This may be the result of chemical reactions, phase changes, kinetic effects or radioactivity.[1]​.

On the largest scale, planetary differentiation is a physical and chemical separation of a planet into chemically distinct regions. For example, terrestrial planets formed iron-rich cores and silicate-rich mantles and crusts.[14] In Earth's mantle, the main source of chemical differentiation is partial melting, particularly near mid-ocean ridges.[15] This can occur when the solid is heterogeneous or a solid solution, and some of the melt separates from the solid. The process is known as "equilibrium" or "batch melting" if the solid and melt remain in equilibrium until the melt is removed, and "fractional or Rayleigh melting" if it is removed continuously.[16]​.

Isotopic fractionation") can have mass-dependent and mass-independent forms. Molecules with heavier isotopes have lower zero-point energies and are therefore more stable. As a result, chemical reactions show little dependence on isotopes, 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. heavy.[17]​ Mass-dependent fractionation is greater in light elements because the difference in masses is a larger fraction of the total mass.[18]​.

The ratios between isotopes are generally compared to a standard. For example, sulfur has four stable isotopes, of which the two most common are S and S.[18]​ The ratio of their concentrations, , is reported as.

where is the same relationship for a standard. Because the differences are small, the proportion is multiplied by 1000 to make it parts per thousand. This is represented by the symbol .[17]​.

Fractionation equilibrium occurs between chemicals or phases that are in equilibrium with each other. In phase equilibrium fractionation, the heavier phases prefer the heavier isotopes. For two phases A and B, the effect can be represented by the factor.

In the liquid-vapor phase transition for water, at 20 °C it is for O and for H. In general, fractionation is greater at lower temperatures. At 0 °C, the factors are and .[17]​.

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 air humidity is less than 100% or the wind moves water vapor. weaker, they tend to react faster and enrich the reaction products.[17]​.

Biological fractionation is a form of kinetic fractionation, since the 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 factors mentioned above, the environment and species of the organism can have a large effect on fractionation.[17]​.

Chemical elements, through a variety of physical and chemical processes, change in concentration and move in what are known as "geochemical cycles." Understanding 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 which is not practical for modeling complete variability. Instead, in an approach borrowed from chemical engineering, geochemists[1] average the concentration over regions of the Earth called "geochemical reservoirs." The choice of deposit depends on the problem; For example, the ocean can be a single reservoir or divided into multiple reservoirs.[19]​ In a type of model called a box model, a reservoir is represented by a box with inputs and outputs.[1][19]​.

Geochemical models generally involve feedback. In the simplest case of a linear cycle, the input or output of a tank is proportional to the concentration. For example, salt is removed from the ocean by the formation of evaporites, and given a constant rate of evaporation in evaporite basins, the rate of salt removal 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 material balance equation is.

This expresses the fact that any change in mass must be balanced by changes in input or output. On a time scale of , the system approaches a steady state in which . The residence time is defined as.

where and are the entry and exit rates. In the example above, the steady-state entry and exit rates are equal to , so .[19]​.

If the entry and exit rates are nonlinear functions of , they can still be very balanced on time scales much longer than the residence time; Otherwise, there would be large fluctuations in . In that case, the system is always close to a stable state and an expansion of lower order than that of the mass balance equation will lead to a linear equation such as equation (1). In most systems, one or both of the input and output depend on , resulting in feedback that tends to maintain the stable state. If an external forcing disturbs the system, it will return to the stable state on a time scale of .[19]​.

Solar system

The composition of the solar system is similar to that of many other stars and, apart from small anomalies, it can be assumed that it formed from a solar nebula that had a uniform composition, and that 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 electromagnetic spectrum to models of the Sun's atmosphere. By far the two largest elements by fraction of the total mass are hydrogen (74.9%) and helium (23.8%), with all remaining elements contributing only 1.3%. There is a general trend of exponentially decreasing abundance related to increasing atomic number, although elements with even atomic number are more common than their odd-numbered neighbors (the Oddo-Harkins rule), compared to the general trend, lithium, boron and beryllium are depleted and iron is anomalously enriched.[23]​.

The pattern of elemental abundance is mainly due to two factors: hydrogen, helium and part of the lithium were formed about 20 minutes after the Big Bang, while the rest was created inside the stars.[4]​.

Meteorites

Meteorites come in a variety of compositions, but chemical analysis can determine whether they were once in planetesimals that merged or differentiated.[21] Chondrites are undifferentiated and have round inclusions of minerals called chondrules. With ages of 4.56 billion years, they date to the early solar system. One particular type, the CI chondrite, has a composition that closely resembles that of the Sun's photosphere, except due to the depletion of some volatiles (H, He, C, N, O) and a group of elements (Li, B, Be) that are destroyed by nucleosynthesis in the Sun.[4]​[21]​ Due to this last group, CI chondrites are considered a better combination for the composition of the early solar system. Furthermore, chemical analysis of CI chondrites is more precise than that of the photosphere, so they are generally used as a source of chemical abundance, despite their rarity (only five have been recovered on Earth).[21]​.

giant planets

The planets of the solar system are divided into two groups: the four inner planets are the terrestrial planets (Mercury "Mercury (planet)"), Venus "Venus (planet)"), Earth and Mars "Mars (planet)")), with relatively small sizes and rocky surfaces. The four outer planets are the giant planets, which are dominated by hydrogen and helium and have lower average densities. These can be subdivided into the gas giants (Jupiter "Jupiter (planet)") and Saturn "Saturn (planet)") and the ice giants (Uranus "Uranus (planet)") and Neptune "Neptune (planet)")) which have large ice cores.[24]​.

