Mineral - Mineral Chemistry

Mineral Chemistry

The abundance and diversity of minerals is controlled directly by their chemistry, in turn dependent on elemental abundances in the Earth. The majority of minerals observed are derived from the Earth's crust. Eight elements account for most of key components of minerals, due to their abundance in the crust. These eight elements, summing to over 98% of the crust by weight, are, in order of decreasing abundance: oxygen, silicon, aluminium, iron, magnesium, calcium, sodium and potassium. Oxygen and silicon are by far the two most important—oxygen composes 46.6% of the crust by weight, and silicon accounts for 27.7%.

The minerals that form are directly controlled by the bulk chemistry of the parent body. For example, a magma rich in iron and magnesium will form mafic minerals, such as olivine and the pyroxenes; in contrast, a more silica-rich magma will crystallize to form minerals than incorporate more SiO₂, such as the feldspars and quartz. In a limestone, calcite or aragonite (both CaCO₃) are present because the rock is rich in calcium and carbonate. A corollary is that a mineral will not be found in a rock whose bulk chemistry does not resemble the bulk chemistry of a given mineral with the exception of trace minerals. For example, kyanite, Al₂SiO₅ forms from the metamorphism of aluminium-rich shales; it would not likely occur in aluminium-poor rock, such quartzite.

The chemical composition may vary between end member species of a mineral series. For example, the plagioclase feldspars comprise a continuous series from sodium-rich end member albite (NaAlSi₃O₈) to calcium-rich anorthite (CaAl₂Si₂O₈) with four recognized intermediate varieties between them (given in order from sodium- to calcium-rich): oligoclase, andesine, labradorite, and bytownite. Other examples of series include the olivine series of magnesium-rich forsterite and iron-rich fayalite, and the wolframite series of manganese-rich hübnerite and iron-rich ferberite.

Chemical substitution and coordination polyhedra explain this common feature of minerals. In nature, minerals are not pure substances, and are contaminated by whatever other elements are present in the given chemical system. As a result, it is possible for one element to be substituted for another. Chemical substitution will occur between ions of a similar size and charge; for example, K+ will not substitute for Si4+ because of chemical and structural incompatibilities caused by a big difference in size and charge. A common example of chemical substitution is that of Si4+ by Al3+, which are close in charge, size, and abundance in the crust. In the example of plagioclase, there are three cases of substitution. Feldspars are all framework silicates, which have a silicon-oxygen ratio of 2:1, and the space for other elements is given by the substitution of Si4+ by Al3+ to give a base unit of -; without the substitution, the formula would be charge-balanced as SiO₂, giving quartz. The significance of this structural property will be explained further by coordination polyhedra. The second substitution occurs between Na+ and Ca2+; however, the difference in charge has to accounted for by making a second substitution of Si4+ by Al3+.

Coordination polyhedra are geometric representation of how a cation is surrounded by an anion. In mineralogy, due its abundance in the crust, coordination polyhedra are usually considered in terms of oxygen. The base unit of silicate minerals is the silica tetrahedron—one Si4+ surrounded by four O2-. An alternate way of describing the coordination of the silicate is by a number: in the case of the silica tetrahedron, the silicon is said to have a coordination number of 4. Various cations have a specific range of possible coordination numbers; for silicon, it is almost always 4, except for very high-pressure minerals where compound is compressed such that silicon is in six-fold (octahedral) coordination by oxygen. Bigger cations have a bigger coordination number because of the increase in relative size as compared to oxygen (the last orbital subshell of heavier atoms is different too). Changes in coordination numbers between leads to physical and mineralogical differences; for example, at high pressure such as in the mantle, many minerals, especially silicates such as olivine and garnet will change to a perovskite structure, where silicon is in octahedral coordination. Another example are the aluminosilicates kyanite, andalusite, and sillimanite (polymorphs, as they share the formula Al₂SiO₅), which differ by the coordination number of the Al3+; these minerals transition from one another as a response to changes in pressure and temperature. In the case of silicate materials, the substitution of Si4+ by Al3+ allows for a variety of minerals because of the need to balance charges.

Changes in temperature and pressure, and composition alter the mineralogy of a rock sample. Changes in composition can be caused by processes such as weathering or metasomatism (hydrothermal alteration). Changes in temperature and pressure occur when the host rock undergoes tectonic or magmatic movement into differing physical regimes. Changes in thermodynamic conditions make it favourable for mineral assemblages to react with each other to produce new minerals; as such, it is possible for two rocks to have an identical or a very similar bulk rock chemistry without having a similar mineralogy. This process of mineralogical alteration is related to the rock cycle. An example of a series of mineral reactions is illustrated as follows.

Orthoclase feldspar (KAlSi₃O₈) is a mineral commonly found in granite, a plutonic igneous rock. When exposed to weathering, it reacts to form kaolinite (Al₂Si₂O₅(OH)₄, a sedimentary mineral, and silicic acid):

2 KAlSi₃O₈ + 5 H₂O + 2 H+ → Al₂Si₂O₅(OH)₄ + 4 H₂SiO₃ + 2 K+

Under low-grade metamorphic conditions, kaolinite reacts with quartz to form pyrophyllite (Al₂Si₄O₁₀(OH)₂):

Al₂Si₂O₅(OH)₄ + SiO₂ → Al₂Si₄O₁₀(OH)₂ + H₂O

As metamorphic grade increases, the pyrophyllite reacts to form kyanite and quartz:

Al₂Si₄O₁₀(OH)₂ → Al₂SiO₅ + 3 SiO₂ + H₂O

Alternatively, a mineral may change its crystal structure as a consequence of changes in temperature and pressure without reacting. For example, quartz will change into a variety of its SiO₂ polymorphs, such as tridymite and cristobalite at high temperatures, and coesite at high pressures.