Which silicate minerals are the main component

Magmas richest in SiO 2 , such as alkali granite, granite, and tonalite are generally deficient in MgO. We term such magmas silicic Si-rich or felsic contraction of feldspar and silica. Light-colored minerals dominate felsic rocks, so many geologists use the term felsic to refer to any light-colored igneous rock, even if the chemical composition is unknown. These include diorite, gabbro, and peridotite. We call them mafic contraction of the words magnesium and ferric and, for peridotite, ultramafic.

They are usually dark in color. The term intermediate describes rocks with compositions between mafic and silicic but no intermediate rock analyses are in the table above. The table above reveals that alkalic rocks can be either silicic or mafic compositions.

Overall, though, the most common alkalic rocks are silicic. We grouped alkali syenite and syenite with the other mafic magmas because alkali syenite and syenite are silica-poor compared with silicic magmas, but their most significant distinctions are their high alkali oxide contents compared with CaO content.

The photo seen here Figure 6. The main minerals are gray potassium feldspar and lighter colored nepheline. The black mineral is hornblende. Rocks melt in many places within Earth, and magma compositions reflect the sources. Mid-ocean ridge and ocean hot-spot magmas are mostly mafic; subduction zone magmas are generally silicic to intermediate.

Continental rifts produce a variety of magma types. Rocks of different compositions have different melting temperatures because some elements combine to promote melting. Silicon and oxygen, in particular, promote melting because they form very stable molten polymers long chains of Si and O that persist even when melted.

Silicic minerals, and SiO 2 -rich rocks, therefore, melt at lower temperatures than mafic minerals and SiO 2 -poor rocks.

Magmas may also contain gases, liquids, or vapors, collectively called volatiles. H 2 O and CO 2 are the most common volatiles, but volatile compounds of sulfur, chlorine, and several other elements may also be present. Sometimes, these compounds separate from a melt to form bubbles, most commonly in cooling lava, creating empty vesicles as the magma solidifies.

Often, vesicles like these become filled with secondary zeolites or other minerals over time. This rock contains green crystals of olivine, but most of it is dark colored volcanic glass. The specimen also contains a piece of black and white gabbro that was plucked from the lower crust as the magma rose to the surface. Water is especially important during eruptions because just a small amount of water can produce large amounts of steam and violent eruptions. This is especially true for silicic magmas.

The explosive May 18, eruption of Mt. Saint Helens in Washington is an example. A series of earthquakes and small eruptions began at Mt. Saint Helens in March, The eruption produced a column of ash that rose 24 km into the sky.

Ash was deposited in a dozen states and two Canadian provinces. Every mineral has a characteristic melting temperature. This can be expected to lead to an orderly and predictable sequence of minerals crystallizing as magma cools and solidifies.

It does, sort of. Complications arise because some minerals do not crystallize at a single temperature, but instead form from other minerals while reacting with magma. This is called incongruent melting. Further complications arise because minerals together may melt and crystallize at lower temperatures than if they if they were alone. This is called eutectic melting.

Crystallization of a cooling magma is a step-wise process with some minerals forming before others. So, crystallization occurs over a range of temperature. As seen in Figure 6. Crystallization continues until the magma reaches the solidus temperature. Above the liquidus, all is melted. Below the solidus, all is solid. Between the two temperatures a rock is partially melted. The differences between liquidus and solidus temperatures can be s of degrees and is different for different composition magmas.

Different magmas produce different minerals, and different minerals crystallize at different temperatures. Consequently, composition is the key factor that determines the temperature at which crystallization begins. For mafic rocks, the first minerals to crystallize are typically olivine, pyroxene, or plagioclase.

For silicic rocks, the first crystals may be alkali feldspars or micas. Eventually, a magma completely solidifies, and the last drop of melt crystallizes at the solidus temperature. Bowen pioneered in the study of magma crystallization, for which he received the Roebling Medal from the Mineralogical Society of America in By studying naturally occurring igneous rocks and conducting laboratory experiments, he derived an idealized model for equilibrium crystallization in a magmatic system.

Some petrologists have developed more precise models for magmas of specific compositions. It contains nine mineral names arranged in a Y-shape. We call the left-hand side of the Y the discontinuous side because abrupt changes occur as different minerals crystallize in sequence. We call the right-hand side the continuous side because plagioclase is continually present during crystallization, starting as Ca-rich plagioclase at high temperature and changing to Na-rich plagioclase as cooling progresses.