Most of our direct information about the composition of giant planets comes from spectroscopy. Since the 1930s, Jupiter was known to contain hydrogen, methane and ammonia. In the 1960s, interferometry greatly increased the resolution and sensitivity of spectral analysis, allowing the identification of a much larger collection of molecules including ethane, acetylene, water, and carbon monoxide.[25] However, Earth-based spectroscopy becomes increasingly difficult with more remote planets, as the reflected light from the Sun is much dimmer; and spectroscopic analysis of light from planets can only be used to detect vibrations of molecules, which are in the infrared frequency range. This restricts the abundance of the elements H, C and N.[25]​ Two other elements are detected: phosphorus in phosphine gas (PH) and germanium in germane gas (GeH).[25]​.

The helium atom has vibrations in the ultraviolet range, which is strongly absorbed by the atmospheres of the outer planets and by the Earth. Therefore, despite its abundance, helium was only detected when spacecraft were sent to the outer planets, and then only indirectly, through collision-induced absorption into hydrogen molecules.[25] More information about Jupiter was obtained with the Galileo probe when it was sent into the atmosphere in 1995; Saturn.[28]​ In Jupiter's atmosphere, He was found to be depleted by a factor of 2 compared to the solar composition and Ne by a factor of 10, a surprising result since the other noble gases and the elements C, N and S increased by factors of 2 to 4 (oxygen was also depleted, but this was attributed to the unusually dry region that Galileo sampled).[27]​.

Spectroscopic methods only penetrate the atmospheres of Jupiter and Saturn to depths where the pressure is approximately equal to 1 bar "Bar (unit)"), approximately the atmospheric pressure of the Earth at sea level.[25]​ The Galileo probe penetrated down to 22 bars.[27]​ This is a small fraction of the planet, which is expected to reach pressures of more than 40 Mbar. To constrain the composition in the interior, thermodynamic models are built using temperature information from infrared emission spectra and equations of state for probable compositions. High-pressure experiments predict that hydrogen will be a metallic liquid in the interior of Jupiter and Saturn, while in Uranus and Neptune it will remain in a molecular state. Estimates also depend on the models for the formation of the planets. Condensation of the presolar nebula would give rise to a gaseous planet with the same composition as the Sun, but planets could also have formed when a solid core captured nebulous gas.[25]​.

In current models, the four giant planets have rock and ice cores that are approximately the same size, but the proportion of hydrogen and helium decreases from about 300 Earth masses on Jupiter, to 75 on Saturn and only a few on Uranus and Neptune.[25] Thus, while the gas giants are composed mainly of hydrogen and helium, the ice giants are composed mainly of heavier elements (O, C, N, S), mainly in the form of water, methane and ammonia. The surfaces are cold enough for molecular hydrogen to be liquid, so much of each planet is probably an ocean of hydrogen lying on top of heavier compounds. 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 on Saturn it can form a separate layer beneath the metallic hydrogen.[25]​.

terrestrial planets

The terrestrial planets are thought to come from the same nebular material as the giant planets, but have lost most of the lighter elements and have different histories. Planets closer to the Sun might be expected to have a higher fraction of refractory elements, but if their later stages of formation involved collisions with large objects with orbits that sampled different parts of the Solar System, there might be little systematic dependence on position.[30]​.

Direct information about Mars, Venus and Mercury comes largely from spacecraft missions. Using gamma ray spectrometers), the composition of the crust of Mars has been measured by the Mars Odyssey orbiter,[31]​ the crust of Venus by some of the Venus Venera missions,[30]​ and the crust of Mercury by the MESSENGER spacecraft.[32]​ Additional information about Mars comes from meteorites that have landed on Earth (shergottites, nakhlites, and chassignites), collectively known as SNC meteorites).[33]​ Abundances are also limited by the masses of the planets, while the internal distribution of elements is constrained by their moments of inertia.[4]​.

The condensed planets of the solar nebula, and many of the details of their composition, are determined by fractionation as they cool. The condensing phases are divided into five groups. The first to condense are materials rich in refractory elements such as Ca and Al. These are followed by nickel and iron, then magnesium silicates. Below 700 kelvins (700 K), metals and silicates rich in volatiles form a fourth group, and the fifth FeO group includes magnesium silicates.[34]​ The compositions of the planets and the Moon are chondritic, which means that within each group the proportions between the elements are the same as in carbonaceous chondrites.[4]​.

Estimates of planetary compositions depend on the model used. In the equilibrium condensation model, each planet formed from a feeding zone in which the compositions of the solids were determined by the temperatures in those zones. Thus, Mercury formed at 1400 K, when iron remains 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 it incorporated FeO into magnesium silicates. The biggest problem with this theory is that the volatiles would not condense, so the planets would have no atmospheres and the Earth would have no atmosphere.[4]​.

In chondritic mixing models, chondrite compositions are used to estimate planetary compositions. For example, one model mixes two components, one with the C1 chondrite composition and another with only the refractory components of C1 chondrites.[4]​ In another model, the abundances of the five fractionation groups are estimated using an index element for each of the groups. For the most refractory group, uranium is used; iron for the second; the proportions of potassium and thallium to uranium for the following two; and the molar ratio FeO/(FeO+MgO) for the latter. Using thermal and seismic models along with heat flow and density, Fe can be restricted within 10% on Earth, Venus and Mercury. U can be restricted to about 30% on Earth, but its abundance on other planets is based on "educated guesses." A difficulty with this model is that there can be significant errors in its prediction of volatile abundances because some volatiles are only partially condensed.[34]​[4]​.

Los constituyentes más comunes de las rocas son casi todos óxidos; los cloruros, sulfuros y fluoruros son las únicas excepciones importantes a esto y su cantidad total en cualquier roca es generalmente mucho menor del 1%. F. W. Clarke ha calculado que un poco más del 47% de la corteza terrestre está compuesta principalmente de oxígeno. Se presenta principalmente en combinación como óxidos, de los que los principales son la sílice, la alúmina, los óxidos de hierro y varios carbonatos (carbonato de calcio, carbonato de magnesio, carbonato de sodio y carbonato de potasio). La sílice funciona principalmente como un ácido, formando silicatos, y todos los minerales más comunes de las rocas ígneas son de esta naturaleza. A partir de un cálculo basado en 1672 análisis de numerosos tipos de rocas, Clarke llegó a la siguiente composición porcentual promedio de la corteza terrestre:.

