What Is Rock?

What Is Rock?

kacimo samy   7:59 ص

What Is Rock? 

To geologists, rock is a coherent, naturally occurring solid, consisting of an aggregate of minerals or, less commonly, of glass. Let’s take this definition apart to see what its components mean. 

• Coherent: A rock holds together, and thus must be broken to be separated into pieces. As a result of its coherence, rock can form cliff or can be carved into sculptures. A pile of unattached mineral grains does not constitute a rock. 
• Naturally occurring: Geologists consider only naturally occurring materials to be rocks, so manufactured materials, such as concrete and brick, do not qualify. 
• An aggregate of minerals or a mass of glass: The vast majority of rocks consist of an aggregate (a collection) of many mineral grains, and/or crystals, stuck or grown together. Some rocks contain only one kind of mineral, whereas others contain several different kinds. A few rock types consist of glass. 

Rocks, aggregates of mineral grains and/or crystals, can be clastic or crystalline.

What holds rock together? Grains in rock stick together to form a coherent mass either because they are bonded by  natural cement, mineral material that precipitates from water and fills the space between grains (figure above a), or because they i nterlock with one another like pieces in a jigsaw puzzle (figure above b). Rocks whose grains are stuck together by cement are called  clastic, whereas rocks whose crystals interlock with one another are called crystalline. Glassy rocks hold together because they originate as a continuous mass (that is, they have no separate grains), because glassy grains were welded together while still hot, or because they were cemented together at a later time. 

Types of rock exposures.

At the surface of the Earth, rock occurs either as broken chunks (pebbles, cobbles, or boulders) that have moved by falling down a slope or by being transported in ice, water, or wind, or as bedrock that is still attached to the Earth’s crust. Geologists refer to an exposure of bedrock as an outcrop. An outcrop may appear as a rounded knob out in a field, as a ledge forming a cliff or ridge, on the face of a stream cut (where running water dug down into bedrock), or along human-made roadcuts and excavations (figure above a–d). 

To people who live in cities or forests or on farmland, outcrops of bedrock may be unfamiliar, since bedrock may be completely covered by vegetation, sand, mud, gravel, soil, water, asphalt, concrete, or buildings. Outcrops are particularly rare in regions such as the midwestern United States, where, during the past million years, ice-age glaciers melted and buried  bedrock under thick deposits of debris. 

Credits: Stephen Marshak (Essentials of Geology)

Something Precious, Gemstones

Something Precious, Gemstones

kacimo samy   7:58 ص

Something Precious, Gemstones

The Hope Diamond.

Mystery and romance follow famous gems. Consider the stone now known as the Hope Diamond, recognized by name the world over (figure above). No one knows who first dug it out of the ground (See Where Do Diamonds Come From?). Was it mined in the 1600s, or was it stolen off an ancient religious monument? What we do know is that in the 1600s, a French trader named Jean Baptiste Tavernier obtained a large (112.5 carats, where 1 carat 200 milligrams), rare blue diamond in India, perhaps from a Hindu statue, and carried it back to France. King Louis XIV bought the diamond and had it fashioned into a jewel of 68 carats. This jewel vanished in 1762 during a burglary. Perhaps it was lost forever perhaps not. In 1830, a 44.5-carat blue diamond mysteriously appeared on the jewel market for sale.

Henry Hope, a British banker, purchased the stone, which then became known as the Hope Diamond. It changed hands several times until 1958, when a famous New York jeweller named Harry Winston donated it to the Smithsonian Institution in Washington, DC, where it now sits behind bulletproof glass in a heavily guarded display. 

What makes stones such as the Hope Diamond so special that people risk life and fortune to obtain them? What is the difference between a gemstone, a gem, and any other mineral? A gemstone is a mineral that has special value because it is rare and people consider it beautiful. A gem, or jewel, is a finished stone ready to be set in jewellery. Jewellers distinguish between precious stones (such as diamond, ruby, sapphire, and emerald), which are particularly rare and expensive, and semiprecious stones (such as topaz, tourmaline, aquamarine, and garnet), which are less rare and less expensive. All the stones mentioned so far are transparent crystals, though most have some colour. The category 

of semiprecious stones also includes opaque or translucent minerals such as lapis, malachite, and opal. 