Although not evident from this diagram, just like plagioclase, minerals on the discontinuous side of the series change composition as cooling proceeds.

At higher temperatures, for example, olivine usually has a greater magnesium to iron ratio Mg:Fe than at lower temperatures. Mineral crystallization temperatures depend on mineral composition. Minerals at the top of the series olivine, pyroxene, and calcium-rich plagioclase are mafic relatively silica poor. Those at the bottom are silicic relatively silica rich.

Silica content is the most significant factor controlling melting and crystallization temperatures. The mafic minerals at the top of the discontinuous series also are deficient in aluminum and alkalis and rich in iron and magnesium, compared with minerals at the bottom.

It will skip all the other minerals and just form quartz. And, although all magmas crystallize different minerals at different temperatures, none follow the complete series. Mafic magmas, which crystallize at high temperature, produce rocks containing minerals at the top of the series. Silicic magmas, which crystallize at lower temperatures, produce rocks that contain minerals at the bottom of the series.

So mafic rocks such as basalt or gabbro commonly contain olivine, pyroxene and Ca-rich plagioclase. Felsic rocks such as rhyolite or granite are generally rich in K-feldspar and quartz. And, intermediate composition rocks may contain pyroxene, amphibole, or biotite with plagioclase. If magma and minerals remain in equilibrium, different minerals crystallize at different temperatures — sometimes more than one at a time, and mineral compositions change as temperature decreases.

Continuous reactions take place as elements move from magma into growing crystals. These relationships are orderly and predictable for a magma of any given composition.

But, they are highly dependent on composition. Granite and gabbro, for example, do not crystallize the same minerals, nor do they crystallize at the same temperatures. However, mineral crystals and magmas do not always remain in equilibrium during crystallization.

Several things may cause disequilibrium. For example, as depicted in Figure 6. The crystals, then, will not stay in equilibrium with the magma above them. As this happens, an original parental magma changes composition and becomes an evolved magma.

Many minerals, including feldspars, pyroxenes, and oxides, can be found in thick, nearly single-mineral cumulates. And, even if a thick cumulate layer does not develop, the composition of the upper part of the magma chamber may not be the same as the lower part. When a melt gets separated from early formed crystals, we call the process partial crystallization , or fractional crystallization.

The photos above in figures 6. The photo on the left Figure 6. The light material around the chromite layers is mostly the plagioclase. Like the Bushveld, the Stillwater Complex sees significant mining activity.

Mafic and ultramafic layered intrusions, often called layered mafic complexes , are rare but are found worldwide — generally in old continental interiors. The largest, the Bushveld Complex of South Africa, is more than 66, km 2. Smaller layered intrusions, such as the Skaergaard Complex Greenland , may be only km 2 or less. These complexes contain mafic rocks with compositional layering caused by fractional crystallization.

They have cumulate layers piled on top of each other. Early formed crystals settled to the bottom of the magma chamber because they were denser than the melt. As the magma cooled and other minerals crystallized, they too settled. Consequently a layered sequence developed with ultramafic high-temperature minerals at the bottom and successively more silicic lower-temperature minerals on top. New pulses of magma added more layers. Besides silicates, chromite and other oxide minerals accumulate in mafic complexes.

Rarely, plagioclase forms a cumulate layer at the top of a magma chamber when it floats on denser magma. Because of their high metal content and natural separation into concentrated layers, mafic complexes often host rich ore deposits. They are especially important for production of chrome and platinum group metals ruthenium, rhodium, palladium, osmium, iridium, platinum.

The Bushveld complex is, perhaps, the most valuable ore deposit in the world, producing significant amounts of platinum group metals, as well as chromium, iron, tin, titanium, and vanadium. The photo seen in Figure 6. The metallic minerals are pyrrhotite and chalcopyrite that contain significant amounts of platinum, palladium, osmium, and other generally rare elements. Besides separation of crystals and melt, disequilibrium occurs for other reasons.

Sometimes large mineral grains do not remain in equilibrium with a surrounding magma. For example, because diffusion of elements through solid crystals is slow, different parts of large crystals may not have time to maintain equilibrium compositions. In principle, mineral crystals should be homogeneous, but in zoned crystals, such as the tourmaline crystals seen in this photo Figure 6.

Consequently, different zones have different compositions and sometimes different colors. Marked chemical zonation often occurs if a magma begins to cool at one depth and then rapidly moves upward to cooler temperatures.