Todos los demás constituyentes aparecen solo en cantidades muy pequeñas, generalmente mucho menores del 1% .[35]​

Estos óxidos se combinan de una manera casual. Por ejemplo, la potasa (carbonato de potasio) y la sosa (carbonato de sodio) se combinan para producir feldespatos. En algunos casos, pueden tomar otras formas, como nefelina, leucita y moscovita, pero en la gran mayoría de los casos se encuentran como feldespato. El ácido fosfórico con cal (carbonato de calcio) forma la apatita. El dióxido de titanio con óxido ferroso da lugar a la ilmenita. Parte de la cal forma feldespato de cal. El carbonato de magnesio y los óxidos de hierro con sílice cristalizan en forma de olivino o enstatita, o con alúmina y cal forman el complejo de silicatos ferro-magnésicos, de los que los piroxenos, anfíboles y biotitas son los principales. Cualquier exceso de sílice por encima de lo que se requiere para neutralizar las bases se separará como cuarzo; el exceso de alúmina cristaliza como corindón. Estas deben considerarse solo como tendencias generales. Es posible, según el análisis de la roca, decir aproximadamente qué minerales contiene la roca, pero hay numerosas excepciones a cualquier regla.[35]​.

mineral constitution

Except in acidic or siliceous igneous rocks containing more than 66% silica, known as felsic rocks, quartz is not abundant in igneous rocks. Basic rocks (containing 20% ​​silica or less) rarely contain as much silicon, and are known as mafic rocks. If magnesium and iron are above average, while silica is low, olivine can be expected; When silica is present in greater quantity on ferro-magnesium minerals, other minerals such as augite, hornblende, enstatite or biotite occur instead of olivine. Unless potash is high and silica relatively low, leucite will not be present, since leucite is not produced with free quartz. Nepheline, likewise, is generally found in rocks with a lot of soda and comparatively little silica. At high alkalis, soda-containing pyroxenes and amphiboles may be present. The lower the percentage of silica and alkalis, the greater the prevalence of plagioclase feldspar as contracted to soda or potassium feldspar.[35]​.

The Earth's crust is composed of 90% silicate minerals and their abundance on Earth is as follows: plagioclase feldspar" (39%), alkali feldspar (12%), quartz (12%), pyroxene (11%), amphiboles (5%), micas (5%), clay minerals (5%); the remaining silicate minerals make up another 3% of the Earth's crust. Only 8% of the Earth is composed of minerals that They are not silicates, like carbonates"), oxides and sulfides.[36]​.

The other determining factor, namely the physical conditions accompanying consolidation, generally plays a smaller part, but is by no means negligible. Certain minerals are practically confined to deep intrusive rocks, for example, microcline, muscovite, diallage. Leucite is very rare in plutonic masses; Many minerals have special peculiarities in their microscopic character depending on whether they crystallized at depth or near the surface, for example, hypersthene, orthoclase, quartz. There are some curious examples of rocks having the same chemical composition, but consisting of completely different minerals, for example, the hornblende from Gran, in Norway, which contains only hornblende, has the same composition as some of the camptonites from the same locality which contain feldspar and hornblende of a different variety. In this sense, what has been said previously about the corrosion of porphyritic minerals in igneous rocks can be repeated. In rhyolites and trachytes, early hornblende and biotite crystals can be found in large quantities partially converted to augite and magnetite. Hornblende and biotite remained stable under pressures and other conditions below the surface, but are unstable at higher levels. In the soil mass of these rocks, augite is almost universally present. But the plutonic representatives of the same magma, granite and syenite contain biotite and hornblende much more commonly than augite.[35]​.

Felsic, intermediate and mafic igneous rocks

Those rocks that contain the greatest amount of silica, and when crystallized leave free quartz, form a group generally called "felsic" rocks. Those that again contain the least silica and the most magnesia and iron, so that quartz is absent, while olivine is usually abundant, form the "mafic" group. "Intermediate" rocks include those characterized by the general absence of both quartz and olivine. An important subdivision of these contains a very high percentage of alkalis, especially soda, and consequently has minerals such as nepheline and leucite that are not common in other rocks. It is often separated from the others as "alkaline" or "soda" rocks, and there is a corresponding series of mafic rocks. Finally, a small subgroup rich in olivine and lacking feldspar has been called “ultramafic” rocks. They have very low percentages of silica but a lot of iron and magnesia.

Except for the latter, practically all rocks contain feldspars or feldspathoid minerals. In acidic rocks, common feldspars are orthoclase, perthite, microcline, and oligoclase, all with lots of silica and alkalis. Labradorite, anorthite and bytownite prevail in mafic rocks, which are rich in lime and poor in silica, potash and soda. Augite is the most common ferromagnesian in mafic rocks, but biotite and hornblende are generally more common in felsic rocks.[35]​.

Rocks containing leucite or nepheline, either partially or completely replacing feldspars, are not included in this table. They are essentially of an intermediate or mafic character. Accordingly, they might be regarded as varieties of syenite, diorite, gabbro, etc., in which feldspathoids occur, and in fact there are many transitions between syenites of the ordinary type and nepheline—or leucite—syenite, and between gabbro or dolerite and mineralite. or sexy. But, since many minerals develop in these "alkaline" rocks that are not common elsewhere, it is convenient in a purely formal classification such as that described here, to treat the whole assemblage as a distinct series.[35]​.