In everyday language, pearls and amber may also be considered gemstones. Unlike diamonds and garnets, which form inorganically in rocks, pearls form in living oysters when the oyster extracts calcium and carbonate ions from water and precipitates them around an impurity, such as a sand grain, embedded in its body. Thus, pearls are a result of biomineralization. Most pearls used in jewellery today are “cultured” pearls, made by artificially introducing round sand grains into oysters in order to stimulate pearl production. Amber is also formed by organic processes it consists of fossilized tree sap. But because amber consists of organic compounds that are not arranged in a crystal structure, it does not meet the definition of a mineral.

Cutting gemstones.

In some cases, gemstones are merely pretty and rare versions of more common minerals. For example, ruby is a special version of the common mineral corundum, and emerald is a special version of the common mineral beryl (figure above a). As for the beauty of a gemstone, this quality lies basically in its colour and, in the case of transparent gems, its “fire” the way the mineral bends and internally reflects the light passing through it, and disperses the light into a spectrum. Fire makes a diamond sparkle more than a similarly cut piece of glass. 

Gemstones form in many ways. Some solidify from a melt, some form by diffusion, some precipitate out of a water solution in cracks, and some are a consequence of the chemical interaction of rock with water near the Earth’s surface. Many gems come from pegmatites, particularly coarse-grained rocks formed by the solidification of steamy melt. 

Most gems used in jewellery are “cut” stones, meaning that they are not raw crystals right from the ground, but rather have been faceted. The smooth facets on a gem are ground and polished surfaces made with a faceting machine (figure above b). Facets are not the natural crystal faces of the mineral, nor are they cleavage planes, though gem cutters sometimes make the facets parallel to cleavage directions and will try to break a large gemstone into smaller pieces by splitting it on a cleavage plane. A faceting machine consists of a doping arm, a device that holds a stone in a specific orientation, and a lap, a rotating disk covered with a wet paste of grinding powder and water. The gem cutter fixes a gemstone to the end of the doping arm and positions the arm so that it holds the stone against the moving lap. The movement of the lap grinds a facet. When the facet is complete, the gem cutter rotates the arm by a specific angle, lowers the stone, and grinds another facet. The geometry of the facets defines the cut of the stone. Different cuts have names, such as “brilliant,” “French,” “star,” and “pear.” Grinding facets is a lot of work a typical engagement-ring diamond with a brilliant cut has 57 facets (figure above c).

Where Do Diamonds Come From? 

Diamonds consist of carbon, which typically accumulates only at or near Earth’s surface. Experiments demonstrate that the pressures needed to form diamond are so extreme that, in nature, they generally occur only at depths of around 150 km below the Earth’s surface. Nowadays, engineers can duplicate these conditions in the laboratory, so corporations manufacture several tons of synthetic diamonds a year. 

Diamond occurrences.

How does carbon get down to depths of 150 km? Geologists speculate that  subduction or collision carries carbon- containing rocks and sediments down to  the depth where it transforms into diamond  beneath continents. But if diamonds form at great depth, how do they return to the surface? Some diamonds rise when rifting cracks the continental crust and causes a small part of the underlying mantle to melt. Magma generated during this process rises to the surface, bringing the diamonds with it. Near the surface, the magma solidifies to form an igneous rock called kimberlite, named for Kimberley, South Africa. Diamonds brought up with the magma are embedded as crystals in solid kimberlite (figure above).  Much of the world’s diamond supply comes from mines in this rock.  But some sources occur in deposits of sediment formed from the breakdown and erosion of kimberlite that had been exposed at the surface. Rivers and glaciers may transport diamond- bearing sediments far from their original bedrock source.

Not all natural diamonds are valuable; the value depends on colour and clarity. Diamonds that contain imperfections (cracks, or specks of other material), or are dark gray in colour, are not used for jewellery. These stones, called industrial diamonds, are used as abrasives. Gem-quality diamonds come in a range of sizes.  Jewellers measure the size of these gems in carats, 

where one carat equals 200 milligrams  (0.2 gram). In English units of measurement, one ounce equals 142 carats. The largest diamond ever found, a stone called the Cullinan Diamond, was discovered in South Africa in 1905, and weighed 3,106 carats (621 grams) before being cut. By comparison, the diamond on a typical engagement ring weighs less than one carat. Gem-quality diamonds are actually more common than you might expect suppliers stockpile the stones in order to avoid flooding the market and lowering the price. 