The result is often a porphyritic rock with large zoned phenocrysts. The zones may be visible with the naked eye, for example the tourmaline in Figure 6. Other examples of visibly zoned minerals include the tourmaline in Figure 4. If the zoning is not visible to the naked eye, it may be visible when a crystal is viewed in thin section with a petrographic microscope.

The thin section photo seen here Figure 6. The colors are artifacts and are not real mineral colors. Zoning can also be detected using a scanning electron microscope. See, for example, the electron microscope images of zoned plagioclase in Figure 4. Silicate minerals dominate igneous rocks because silicon and oxygen are the most common elements in magmas. So, in the following discussion we systematically consider the important silicate minerals and groups.

Recall from Chapter 1 that the fundamental building block in silicate minerals is an SiO 4 4- tetrahedron with oxygen at the corners and silicon in the center. Individual tetrahedra bond to other tetrahedra or to cations to make a wide variety of atomic arrangements. The top drawing in this chart Figure 6. Sometimes aluminum substitutes for silicon, so silicate minerals may contain both SiO 4 and AlO 4 tetrahedra.

A key property of both silicon and aluminum tetrahedra is that they can share the oxygen anions O 2- at their corners. When they do this, they form polymers that are various kinds of rings, chains, sheets, or three-dimensional arrangements.

For a slightly different view, see Figure 1. In some minerals, SiO 4 4- tetrahedra are not polymerized. They do not share oxygen atoms, and instead are joined together by bonds to other cations. These minerals, called isolated tetrahedra silicates , or island silicates , include most importantly minerals of the olivine group.

In a few minerals, especially minerals of the epidote group, tetrahedra join to form pairs. These are the paired tetrahedral silicates , sometimes called sorosilicates or butterfly silicates.

In the pyroxenes and other single chain silicates , tetrahedra link to create zigzag chains. The amphibole minerals are double chain silicates. Micas and clays are sheet silicates. In sheet silicates, tetrahedra share three of their four oxygen with other tetrahedra, creating sheets and minerals with layered atomic arrangements.

Finally, the feldspars and quartz are examples of network silicates , sometimes called framework silicates , in which every oxygen is shared between two tetrahedra, creating a three-dimensional network. The drawing in Figure 6. The most important framework silicates are quartz and other SiO 2 minerals, and the feldspars. In the feldspars, and a related group of minerals called feldspathoids , alkali and alkaline earth elements — mostly Na, K, or Ca — occupy large sites between tetrahedra.

We will start our review of silicate minerals by looking at the SiO 2 polymorphs, the feldspars, and the feldspathoids, and then work our way upwards in the figure above, towards minerals with less polymerization.

Silica Minerals SiO 2. Quartz, like many other minerals, is polymorphic. Mineralogists and chemists have identified more than 10 different silica SiO 2 polymorphs, but some do not occur as minerals. We briefly looked at the most common of these polymorphs in Chapter 4.

All the silica minerals except stishovite are framework silicates; the differences between the minerals are the angular relationships between the tetrahedra that comprise them. The blue spheres are oxygen atoms and silicon atoms are at the centers of every gray tetrahedron. For some spectacular scanning electron microscope images of several silica polymorphs, see Figure Common quartz, more properly called low quartz because it has lower symmetry than high quartz , is the only polymorph stable under normal Earth surface conditions, but it has many different appearances.

The anhedral specimen seen in Figure 6. But euhedral crystals, such as those shown in Figure 6. Some of the different quartz varieties have specific names, such as milky quartz , rose quartz , Herkimer diamond , amethyst , and citrine. For several decades, petrologists have understood that different silica polymorphs occur in different geological settings because they are stable under different pressure-temperature conditions. This phase diagram Figure 6. The horizontal axis is temperature.

The vertical scale on the left gives pressures in gigapascals GPa , and the scale on the right shows the depths in Earth corresponding to those pressures. Pressure-temperature P-T phase diagrams such as the one seen here show which mineral is stable for any combination of P-T. Low quartz is therefore the most common polymorph.

If all rocks maintained and stayed at equilibrium, we would have no samples of any other silica polymorphs to study. Although given enough time, they usually change into low quartz.

They are usually associated with meteorite impact craters. Tridymite and cristobalite only exist in certain high temperature silicic volcanic rocks. Although not shown in this diagram, just like quartz, tridymite and cristobalite have both high- and low-symmetry polymorphs.

The red field in the phase diagram above Figure 6. Melting temperature is greatest at high pressure, and is different for the different polymorphs.