This classification is essentially based on the mineralogical constitution of igneous rocks. Any chemical distinction between the different groups, although implicit, is relegated to a subordinate position. It is true that it is artificial, but it has grown with the growth of science and is still adopted as the basis on which still smaller subdivisions are erected. The subdivisions are in no way of equal value. Syenites, for example, and peridotites, are much less important than granites, diorites and gabbros. Furthermore, effusive andesites do not always correspond to plutonic diorites but partly also to gabbros. As different types of rock, considered aggregates of minerals, gradually pass over one another, transitional types are very common and are often important enough to be given special names. Quartz syenites and nordmarkites") can come between granite and syenite, tonalites and adamelites between granite and diorite, monzoaites between syenite and diorite, norites and hyperites") between diorite and gabbro, and so on.[35]​.

Trace metals readily form "Complex (chemistry)") complexes with major ions in the ocean, including hydroxide, carbonate, and chloride, and their changes in chemical speciation depend 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 water molecules and form strong bonds with metals that have a charge of more than two. A consequence of the lower reactivity of complexed metals compared to the same concentration of free metal is that chelation tends to stabilize the metals in aqueous solution rather than in solids.[38]​.

The concentrations of some trace metals—cadmium, copper, molybdenum, manganese, rhenium, uranium, and vanadium—in certain sediments record the redox history of the oceans. In aquatic environments, cadmium(II) can be in the form CdCl in oxic waters or CdS(s) in a reduced environment. Therefore, higher concentrations of Cd in marine sediments may indicate conditions of low redox potential in the past. For copper (II), a prevalent form is CuCl(aq) in oxic environments and CuS(s) and CuS in reduced environments. The reduced seawater environment leads to two possible oxidation states of copper, Cu(I) and Cu(II). Molybdenum is present as the oxidation state of Mo(VI) as MoO in oxic environments. Rhenium is present as the oxidation state Re (VII) as ReO under oxic conditions, but is reduced to Re (IV) which can form ReO or ReS Uranium is in the oxidation state VI in UUO(CO)(aq) and is found in the reduced form UO(s). Vanadium occurs in various forms in the oxidation state V (V) and HVO; This relative predominance of these species depends on pH.

In the water column of the ocean or deep lakes, the vertical profiles of dissolved trace metals are characterized by following the following distributions: conservative type (conservative–type), nutrient type (nutrient–type) or scavenged–type*. Across these three distributions, trace metals have different residence times and are used to different extents by planktonic microorganisms. Trace metals with conservative distributions have high concentrations relative to their biological use. An example of a trace metal with a conservative distribution is molybdenum. It has a residence time in the oceans of about 8 x 10 years and is generally present as the anion molybdate (MoO). Molybdenum interacts weakly with particles and shows a nearly uniform vertical profile. in the ocean. Relative to the abundance of molybdenum in the ocean, the amount required as a metal cofactor for enzymes in marine phytoplankton is negligible.[39]​.

Trace metals with nutrient-like distributions are strongly associated with the internal cycles of particulate organic matter, especially assimilation by plankton. The lowest dissolved concentrations of these metals are found at the ocean surface, where they are assimilated by plankton. As dissolution and decomposition occur at greater depths, the concentrations of these trace metals increase. The residence times of these metals, such as zinc, are several thousand to one hundred thousand years. Finally, an example of a swept-type trace metal is aluminum, which has strong interactions with particles and a short residence time in the ocean. Residence times for sweep-type trace metals are around 100 to 1000 years. Concentrations of these metals are highest around bottom sediments, hydrothermal vents, and rivers. For aluminum, atmospheric dust provides the largest source of external inputs to the ocean.[39]​.

Iron and copper show hybrid distributions in the ocean. They are influenced by recycling and intense cleaning. Iron is a limiting nutrient in vast areas of the oceans, and is found in great abundance along with manganese near hydrothermal vents. Here, many iron precipitates are found, mainly in the form of iron sulfides and oxidized iron oxyhydroxide compounds. Iron concentrations near hydrothermal vents can be up to a million times the concentrations found in the open ocean.[39]​.

Using electrochemical techniques, it is possible to show that bioactive trace metals (zinc, cobalt, cadmium, iron and copper) are bound by organic ligands on the surface of seawater. These ligand complexes serve to decrease the bioavailability of trace metals in the ocean. For example, copper, which can be toxic to phytoplankton and bacteria in the open ocean, can form organic complexes. The formation of these complexes reduces concentrations of bioavailable inorganic copper complexes that could be toxic to marine life at high concentrations. Unlike copper, the toxicity of zinc in marine phytoplankton is low and there is no advantage to increasing organic binding of Zn. In high-chlorophyll, low-nutrient regions, iron is the limiting nutrient, and the dominant species are strong organic Fe(III) complexes.[39]​.

Some subfields of geochemistry are:[7]​.

The building blocks of materials are 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 numbers is the mass number, which is approximately equal to the atomic mass. Atoms with the same atomic number but different numbers of neutrons are called isotopes. A given isotope is identified by the element symbol preceded by a superscript mass number. For example, two common isotopes of chlorine are Cl and Cl. There are about 1700 known combinations of Z and N, of which only 260 are stable. However, most unstable isotopes do not occur in nature. In geochemistry, stable isotopes are used to trace chemical pathways and reactions, while isotopes are primarily used to date samples.[4]​.

The chemical behavior of an atom—its affinity for other elements and the type of bonds it forms—is determined by the arrangement of electrons in orbitals, particularly the outermost electrons (valence "Valence (chemistry)"). These arrangements are reflected in the position of the elements in the periodic table.[4]​ Depending on the position, the elements are divided into broad groups of alkali metals, alkaline earth metals, transition metals, semimetals (also known as metalloids), halogens, noble gases, lanthanides and actinides.[4]​.

