Silicon and oxygen bond covalently to create a silica tetrahedron, which is a four-sided pyramid shape with O at each corner and Si in the middle (Figure 5.14). This structure is the building block of the many important silicate minerals. The bonds in a silica tetrahedron have some of the properties of covalent bonds and some of the properties of ionic bonds. As a result of the ionic character, silicon becomes a cation (with a charge of +4) and oxygen becomes an anion (with a charge of –2). The net charge of a silica tetrahedron (SiO4) is –4.
In silicate minerals, these tetrahedra are arranged and linked together in a variety of ways, from single units to chains, rings, and more complex frameworks. In the rest of this section we will discuss the structures of the most common silicate minerals in Earth’s crust and mantle.
Exercise 5.3 Make a Tetrahedron
Cut around the outside of the shape (solid lines and dotted lines), and then fold along the solid lines to form a tetrahedron.
If you have glue or tape, secure the tabs to the tetrahedron to hold it together. If you don’t have glue or tape, make a slice along the thin grey line and insert the pointed tab into the slit.
If you’re feeling ambitious, make several tetrahedra and try joining them into pairs, rings, single and double chains, sheets, and even three-dimensional frameworks.
Olivine Consists of Isolated Tetrahedra Linked by Cations
The simplest silicate structure, that of the mineral olivine (Figure 5.15), is composed of isolated tetrahedra bonded to iron and/or magnesium ions (Figure 5.15 left). In olivine, the –4 charge of each silica tetrahedron is balanced by two iron or magnesium cations, each with a charge of +2. Olivine can be pure Mg2SiO4 or Fe2SiO4, or some combination of the two, written as (Mg,Fe)2SiO4. The magnesium and iron cations are close in radius (0.73 Å versus 0.62 Å ). Because of this size similarity, and because they both have a charge of +2, iron and magnesium can readily substitute for each other in olivine and in many other minerals.
Although the iron and magnesium ions are similar in size, allowing them to substitute for each other in some silicate minerals, the common ions in silicate minerals have a wide range of sizes (Figure 5.16). Ionic radii are critical to the composition of silicate minerals, so we’ll be referring to this diagram again.
Pyroxene and Amphibole Are Chain Silicates
Pyroxene (Figure 5.17 bottom left) is an example of a single-chain silicate. The structure of chain silicates is shown in Figure 5.17 (top). In pyroxene, silica tetrahedra form a chain because one oxygen from each tetrahedron is shared with the adjacent tetrahedron. This means there are fewer oxygens in the structure. This can be expressed as an oxygen-to-silicon ratio (O:Si). The O:Si is lower than in olivine (3:1 instead of 4:1), and the net charge per silicon atom is less (–2 instead of –4), since fewer cations are necessary to balance that charge. Pyroxene compositions have the silica tetrahedra represented as SiO3 (e.g., MgSiO3, FeSiO3, and CaSiO3.) In other words, pyroxene has one cation for each silica tetrahedron (e.g., MgSiO3) while olivine has two (e.g., Mg2SiO4). The structure of pyroxene is more “permissive” than that of olivine, meaning that cations with a wider range of ionic radii can fit into it. That’s why pyroxenes can have calcium cations (radius 1.00 Å) substitute for iron (0.63 Å) and magnesium (0.72 Å) .
Exercise 5.4 Silicon to Oxygen Ratios
The diagram below represents a single chain in a silicate mineral. Count the number of tetrahedra versus the number of oxygen ions (yellow spheres). Each tetrahedron has one silicon ion. Confirm for yourself that the ratio of silicon to oxygen for the single chain is 3:1.
The diagram below represents a double chain in a silicate mineral. Again, count the number of tetrahedra versus the number of oxygen ions. What is the ratio of Si to O in double-chain silicates (e.g., amphibole)?
Mica and Clay Minerals Are Sheet Silicates
In mica structures the silica tetrahedra are arranged in continuous sheets (Figure 5.18 top), where each tetrahedron shares three oxygen anions with adjacent tetrahedra. Because more oxygens are shared between adjacent tetrahedra, fewer charge-balancing cations are needed for sheet silicate minerals. Bonding between sheets is relatively weak, and this accounts for the tendency of mica minerals to split apart in sheets (referred to as one-directional cleavage; Figure 5.18 bottom right). Two common micas in silicate rocks are biotite (Figure 5.18 bottom left), which contains iron and/or magnesium, making it a dark mineral; and muscovite (Figure 5.18 right) which contains aluminum and potassium, and is light in colour.
Apart from muscovite, biotite, and chlorite, there are many other sheet silicates (or phyllosilicates), which usually exist as clay-sized fragments (i.e., less than 0.004 mm). These include the clay minerals kaolinite, illite, and smectite, and although they are difficult to study because they are occur as very small particles, they are extremely important components of rocks and especially of soils. All of the sheet silicate minerals also have water in their structure. It occurs as the OH– anion.
Quartz and Feldspar Are Framework Silicates
In framework silicates, three or four oxygen on a silica tetrahedron are shared with other silica tetrahedra. This means the tetrahedra are connected to each other in three dimensions rather than in two-dimensional chains and sheets.
Feldspars (Figure 5.19) are a very important group of minerals with three-dimensional frameworks. The three main feldspar minerals are potassium feldspar, (also referred to as K-feldspar or K-spar; Figure 5.19 right) and two types of plagioclase feldspar: albite (NaAlSi3O8; Figure 5.19 left) and anorthite (CaAl2Si3O8; Figure 5.19 centre). As is the case for iron and magnesium in olivine, there is a continuous range of compositions (referred to as a solid solution series) between albite and anorthite in plagioclase. This is because the calcium and sodium ions are almost identical in size (1.00 Å versus 0.99 Å). Any intermediate compositions between CaAl2Si3O8 and NaAlSi3O8 can exist. This is a little bit surprising because calcium and sodium ions don’t have the same charge (Ca2+ versus Na+). This problem is accounted for substituting some Al3+ for Si4+. Albite has one Al and three Si, while anorthite is has two Al and two Si. Plagioclase feldspars of intermediate composition have intermediate proportions of Al and Si.
K-feldspar (KAlSi3O8; Figure 5.19 right) has a slightly different structure than that of plagioclase, owing to the larger size of the potassium ion (1.37 Å). It is possible for potassium and sodium to substitute for each other, but this requires very high temperatures.
Quartz (SiO2; Figure 5.20) contains only silica tetrahedra. In quartz, each silica tetrahedron is bonded to four other tetrahedra (with an oxygen shared at every corner of each tetrahedron), making a three-dimensional framework. As a result, the ratio of silicon to oxygen is 1:2. Since the one silicon cation has a +4 charge and the two oxygen anions each have a –2 charge, the charge is balanced. There is no need to add cations to balance the charge. The hardness of quartz and the fact that it breaks irregularly (notice the bottom of the crystal in Figure 5.20 right) and not along smooth planes result from the strong covalent/ionic bonds characteristic of the silica tetrahedron.