Consider what happens when a volcano erupts and silica-rich magma cools. This happens most of the time but occasionally metastable polymorphs can be found in volcanic rocks.

Essential minerals are minerals that must be present for a rock to have the name that it does, and quartz is an essential mineral in silicic and intermediate igneous rocks, many sediments, and many metamorphic rocks. Quartz is not normally found in mafic igneous rocks because crystallization of mafic minerals such as olivine or pyroxene generally consumes all silica that is available, so there is none left over to form quartz.

Granites contain essential quartz. In silicic plutonic rocks such as granite, quartz is always associated with K-feldspar, commonly in a mosaic pattern similar to what is seen in this photograph Figure 6.

The largest of the grains in this view are pinkish K-feldspar about 1 cm across. Quartz is glassy gray. White plagioclase and black biotite are also present. Quartz is also an essential mineral in sandstone and some other sedimentary rocks. Quartz is the only mineral present in some sandstones or cherts. But, sandstone may also contain significant amounts of other minerals including feldspar or clay, and sometimes pebbles or rock fragments.

Quartz cannot exist in rocks containing corundum Al 2 O 3 , because the two minerals would react to form an aluminosilicate mineral of some sort.

It cannot exist in rocks containing feldspathoids leucite, nepheline, or analcime because quartz and feldspathoids react to give feldspars. For similar reasons, quartz is absent or minor in many alkali-rich igneous rocks and in rocks containing the oxide mineral spinel MgAl 2 O 4.

The photograph on the left above Figure 6. The quartz formed when hot hydrothermal water infiltrated the rock along fractures before the rock was uplifted to the surface and eventually weathered. Temperature need not be high, however, for quartz to precipitate.

For example, the amethyst purple quartz in the geode shown in Figure 6. Figure 4. Most quartz crystals are twinned, but the twinning can be impossible to see or easily overlooked. The drawings seen here Figue 6. Brazil twins and Dauphine twins are penetration twins, and Japanese twins are contact twins. All three are generally growth twins but can also form in other ways. Some quartz crystals exhibit more than one kind of twinning.

Brazil and Dauphine twins are distinguished by symmetry relationships between crystal faces of particular shapes, sometimes by identifying striations fine lines on crystal faces that developed when the crystal formed , but sometimes are difficult to tell apart. Click on the crystal drawing right to see a 15 second video showing Dauphine twinning; the striations on the crystal faces are also apparent.

The two large quartz crystals in Figure 6. They are widespread and are essential minerals in many igneous, metamorphic, and sedimentary rocks. Rarely, they contain significant amounts of other elements such as Ba, Sr, B, or Fe. For most purposes, we consider them to be ternary solutions, which means we can describe their composition in terms of three end members and plot them on diagrams such as the one shown in Figure 6. In Figure 6. Natural feldspars form two distinct series, the alkali feldspar series and plagioclase , both labeled on this triangular diagram.

We call any feldspar with composition near NaAlSi 3 O 8 , albite , and one with composition near CaAl 2 Si 2 O 8 , anorthite , even if other components are present. Intermediate plagioclase compositions are commonly given specific names labeled in the figure : oligoclase , andesine , labradorite , and bytownite. Labradorite may also contain a small amount of orthoclase.

Compositions between plagioclase and alkali feldspar that would plot in the white part of the triangle are rare or do not exist. Confusion sometimes arises because the names of some composition ranges are the same as the names of feldspar end members albite, anorthite, orthoclase.

Orthoclase, for example, is the name given to end member KAlSi 3 O 8. The triangular diagram in Figure 6. This feldspar is an example of anorthoclase. It has a composition of about Ca 0. Often, we describe the compositions of feldspars by using abbreviations with subscripts. Thus, the Grorud feldspar has composition An 3 Ab 62 Or Like quartz, feldspars are framework silicates. Unlike quartz, feldspars contain both SiO 4 and AlO 4 tetrahedra.

As an example, Figure 6. The purple atoms are sodium, and the yellow and green tetrahedra are SiO 4 and AlO 4 , respectively. In all feldspars, Na, K, or Ca occupy spaces between tetrahedra.

Most igneous rocks contain feldspar of some sort, but the kind of feldspar varies with rock composition. In silicic igneous rocks, such as granite, plagioclase is absent or subordinate to K-rich alkali feldspar.

If plagioclase is present, it is always albite-rich. Similarly, in mafic rocks, alkali feldspar is not normally present but plagioclase is common. Because mafic rocks contain much more Ca than Na, the plagioclase in them is generally anorthite-rich.