Another useful classification scheme for geochemistry is the Goldschmidt classification, which places elements into four main groups:

Within each group, some elements are refractory, remaining stable at high temperatures, while others are volatile&action=edit&redlink=1 "Volatiity (chemistry) (not yet written)"), evaporating more easily, so heating can separate them.[1]​[4]​.

Contenido

La composición química de la Tierra y de otros cuerpos está determinada por dos procesos opuestos: la diferenciación y la mezcla. En el manto "Manto (geología)") de la Tierra, la diferenciación se produce en las dorsales mediooceánicas a través de la fusión parcial, con más materiales refractarios que permanecen en la base de la litosfera mientras que el resto se eleva para formar basalto. Después de que una placa oceánica descienda en el manto, la convección finalmente mezclará las dos partes. La erosión diferenciará el granito separándolo en arcilla, en el fondo del océano, arenisca, en el borde del continente, y minerales, disueltos en las aguas del océano. El metamorfismo y la anatexia (fusión parcial de las rocas de la corteza) pueden volver a mezclar esos elementos. En el océano, los organismos biológicos pueden causar una diferenciación química, mientras que la disolución de los organismos y sus desechos puede mezclar los materiales nuevamente.[1]​.

Division

An important source of differentiation is fractionation, an unequal distribution of elements and isotopes. This may be the result of chemical reactions, phase changes, kinetic effects or radioactivity.[1]​.

On the largest scale, planetary differentiation is a physical and chemical separation of a planet into chemically distinct regions. For example, terrestrial planets formed iron-rich cores and silicate-rich mantles and crusts.[14] In Earth's mantle, the main source of chemical differentiation is partial melting, particularly near mid-ocean ridges.[15] This can occur when the solid is heterogeneous or a solid solution, and some of the melt separates from the solid. The process is known as "equilibrium" or "batch melting" if the solid and melt remain in equilibrium until the melt is removed, and "fractional or Rayleigh melting" if it is removed continuously.[16]​.

Isotopic fractionation") can have mass-dependent and mass-independent forms. Molecules with heavier isotopes have lower zero-point energies and are therefore more stable. As a result, chemical reactions show little dependence on isotopes, 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. heavy.[17]​ Mass-dependent fractionation is greater in light elements because the difference in masses is a larger fraction of the total mass.[18]​.

The ratios between isotopes are generally compared to a standard. For example, sulfur has four stable isotopes, of which the two most common are S and S.[18]​ The ratio of their concentrations, , is reported as.

where is the same relationship for a standard. Because the differences are small, the proportion is multiplied by 1000 to make it parts per thousand. This is represented by the symbol .[17]​.

Fractionation equilibrium occurs between chemicals or phases that are in equilibrium with each other. In phase equilibrium fractionation, the heavier phases prefer the heavier isotopes. For two phases A and B, the effect can be represented by the factor.

In the liquid-vapor phase transition for water, at 20 °C it is for O and for H. In general, fractionation is greater at lower temperatures. At 0 °C, the factors are and .[17]​.

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 air humidity is less than 100% or the wind moves water vapor. weaker, they tend to react faster and enrich the reaction products.[17]​.

Biological fractionation is a form of kinetic fractionation, since the 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 factors mentioned above, the environment and species of the organism can have a large effect on fractionation.[17]​.

Chemical elements, through a variety of physical and chemical processes, change in concentration and move in what are known as "geochemical cycles." Understanding 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 which is not practical for modeling complete variability. Instead, in an approach borrowed from chemical engineering, geochemists[1] average the concentration over regions of the Earth called "geochemical reservoirs." The choice of deposit depends on the problem; For example, the ocean can be a single reservoir or divided into multiple reservoirs.[19]​ In a type of model called a box model, a reservoir is represented by a box with inputs and outputs.[1][19]​.

Geochemical models generally involve feedback. In the simplest case of a linear cycle, the input or output of a tank is proportional to the concentration. For example, salt is removed from the ocean by the formation of evaporites, and given a constant rate of evaporation in evaporite basins, the rate of salt removal 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 material balance equation is.

This expresses the fact that any change in mass must be balanced by changes in input or output. On a time scale of , the system approaches a steady state in which . The residence time is defined as.

where and are the entry and exit rates. In the example above, the steady-state entry and exit rates are equal to , so .[19]​.

If the entry and exit rates are nonlinear functions of , they can still be very balanced on time scales much longer than the residence time; Otherwise, there would be large fluctuations in . In that case, the system is always close to a stable state and an expansion of lower order than that of the mass balance equation will lead to a linear equation such as equation (1). In most systems, one or both of the input and output depend on , resulting in feedback that tends to maintain the stable state. If an external forcing disturbs the system, it will return to the stable state on a time scale of .[19]​.

Solar system

The composition of the solar system is similar to that of many other stars and, apart from small anomalies, it can be assumed that it formed from a solar nebula that had a uniform composition, and that 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 electromagnetic spectrum to models of the Sun's atmosphere. By far the two largest elements by fraction of the total mass are hydrogen (74.9%) and helium (23.8%), with all remaining elements contributing only 1.3%. There is a general trend of exponentially decreasing abundance related to increasing atomic number, although elements with even atomic number are more common than their odd-numbered neighbors (the Oddo-Harkins rule), compared to the general trend, lithium, boron and beryllium are depleted and iron is anomalously enriched.[23]​.

The pattern of elemental abundance is mainly due to two factors: hydrogen, helium and part of the lithium were formed about 20 minutes after the Big Bang, while the rest was created inside the stars.[4]​.

