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Although color is generally the most conspicuous characteristic of any mineral, it is considered a diagnostic property of only a few minerals. Slight impurities in the common mineral quartz, for example, give it a variety of tints including pink, purple, yellow, white, gray, and even black. Other minerals, such as tourmaline, also exhibit a variety of hues, with multiple colors sometimes occurring in the same sample. thus, the use of color as a means of identification is often ambiguous or even misleading. The term tenacity describes a mineral's toughness, or its resistance to breaking or deforming. Minerals that are ionically bonded, such as fluorite and halite, tend to be brittle and shatter into small pieces when struck. By contrast, minerals with metallic bonds, such as native copper, are malleable, or easily hammered into different shapes. Minerals, including gypsum and talc, that can be cut into thin shavings are described as sectile. Still others, notably the micas, are elastic and will bend and snap back to their original shape after the stress is released. One of the most useful diagnostic properties is hardness, a measure of the resistance of a mineral to abrasion or scratching. This property is determined by rubbing a mineral of unknown hardness against one of the known hardness, or vice versa. A numerical value of hardness can be obtained by using the Mohs Scale of hardness, which consists of 10 minerals arranged in order from 1 (softest) to 10 (hardest). It should be noted that the Mohs scale is relative in ranking, and it does not imply that mineral number 2, gypsum, is twice as hard as mineral 1, talc. In fact, gypsum is only slightly harder than talc. In the laboratory, other common objects can be used to determine the hardness of a mineral. These include a human fingernail, which has a hardness of about 2.5, a copper penny 3.5, and a piece of glass 5.5. The mineral gypsum which has a hardness of 2, can be easily scratched with a fingernail. On the other hand, the mineral calcite, which has a hardness of 3, will scratch a fingernail but will not scratch glass. Quartz, one of the hardest common minerals, will easily scratch glass. diamonds, hardest of all, scratch anything, including other diamonds. In the crystal structure of many minerals, some atomic bonds are weaker than others. It is along these weak bonds that minerals tend to break when they are stressed. Cleavage (kleiben=carve) is the tendency of a mineral to break (cleave) along the planes of weak bonding. Not all minerals have cleavage, but those that do can be identified by the relatively smooth, flat surfaces that are produced when the mineral is broken. The simplest type of cleavage is exhibited by the micas. Because these minerals have very weak bonds in one direction, they cleave to form thin, flat sheets. Some minerals have excellent cleavage in one, two, three, or more directions, whereas others exhibit fair or poor cleavage, and still others have no cleavage at all. When minerals break evenly in more than one direction, cleavage is described by the number of cleavage directions and the angle(s) at which they meet. Each cleavage surface that has a different orientation is counted as a different direction of cleavage. For example, some minerals cleave to form six-sided cubes. Because cubes are defined by three different sets of parallel planes that intersect at 90 degree angles, cleavage is described as three directions of cleavage that meet at 90 degrees. Do not confuse cleavage with crystal shape. When a mineral exhibits cleavage it will break into pieces that all have the same geometry. By contrast, the smooth sided quartz crystals in figure 2.1 do not have cleavage. If broken, they fracture into shapes that do not resemble one another or the original crystals. Density, an important property of matter, is defined as mass per unit volume. Mineralogists often use a related measure called specific gravity to describe the density of minerals. Specific gravity is a number representing the ratio of a mineral's weight to the weight of an equal volume of water. Most common rock forming minerals have a specific gravity of between 2 and 3. For example, quartz has a specific gravity of 2.65. By contrast, some metallic minerals such as pyrite, native copper, and magnetite are more than twice as dense and thus have more than twice the specific gravity as quartz. Galena, an ore of lead, has a specific gravity of roughly 7.5, whereas the specific gravity of 24-karat gold is approximately 20. With a little practice, you can estimate the specific gravity of a mineral by hefting it in your hand. as yourself, does this mineral feel about as "heavy" as similar sized rocks you have handled? If the answer is "yes," the specific gravity of the sample will likely be between 2.5 and 3. In addition to the properties discussed thus far, some minerals can be recognized by other distinctive properties. For example, halite is ordinary salt, so it can be quickly identified through taste. Talc and graphite both have distinctive feels; talc feels soapy, and graphite feels greasy. Further, the streaks of many sulfur-bearing minerals emit odors like rotten eggs. A few