Intermediate igneous rocks nearly always contain both feldspars. Vesuvius, Italy. The two photos above show light colored albite crystals from a classic locality in Austria and anorthite crystals from near Naples, Italy. These are both examples of plagioclase. But, K-feldspar has three polymorphs with slightly different atomic arrangements: orthoclase , microcline , and sanidine. Photos of each are shown below in Figures 6.

In these photos the polymorphs have different colors, but color is not a good diagnostic property because the different polymorphs all come in many colors. Note the well developed penetration twins in the sanidine specimen. Sanidine forms in high temperature rocks; orthoclase and microcline form in lower temperature rocks Figure 6. Sanidine commonly crystallizes from silicic lava but may change into orthoclase and, perhaps, microcline if the lava cools relatively slowly.

Just about all microcline forms by recrystallization of sanidine or orthoclase. Like K-feldspar, Na-feldspar forms different polymorphs, with different atomic arrangements, depending on temperature Figure 6.

Three other polymorphs exist at lower temperatures, all generally called albite. Unlike the SiO 2 polymorphs, the differences in atomic structure between the feldspar polymorphs are not great and the boundaries on phase diagrams are poorly known.

Distinguishing between the different Na-feldspar polymorphs can be difficult or impossible without X-ray analysis. The striations derive from a type of polysynthetic twinning called albite twinning. The atomic arrangements in alternating domains are slightly different and are reflections of each other. See the drawing of albite twinning in Figure 4. Twinning occurs in all kinds of feldspars but is generally not visible with the naked eye.

So, although a feldspar crystal appears homogeneous, it may be composed of two or more twin domains. We most easily see feldspar twinning when using a polarizing microscope to view a feldspar in a thin section. The different domains are visible because they have different optical properties. Feldspars twin in several different ways that have different names, for example Carlsbad, Baveno, or Manebach twinning, shown in the three photos above Figures 6.

The crystals seen in these figures are about 4 cm, 4 cm, and 7 cm tall, respectively. These three kinds of twins are simple twins involving only two twin domains. Carlsbad twins are penetration twins common in orthoclase, and sometimes in plagioclase.

The photo above left shows Carlsbad twinning with two orthoclase crystals that appear to have grown through each other. Figures 4. Baveno twins are contact twins that occur most often in orthoclase and microcline. The photo above center shows a feldspar with a Baveno twin. A vertical composition plane separates the left and right sides of the crystal. Manebach twins are most common in orthoclase. The photo on the right, above, is a crystal with a Manebach twin.

The reentrant angle two small crystal faces forming a vee shape that points into the crystal in its top shows where the near-vertical composition plane passes through the crystal.

Both Baveno and Manebach twins are rarer than Carlsbad twins. When both are present, Carlsbad and Manebach twinning are diagnostic for orthoclase. Polysynthetic twins, such as the twins in Figure 6. Albite twinning is one kind; pericline twinning is another.

The difference between the two is that domains in albite twins are related by reflection, and domains in pericline twins are related by rotation. Albite twinning is a diagnostic property for plagioclase. Sometimes albite twinning combines with Carlsbad twinning to produce complex crystals, but it is still generally recognizable.

Microcline is the stable K-feldspar polymorph at normal Earth surface conditions. So, other K-feldspars may turn into microcline. Pericline twinning is produced during the polymorphic transformation of sanidine or orthoclase to microcline, and most microcline exhibits both pericline and albite twinning.

The combined twinning produces typical microcline twinning , sometimes called cross-hatched twinning or tartan twinning , but we can only see this twinning with a polarizing microscope. The colors are artifacts and are not the true colors of the domains. The field of view is 2 mm across, and the veins that cut across the field of view contain albite.

At high temperature, as long as they are not so hot that they melt, these feldspars can have any composition between sanidine KAlSi 3 O 8 and albite NaAlSi 3 O 8 , and may be anorthoclase part way between the end members.

However, as shown in this phase diagram Figure 6. At lower temperatures, instead of having one intermediate feldspar, anorthoclase unmixes to produce two separate feldspars, one K-rich and the other Na-rich. This happens because a miscibility gap exists between albite and orthoclase. A miscibility gap is a composition range within which no single mineral is stable under a particular set of pressure-temperature conditions.

It is the gray region shown in the phase diagram. The curve above and around the miscibility gap is called the solvus. Miscibility gaps are common in mineral systems because some elements do not mix well under all conditions. Many mineral series show complete miscibility at high temperatures, meaning that all compositions are stable.