Meteorites

Meteorites come in a variety of compositions, but chemical analysis can determine whether they were once in planetesimals that merged or differentiated.[21] Chondrites are undifferentiated and have round inclusions of minerals called chondrules. With ages of 4.56 billion years, they date to the early solar system. One particular type, the CI chondrite, has a composition that closely resembles that of the Sun's photosphere, except due to the depletion of some volatiles (H, He, C, N, O) and a group of elements (Li, B, Be) that are destroyed by nucleosynthesis in the Sun.[4]​[21]​ Due to this last group, CI chondrites are considered a better combination for the composition of the early solar system. Furthermore, chemical analysis of CI chondrites is more precise than that of the photosphere, so they are generally used as a source of chemical abundance, despite their rarity (only five have been recovered on Earth).[21]​.

giant planets

The planets of the solar system are divided into two groups: the four inner planets are the terrestrial planets (Mercury "Mercury (planet)"), Venus "Venus (planet)"), Earth and Mars "Mars (planet)")), with relatively small sizes and rocky surfaces. The four outer planets are the giant planets, which are dominated by hydrogen and helium and have lower average densities. These can be subdivided into the gas giants (Jupiter "Jupiter (planet)") and Saturn "Saturn (planet)") and the ice giants (Uranus "Uranus (planet)") and Neptune "Neptune (planet)")) which have large ice cores.[24]​.

Most of our direct information about the composition of giant planets comes from spectroscopy. Since the 1930s, Jupiter was known to contain hydrogen, methane and ammonia. In the 1960s, interferometry greatly increased the resolution and sensitivity of spectral analysis, allowing the identification of a much larger collection of molecules including ethane, acetylene, water, and carbon monoxide.[25] However, Earth-based spectroscopy becomes increasingly difficult with more remote planets, as the reflected light from the Sun is much dimmer; and spectroscopic analysis of light from planets can only be used to detect vibrations of molecules, which are in the infrared frequency range. This restricts the abundance of the elements H, C and N.[25]​ Two other elements are detected: phosphorus in phosphine gas (PH) and germanium in germane gas (GeH).[25]​.

The helium atom has vibrations in the ultraviolet range, which is strongly absorbed by the atmospheres of the outer planets and by the Earth. Therefore, despite its abundance, helium was only detected when spacecraft were sent to the outer planets, and then only indirectly, through collision-induced absorption into hydrogen molecules.[25] More information about Jupiter was obtained with the Galileo probe when it was sent into the atmosphere in 1995; Saturn.[28]​ In Jupiter's atmosphere, He was found to be depleted by a factor of 2 compared to the solar composition and Ne by a factor of 10, a surprising result since the other noble gases and the elements C, N and S increased by factors of 2 to 4 (oxygen was also depleted, but this was attributed to the unusually dry region that Galileo sampled).[27]​.

Spectroscopic methods only penetrate the atmospheres of Jupiter and Saturn to depths where the pressure is approximately equal to 1 bar "Bar (unit)"), approximately the atmospheric pressure of the Earth at sea level.[25]​ The Galileo probe penetrated down to 22 bars.[27]​ This is a small fraction of the planet, which is expected to reach pressures of more than 40 Mbar. To constrain the composition in the interior, thermodynamic models are built using temperature information from infrared emission spectra and equations of state for probable compositions. High-pressure experiments predict that hydrogen will be a metallic liquid in the interior of Jupiter and Saturn, while in Uranus and Neptune it will remain in a molecular state. Estimates also depend on the models for the formation of the planets. Condensation of the presolar nebula would give rise to a gaseous planet with the same composition as the Sun, but planets could also have formed when a solid core captured nebulous gas.[25]​.

In current models, the four giant planets have rock and ice cores that are approximately the same size, but the proportion of hydrogen and helium decreases from about 300 Earth masses on Jupiter, to 75 on Saturn and only a few on Uranus and Neptune.[25] Thus, while the gas giants are composed mainly of hydrogen and helium, the ice giants are composed mainly of heavier elements (O, C, N, S), mainly in the form of water, methane and ammonia. The surfaces are cold enough for molecular hydrogen to be liquid, so much of each planet is probably an ocean of hydrogen lying on top of heavier compounds. 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 on Saturn it can form a separate layer beneath the metallic hydrogen.[25]​.

terrestrial planets

The terrestrial planets are thought to come from the same nebular material as the giant planets, but have lost most of the lighter elements and have different histories. Planets closer to the Sun might be expected to have a higher fraction of refractory elements, but if their later stages of formation involved collisions with large objects with orbits that sampled different parts of the Solar System, there might be little systematic dependence on position.[30]​.

Direct information about Mars, Venus and Mercury comes largely from spacecraft missions. Using gamma ray spectrometers), the composition of the crust of Mars has been measured by the Mars Odyssey orbiter,[31]​ the crust of Venus by some of the Venus Venera missions,[30]​ and the crust of Mercury by the MESSENGER spacecraft.[32]​ Additional information about Mars comes from meteorites that have landed on Earth (shergottites, nakhlites, and chassignites), collectively known as SNC meteorites).[33]​ Abundances are also limited by the masses of the planets, while the internal distribution of elements is constrained by their moments of inertia.[4]​.

The condensed planets of the solar nebula, and many of the details of their composition, are determined by fractionation as they cool. The condensing phases are divided into five groups. The first to condense are materials rich in refractory elements such as Ca and Al. These are followed by nickel and iron, then magnesium silicates. Below 700 kelvins (700 K), metals and silicates rich in volatiles form a fourth group, and the fifth FeO group includes magnesium silicates.[34]​ The compositions of the planets and the Moon are chondritic, which means that within each group the proportions between the elements are the same as in carbonaceous chondrites.[4]​.

Estimates of planetary compositions depend on the model used. In the equilibrium condensation model, each planet formed from a feeding zone in which the compositions of the solids were determined by the temperatures in those zones. Thus, Mercury formed at 1400 K, when iron remains 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 it incorporated FeO into magnesium silicates. The biggest problem with this theory is that the volatiles would not condense, so the planets would have no atmospheres and the Earth would have no atmosphere.[4]​.