minerals, such as magnetite, have a high iron content and can be picked up with a magnet, while some varieties (lodestone) are natural magnets and will pick up small iron based objects such as pins and paper clips. Moreover, some minerals exhibit special optical properties. For example, when a transparent piece of calcite is placed over printed text, the letters appear twice. This optical property is known as double refraction. One very simple chemical test involves placing a drop of dilute hydrochloric acid from a dropper bottle onto a freshly broken mineral surface. Using this technique, certain minerals, called carbonates, will effervesce (fizz) as carbon dioxide gas is released. This test is especially useful in identifying the common carbonate mineral calcite. One of the simplest silicate structures consists of independent tetrahedral that have their four oxygen ions bonded to positive ions, such as mg2+, Fe2+, and Ca2+. the mineral olivine, with the formula MfFe2SiO4 is a good example. In olivine, magnesium (Mg2+) and/or iron (Fe2+) ions pack between comparatively large independent SiO4 tetrahedral, forming a dense three-dimensional structure. garnet, another common silicate, is also composed of independent tetrahedral ionically bonded by positive ions. Both olivine and garnet form dense, hard, equidimensional crystals that lack cleavage. In the most common silicate structure, all four oxygen ions are shared, producing a complex three-dimensional framework. Quartz, a hard, durable mineral, has the simplest structure in which all of the oxygens are shared. Because its structure is neutral, quartz (SiO2) contains no positive ions--other than silicon. The ratio of oxygen ions to silicon ions differs in each type of silicate structure. In independent tetrahedra (SiO4) there are four oxygen ions for every silicon ion. In single chains, the oxygen-to-silicon ratio is 3:1 (SiO3), and in three-dimensional frame-works as found in quartz the ratio is 2:1 (SiO2). As more oxygen ions are shared, the percentage of silicon in the structure increases. Silicate minerals are, therefore, described as having a low or high silicon content based on their ratio of oxygen to silicon. Minerals with three-dimensional structures in which all four oxygen ions are shared have the highest silicon content. Minerals composed of independent tetrahedra have the lowest. This difference in silicon content is important, as you will see in chapter 3. Except for quartz (SiO2) the basic structure (chains, sheets, or three-dimensional frame-works) of most silicate minerals has a net negative charge. Therefore, metal ions are required to bring the overall charge into balance and to serve as the "mortar" that holds these structures together. The positive ions that most often link silicate structures are iron (Fe2+), magnesium (Mg2+), potassium (K1+), sodium (Na1+), aluminum (Al3+), and calcium (Ca2+). These positively charged ions bond with the unshared oxygen ions that occupy the corners of the silicate tetrahedra. As a general rule, the hybrid covalent bonds between silicon and oxygen are stronger than the ionic bonds that hold one silicate structure to the next. Consequently, properties such as cleavage, and to some extent hardness, are controlled by the nature of the silicate framework. Quartz (SiO2), which has only silicon-oxygen bonds, has great hardness and lacks cleavage, mainly because of equally strong bonds in all directions. By contrast, the mineral talc (the source of talcum powder), has a sheet structure. Magnesium ions occur between the sheets and weakly join them together. The slippery feel of talcum powder is due to the silicate sheets sliding relative to one another, in much the same way sheets of carbon atoms slide in graphite, giving it its lubricating properties. Recall that atoms of similar size can substitute freely for one another without altering a mineral's structure. For example, in the mineral olivine, iron (Fe2+) and magnesium (Mg2+) substitute for each other. This also holds true for the third most common element in earth's crust, aluminum (Al3+), which often substitutes for silicon (Si) in the center of silicon-oxygen tetrahedra. Because most silicates will readily accommodate two or more different positive ions at a given bonding site, individual specimens of a particular mineral may contain varying amounts of certain elements. As a result, many silicate minerals form a mineral group that exhibits a range of compositions between tow end members. Examples include the olivine's, pyroxenes, amphiboles, micas, and feldspars. The feldspars are, by far, the most plentiful silicate group, comprising more than 50 percent of earth's crust. Quartz, the second most abundant mineral in the continental crust, is the only common mineral made completely of silicon and oxygen. Most silicate minerals form when molten rock cools and crystallizes. Cooling can occur at or near earth's surface (low temperature and pressure) or at great depths (high temperature and pressure). The environment during crystallization and the chemical composition of the molten rock determine, to a large degree, which minerals are produced. For example, the silicate mineral olivine crystallizes at high temperatures, whereas quartz