At low temperatures, partial or complete immiscibility restricts possible compositions. We might make an analogy to a pot of homemade chicken soup that separates into two compositions fat and chicken broth with cooling Figure 6. We call the process of a single feldspar separating into two compositions exsolution , equivalent to unmixing. If an intermediate-composition alkali feldspar cools rapidly, it may not have time to exsolve.

Thus, we have examples of anorthoclase to study. On the other hand, if cooling is slow, the feldspar will unmix. This may result in separate grains of K-feldspar orthoclase or microcline and Na-feldspar albite developing in a rock. More often, it results in alternating layers or irregular zones of orthoclase and albite within individual crystals. If the layers or zones are planar or nearly so appearing long and thin in thin section , we call them exsolution lamellae.

The lamellae are sometimes visible with the naked eye, but frequently require a microscope to detect. 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 so this should give the ratio of Si to O in single-chain silicates e. The diagram below represents a double chain in a silicate mineral.

Again, count the number of tetrahedra versus the number of oxygen ions. This should give you the ratio of Si to O in double-chain silicates e. In amphibole structures, the silica tetrahedra are linked in a double chain that has an oxygen-to-silicon ratio lower than that of pyroxene, and hence still fewer cations are necessary to balance the charge. Amphibole is even more permissive than pyroxene and its compositions can be very complex.

Hornblende, for example, can include sodium, potassium, calcium, magnesium, iron, aluminum, silicon, oxygen, fluorine, and the hydroxyl ion OH —. In mica structures, the silica tetrahedra are arranged in continuous sheets, where each tetrahedron shares three oxygen anions with adjacent tetrahedra. There is even more sharing of oxygens between adjacent tetrahedra and hence fewer charge-balancing cations are needed for sheet silicate minerals.

Bonding between sheets is relatively weak, and this accounts for the well-developed one-directional cleavage Figure 2. Chlorite is another similar mineral that commonly includes magnesium. In muscovite mica, the only cations present are aluminum and potassium; hence it is a non-ferromagnesian silicate mineral. Apart from muscovite, biotite, and chlorite, there are many other sheet silicates or phyllosilicates , which usually exist as clay-sized fragments i. These include the clay minerals kaolinite , illite, and smectite , and although they are difficult to study because of their very small size, they are extremely important components of rocks and especially of soils.

Silica tetrahedra are bonded in three-dimensional frameworks in both the feldspars and quartz. In addition to silica tetrahedra, feldspars include the cations aluminum, potassium, sodium, and calcium in various combinations.

Coal is made of plant and animal remains. Is it a mineral? Coal is a classified as a sedimentary rock but is not a mineral.

Minerals are made by natural processes, those that occur in or on Earth. Is a diamond created in a laboratory by placing carbon under high pressures a mineral? Nearly all All minerals have a specific chemical composition. The mineral silver is made up of only silver atoms and diamond is made only of carbon atoms, but most minerals are made up of chemical compounds.

Each mineral has its own chemical formula. Halite, pictured in the Figure 2. Quartz is always made of two oxygen atoms bonded to a silicon atom, SiO 2. A hard mineral containing covalently bonded carbon is diamond, but a softer mineral that also contains calcium and oxygen along with carbon is calcite Figure below. Some minerals have a range of chemical composition. Olivine always has silicon and oxygen as well as iron or magnesium or both, Mg, Fe 2 SiO 4.

How physical properties are used to identify minerals is described in the lesson on Mineral Formation. Minerals are divided into groups based on chemical composition. Most minerals fit into one of eight mineral groups. Silicates are by far the largest mineral group. Feldspar and quartz are the two most common silicate minerals. Both are extremely common rock-forming minerals.

The basic building block for all silicate minerals is the silica tetrahedron, which is illustrated in Figure below. To create the wide variety of silicate minerals, this pyramid-shaped structure is often bound to other elements, such as calcium, iron, and magnesium.

Silica tetrahedrons combine together in six different ways to create different types of silicates Figure below. Tetrahedrons can stand alone, form connected circles called rings, link into single and double chains, form large flat sheets of pyramids, or join in three dimensions. The different ways that silica tetrahedrons can join together cause these two minerals to look very different.

Native elements contain atoms of only one type of element. Only a small number of minerals are found in this category. Some of the minerals in this group are rare and valuable.

Gold, silver, sulfur, and diamond are examples of native elements.