In chondritic mixing models, chondrite compositions are used to estimate planetary compositions. For example, one model mixes two components, one with the C1 chondrite composition and another with only the refractory components of C1 chondrites.[4]​ In another model, the abundances of the five fractionation groups are estimated using an index element for each of the groups. For the most refractory group, uranium is used; iron for the second; the proportions of potassium and thallium to uranium for the following two; and the molar ratio FeO/(FeO+MgO) for the latter. Using thermal and seismic models along with heat flow and density, Fe can be restricted within 10% on Earth, Venus and Mercury. U can be restricted to about 30% on Earth, but its abundance on other planets is based on "educated guesses." A difficulty with this model is that there can be significant errors in its prediction of volatile abundances because some volatiles are only partially condensed.[34]​[4]​.

Los constituyentes más comunes de las rocas son casi todos óxidos; los cloruros, sulfuros y fluoruros son las únicas excepciones importantes a esto y su cantidad total en cualquier roca es generalmente mucho menor del 1%. F. W. Clarke ha calculado que un poco más del 47% de la corteza terrestre está compuesta principalmente de oxígeno. Se presenta principalmente en combinación como óxidos, de los que los principales son la sílice, la alúmina, los óxidos de hierro y varios carbonatos (carbonato de calcio, carbonato de magnesio, carbonato de sodio y carbonato de potasio). La sílice funciona principalmente como un ácido, formando silicatos, y todos los minerales más comunes de las rocas ígneas son de esta naturaleza. A partir de un cálculo basado en 1672 análisis de numerosos tipos de rocas, Clarke llegó a la siguiente composición porcentual promedio de la corteza terrestre:.

Todos los demás constituyentes aparecen solo en cantidades muy pequeñas, generalmente mucho menores del 1% .[35]​

Estos óxidos se combinan de una manera casual. Por ejemplo, la potasa (carbonato de potasio) y la sosa (carbonato de sodio) se combinan para producir feldespatos. En algunos casos, pueden tomar otras formas, como nefelina, leucita y moscovita, pero en la gran mayoría de los casos se encuentran como feldespato. El ácido fosfórico con cal (carbonato de calcio) forma la apatita. El dióxido de titanio con óxido ferroso da lugar a la ilmenita. Parte de la cal forma feldespato de cal. El carbonato de magnesio y los óxidos de hierro con sílice cristalizan en forma de olivino o enstatita, o con alúmina y cal forman el complejo de silicatos ferro-magnésicos, de los que los piroxenos, anfíboles y biotitas son los principales. Cualquier exceso de sílice por encima de lo que se requiere para neutralizar las bases se separará como cuarzo; el exceso de alúmina cristaliza como corindón. Estas deben considerarse solo como tendencias generales. Es posible, según el análisis de la roca, decir aproximadamente qué minerales contiene la roca, pero hay numerosas excepciones a cualquier regla.[35]​.

mineral constitution

Except in acidic or siliceous igneous rocks containing more than 66% silica, known as felsic rocks, quartz is not abundant in igneous rocks. Basic rocks (containing 20% ​​silica or less) rarely contain as much silicon, and are known as mafic rocks. If magnesium and iron are above average, while silica is low, olivine can be expected; When silica is present in greater quantity on ferro-magnesium minerals, other minerals such as augite, hornblende, enstatite or biotite occur instead of olivine. Unless potash is high and silica relatively low, leucite will not be present, since leucite is not produced with free quartz. Nepheline, likewise, is generally found in rocks with a lot of soda and comparatively little silica. At high alkalis, soda-containing pyroxenes and amphiboles may be present. The lower the percentage of silica and alkalis, the greater the prevalence of plagioclase feldspar as contracted to soda or potassium feldspar.[35]​.

The Earth's crust is composed of 90% silicate minerals and their abundance on Earth is as follows: plagioclase feldspar" (39%), alkali feldspar (12%), quartz (12%), pyroxene (11%), amphiboles (5%), micas (5%), clay minerals (5%); the remaining silicate minerals make up another 3% of the Earth's crust. Only 8% of the Earth is composed of minerals that They are not silicates, like carbonates"), oxides and sulfides.[36]​.

The other determining factor, namely the physical conditions accompanying consolidation, generally plays a smaller part, but is by no means negligible. Certain minerals are practically confined to deep intrusive rocks, for example, microcline, muscovite, diallage. Leucite is very rare in plutonic masses; Many minerals have special peculiarities in their microscopic character depending on whether they crystallized at depth or near the surface, for example, hypersthene, orthoclase, quartz. There are some curious examples of rocks having the same chemical composition, but consisting of completely different minerals, for example, the hornblende from Gran, in Norway, which contains only hornblende, has the same composition as some of the camptonites from the same locality which contain feldspar and hornblende of a different variety. In this sense, what has been said previously about the corrosion of porphyritic minerals in igneous rocks can be repeated. In rhyolites and trachytes, early hornblende and biotite crystals can be found in large quantities partially converted to augite and magnetite. Hornblende and biotite remained stable under pressures and other conditions below the surface, but are unstable at higher levels. In the soil mass of these rocks, augite is almost universally present. But the plutonic representatives of the same magma, granite and syenite contain biotite and hornblende much more commonly than augite.[35]​.

Felsic, intermediate and mafic igneous rocks

Those rocks that contain the greatest amount of silica, and when crystallized leave free quartz, form a group generally called "felsic" rocks. Those that again contain the least silica and the most magnesia and iron, so that quartz is absent, while olivine is usually abundant, form the "mafic" group. "Intermediate" rocks include those characterized by the general absence of both quartz and olivine. An important subdivision of these contains a very high percentage of alkalis, especially soda, and consequently has minerals such as nepheline and leucite that are not common in other rocks. It is often separated from the others as "alkaline" or "soda" rocks, and there is a corresponding series of mafic rocks. Finally, a small subgroup rich in olivine and lacking feldspar has been called “ultramafic” rocks. They have very low percentages of silica but a lot of iron and magnesia.