crystallizes at much lower temperatures. In addition, some silicate minerals form at earth's surface from the weathered products of other silicate minerals. Still others are formed under the extreme pressures associated with mountain building. Each silicate mineral, therefore, has a structure and a chemical composition that indicate the conditions under which it formed. by carefully examining the mineral constituents of rocks, geologists can usually determine the circumstances under which the rocks formed. Feldspar, the most common mineral group, can form under a wide range of temperatures and pressures, which partially accounts for its abundance. All feldspars have similar physical properties. They have two planes of cleavage meeting at or near 90-degree angles, are relatively hard (6 on the Mohs scale), and have a luster that range from glassy to pearly. As one component in a rock, feldspar crystals can be identified by their rectangular shape and rather smooth shiny faces. Two different feldspar structures exist. One group of feldspar minerals contains potassium ions in its structure and is therefore referred to as potassium feldspar. (Orthoclase and microcline are common members of the potassium feldspar group.) The other group, called plagioclase feldspar, contains both sodium and calcium ions that freely substitute for one another depending on the environment during crystallization. Potassium feldspar is usually light cream, salmon pink, or occasionally bluish green in color. The plagioclase feldspars, on the other hand, range in color from white to medium gray. However, color should not be used to distinguish these groups. The only way to distinguish the feldspars physically is to look for a multitude of fine parallel lines, called striations. Striations are found on some cleavage planes of plagioclase feldspar but are not present on potassium feldspar. Quartz is the only common silicate mineral consisting entirely of silicon and oxygen. As such, the term silica is applied to quartz, which has the chemical formula SiO2. Because the structure of quartz contains a ratio of two oxygen ions (O2-) for every silicon ion (Si4+), no other positive ions are needed to attain neutrality. In quartz, a three-dimensional framework is developed through the compete sharing of oxygen by adjacent silicon atoms. Thus, all of the bonds in quartz are of the strong silicon-oxygen type. Consequently, quartz is hard, resistant to weathering, and does not have cleavage. When broken, quartz generally exhibits conchoidal fracture. When pure, quartz is clear and, if allowed to grow without interference, will develop hexagonal crystals that develop pyramid-shaped ends. However, like most other clear minerals, quartz is often colored by inclusions of various ions (impurities) and forms without developing good crystal faces. The most common varieties of quartz are milky (white), smoky (gray), rose (pink), amethyst (purple), and rock crystal (clear). Clay is a term used to describe a category of complex minerals that, like the micas, have a sheet structure. Unlike other common silicates, such as quartz and feldspar, most clay minerals originate as products of the chemical weathering of other silicate minerals. Thus, clay minerals make up a large percentage of the surface material we call soil. Because of the importance of soil in agriculture, and because of its role as a supporting material for buildings, clay minerals are extremely important to humans. In addition, clays account for nearly half the volume of sedimentary rocks. clay minerals are generally very fine grained, which makes identification difficult, unless studied microscopically. Their layered structure and weak bonding between layers give them a characteristic feel when wet. Clays are common in shales, mudstones, and other sedimentary rocks. One of the most common clay minerals is kaolinite, which is used in the manufacture of fine china and as a coating for high gloss paper, such as that used in this textbook. further, some clay minerals absorb large amounts of water, which allows them to swell to several times their normal size. These clays have been used commercially in a variety of ingenious ways, including as an additive to thicken milkshakes in fast-food restaurants. Nonsilicate minerals are typically divided into groups, based on the negatively charged ion or complex ion that the members have in common. For example, the oxides contain the negative oxygen ion (O2-), which is bonded to one or more kinds of positive ions. Thus, within each mineral group, the basic structure and tyupe of bonding is similar. as a result, the minerals in each group have similar physical properties that are useful in mineral identification. Although the nonsilicates make up only about 8 percent of earth's crust, some minerals, such as gypsum, calcite, and halite, occur as constituents in sedimentary rocks in significant amounts. Furthermore, many others are important economically. Some of the most common nonsilicate minerals belong to one of three classes of minerals--the carbonates (CO32-), the sulfates (SO42-), and the halides (Cl1-,F1-,Br1-). The carbonate minerals are much simpler structurally than the silicates. This mineral group is composed of