Except for the latter, practically all rocks contain feldspars or feldspathoid minerals. In acidic rocks, common feldspars are orthoclase, perthite, microcline, and oligoclase, all with lots of silica and alkalis. Labradorite, anorthite and bytownite prevail in mafic rocks, which are rich in lime and poor in silica, potash and soda. Augite is the most common ferromagnesian in mafic rocks, but biotite and hornblende are generally more common in felsic rocks.[35]​.

Rocks containing leucite or nepheline, either partially or completely replacing feldspars, are not included in this table. They are essentially of an intermediate or mafic character. Accordingly, they might be regarded as varieties of syenite, diorite, gabbro, etc., in which feldspathoids occur, and in fact there are many transitions between syenites of the ordinary type and nepheline—or leucite—syenite, and between gabbro or dolerite and mineralite. or sexy. But, since many minerals develop in these "alkaline" rocks that are not common elsewhere, it is convenient in a purely formal classification such as that described here, to treat the whole assemblage as a distinct series.[35]​.

This classification is essentially based on the mineralogical constitution of igneous rocks. Any chemical distinction between the different groups, although implicit, is relegated to a subordinate position. It is true that it is artificial, but it has grown with the growth of science and is still adopted as the basis on which still smaller subdivisions are erected. The subdivisions are in no way of equal value. Syenites, for example, and peridotites, are much less important than granites, diorites and gabbros. Furthermore, effusive andesites do not always correspond to plutonic diorites but partly also to gabbros. As different types of rock, considered aggregates of minerals, gradually pass over one another, transitional types are very common and are often important enough to be given special names. Quartz syenites and nordmarkites") can come between granite and syenite, tonalites and adamelites between granite and diorite, monzoaites between syenite and diorite, norites and hyperites") between diorite and gabbro, and so on.[35]​.

Trace metals readily form "Complex (chemistry)") complexes with major ions in the ocean, including hydroxide, carbonate, and chloride, and their changes in chemical speciation depend 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 water molecules and form strong bonds with metals that have a charge of more than two. A consequence of the lower reactivity of complexed metals compared to the same concentration of free metal is that chelation tends to stabilize the metals in aqueous solution rather than in solids.[38]​.

The concentrations of some trace metals—cadmium, copper, molybdenum, manganese, rhenium, uranium, and vanadium—in certain sediments record the redox history of the oceans. In aquatic environments, cadmium(II) can be in the form CdCl in oxic waters or CdS(s) in a reduced environment. Therefore, higher concentrations of Cd in marine sediments may indicate conditions of low redox potential in the past. For copper (II), a prevalent form is CuCl(aq) in oxic environments and CuS(s) and CuS in reduced environments. The reduced seawater environment leads to two possible oxidation states of copper, Cu(I) and Cu(II). Molybdenum is present as the oxidation state of Mo(VI) as MoO in oxic environments. Rhenium is present as the oxidation state Re (VII) as ReO under oxic conditions, but is reduced to Re (IV) which can form ReO or ReS Uranium is in the oxidation state VI in UUO(CO)(aq) and is found in the reduced form UO(s). Vanadium occurs in various forms in the oxidation state V (V) and HVO; This relative predominance of these species depends on pH.

In the water column of the ocean or deep lakes, the vertical profiles of dissolved trace metals are characterized by following the following distributions: conservative type (conservative–type), nutrient type (nutrient–type) or scavenged–type*. Across these three distributions, trace metals have different residence times and are used to different extents by planktonic microorganisms. Trace metals with conservative distributions have high concentrations relative to their biological use. An example of a trace metal with a conservative distribution is molybdenum. It has a residence time in the oceans of about 8 x 10 years and is generally present as the anion molybdate (MoO). Molybdenum interacts weakly with particles and shows a nearly uniform vertical profile. in the ocean. Relative to the abundance of molybdenum in the ocean, the amount required as a metal cofactor for enzymes in marine phytoplankton is negligible.[39]​.

Trace metals with nutrient-like distributions are strongly associated with the internal cycles of particulate organic matter, especially assimilation by plankton. The lowest dissolved concentrations of these metals are found at the ocean surface, where they are assimilated by plankton. As dissolution and decomposition occur at greater depths, the concentrations of these trace metals increase. The residence times of these metals, such as zinc, are several thousand to one hundred thousand years. Finally, an example of a swept-type trace metal is aluminum, which has strong interactions with particles and a short residence time in the ocean. Residence times for sweep-type trace metals are around 100 to 1000 years. Concentrations of these metals are highest around bottom sediments, hydrothermal vents, and rivers. For aluminum, atmospheric dust provides the largest source of external inputs to the ocean.[39]​.

Iron and copper show hybrid distributions in the ocean. They are influenced by recycling and intense cleaning. Iron is a limiting nutrient in vast areas of the oceans, and is found in great abundance along with manganese near hydrothermal vents. Here, many iron precipitates are found, mainly in the form of iron sulfides and oxidized iron oxyhydroxide compounds. Iron concentrations near hydrothermal vents can be up to a million times the concentrations found in the open ocean.[39]​.

Using electrochemical techniques, it is possible to show that bioactive trace metals (zinc, cobalt, cadmium, iron and copper) are bound by organic ligands on the surface of seawater. These ligand complexes serve to decrease the bioavailability of trace metals in the ocean. For example, copper, which can be toxic to phytoplankton and bacteria in the open ocean, can form organic complexes. The formation of these complexes reduces concentrations of bioavailable inorganic copper complexes that could be toxic to marine life at high concentrations. Unlike copper, the toxicity of zinc in marine phytoplankton is low and there is no advantage to increasing organic binding of Zn. In high-chlorophyll, low-nutrient regions, iron is the limiting nutrient, and the dominant species are strong organic Fe(III) complexes.[39]​.