the carbonate ion (CO32-) and one or more kinds of positive ions. The two most common carbonate minerals are calcite, CaCO3 (Calcium carbonate), and dolomite, CaMg (CO3)2 (calcium/magnesium carbonate). Because these minerals are similar both physically and chemically, they are difficult to distinguish from each other. Both have a vitreous luster, a hardness between 3 and 4, and nearly perfect rhombic cleavage. They can, however, be distinguished by using dilute hydrochloric acid. Calcite reacts vigorously with this acid, whereas dolomite reacts much more slowly. Calcite and dolomite are usually found together as the primary constituents in the sedimentary rocks limestone and dolostone. When calcite is the dominant mineral, the rock is called limestone, whereas dolostone results from a predominance of dolomite. Limestone has many uses, including as road aggregate, as building stone, and as the man ingredient in Portland cement. Two other nonsilicate minerals frequently found in sedimentary rocks are halite and gypsum. Both minerals are commonly found in thick layers that are the last vestiges of ancient seas that have long since evaporated. Like limestone, both are important nonmetallic resources. Halite is the mineral name for common table salt (NaCl). Gypsum *CaSO4.2H2O), which is calcium sulfate with water bound into the structure, is the mineral of which plaster and other similar building materials are composed. Most nonsilicate mineral classes contain members that are prized for their economic value. This includes the oxides whose members hematite and magnetite are important ores of iron. Also significant are the sulfides, which are basically compounds of sulfur (S) and one or more metals. Examples of important sulfide minerals include galena (lead), sphalerite (zinc), and chalcopyrite (copper). In addition, native elements, including gold, silver, and carbon (diamonds), plus a host of other nonsilicate minerals--fluorite (flux in making steel), corundum (gemstone, abrasive), and uraninite (a uranium source)---are important economically. Useful metallic minerals that can be mined at a profit. In common usage, the term ore is also applied to some non-metallic minerals such as fluorite and sulfur. However, materials used for such purposes as building stone road aggregate, abrasives, ceramics, and fertilizers are not usually called ores; rather, they are classified as industrial rocks and minerals. Recall that more than 98 percent of earth's crust is compsed of only eight elements, and except for oxygen and silicon, all other elements make up a relatively small fraction of common crustal rocks. Indeed, the natural concentrations of many elements are exceedingly small. A deposit containing the average percentage of a valuable element such as gold has no economic value, because the cost of extracting it greatly exceeds the value of the gold that could be recovered. To have economic value, an element must be concentrated above the level of its average crustal abundance. For example, copper makes up about 0.0135 percent of the crust. For a deposit to be considered as copper ore, it must contain a concentration that is about 100 times this amount. Aluminum, on the other hand, represents 8.13 percent of the crust and can be extracted profitably when it is found in concentrations only about four times its average crustal percentage. It is important to realize that a deposit may become profitable to extract or lose its profitability because of economic changes. If demand for a metal increases and prices rise sufficiently, the status of a previously unprofitable deposit changes, and it becomes an ore. The status of unprofitable deposits may also change if a technological advance allows the ore to be extracted at a lower cost than before. Conversely, changing economic factors can turn a once profitable ore deposit into an unprofitable deposit that can no longer be called an ore. This situation was illustrated at the copper mining operation located at Bingham Canyon, Utah, one of the largest open-pit mines on earth. Mining was halted there in 1985 because outmoded equipment had driven the cost of extracting the copper beyond the current selling price. The owners responded by replacing an antiquated 1000-car railroad with conveyor belts and pipelines for transporting the ore and waste. These devices achieved a cost reduction of nearly 30 percent and returned this mining operation to profitability. Over the years, geologists have been keenly interested in learning how natural processes produce localized concentrations of essential minerals. One well-established fact is that occurrences of valuable mineral resources are closely related to the rock cycle. That is, the mechanisms that generate igneous, sedimentary, and metamorphic rocks, including the processes of weathering and erosion, play a major role in producing concentrated accumulations of useful elements. Moreover, with the development of the theory of plate tectonics, geologists have added another tool for understanding the processes by which one rock is transformed into another. As these rock-forming processes are examined in the following chapters, we will consider their role in producing some of our important mineral resources. 