Credits: Stephen Marshak (Essentials of Geology)

Mineral Classification

Mineral Classification

kacimo samy   2:24 ص

Mineral Classification 

The 4,000 known minerals can be separated into a small number of groups, or mineral classes. You may think, “Why bother?” Classification schemes are useful because they help organize information and streamline discussion. Biologists, for example, classify animals into groups based on how they feed their young and on the architecture of their skeletons, and botanists classify plants according to the way they reproduce and by the shape of their leaves. In the case of minerals, a good means of classification eluded researchers until it became possible to determine the chemical makeup of minerals. A Swedish chemist, Baron Jöns Jacob Berzelius (1779–1848), analyzed minerals and noted chemical similarities among many of them. Berzelius, along with his students, established that most minerals can be classified by specifying the principal anion (negative ion) or anionic group (negative molecule) within the mineral. We now take a look at principal mineral classes, focusing especially on silicates, the class that constitutes most of the rock in the Earth.

The Mineral Classes 

Mineralogists distinguish several principal classes of minerals. Here are some of the major ones.

Physical characteristics of minerals.

• Silicates: The fundamental component of most silicates in the Earth’s crust is the SiO44– anionic group. A well-known example, quartz (figure above a), has the formula SiO2. 
• Oxides: Oxides consist of metal cations bonded to oxygen anions. Typical oxide minerals include hematite (Fe2O3; figure above b) and magnetite (Fe3O4; figure above g). 
• Sulfides: Sulfides consist of a metal cation bonded to a sulfide anion (S2–). Examples include galena (PbS) and pyrite (FeS2; figure above c). 
• Sulfates: Sulfates consist of a metal cation bonded to the SO42– anionic group. Many sulfates form by precipitation out of water at or near the Earth’s surface. An example is gypsum (CaSO4s (2O).
• Halides: The anion in a halide is a halogen ion (such as chloride [Cl–] or fluoride [F –]), an element from the second column from the right in the periodic table (see Appendix). Halite, or rock salt (NaCl; Fig. 3.8d), and fluorite (CaF2), a source of fluoride, are common examples. 
• Carbonates: In carbonates, the molecule CO32– serves as the anionic group. Elements such as calcium or magnesium bond to this group. The two most common carbonates are calcite (CaCO3; Fig. 3.8e) and dolomite (CaMg[CO3]2). 
• Native metals: Native metals consist of pure masses of a single metal. The metal atoms are bonded by metallic bonds. Copper and gold, for example, may occur as native metals. 

The nature of mineral cleavage and fracture.

Silicates: The Major Rock-Forming Minerals 

 The structure of silicate minerals.

Silicate minerals, or silicates, make up over 95% of the continental crust and almost 100% of the oceanic crust and of the Earth’s mantle consist almost entirely of silicates. Thus, silicates are the most common minerals on Earth. As we've noted, silicates in the Earth’s crust and upper mantle contain the SiO44– anionic group. In this group, four oxygen atoms surround a single silicon atom, thereby defining the corners of a tetrahedron, a pyramid-like shape with four triangular faces (figure above a). We refer to this anionic group as the silicon oxygen tetrahedron (or, informally, as the silica tetrahedron), and it acts, in effect, as the building block of silicate minerals. 

Mineralogists distinguish among several groups of silicate minerals based on the way in which silica tetrahedra are arranged (figure above b). The arrangement, in turn, determines the degree to which tetrahedra share oxygen atoms. Note that the number of shared oxygen determines the ratio of silicon (Si) to oxygen (O) in the mineral. Here are the groups, in order from fewer shared oxygen to more shared oxygen: 

• Independent tetrahedra: In this group, the tetrahedra are independent and do not share any oxygen atoms. The attraction between the tetrahedra and positive ions holds such minerals together. This group includes olivine, a glassy green mineral, and garnet ( 2nd figure above f). 
• Single chains: In a single-chain silicate, the tetrahedra link to form a chain by sharing two oxygen atoms. The most common of the many different types of single-chain silicates are pyroxenes (2nd figure above b). 
• Double chains: In a double-chain silicate, the tetrahedra link to form a double chain by sharing two or three oxygen atoms. Amphiboles are the most common type (2nd figure above c). 
• Sheet silicates: The tetrahedra in this group share three oxygen atoms and therefore link to form two- dimensional sheets. Other ions and, in some cases, water molecules fit between the sheets in some sheet silicates. Because of their structure, sheet silicates have cleavage in one direction, and they occur in books of very thin sheets. In this group, we find micas (2nd figure above a)  and clays. Clays occur only in extremely tiny flakes. 
• Framework silicates: In a framework silicate, each tetrahedron shares all four oxygen atoms with its neighbours, forming a three-dimensional structure. Examples include feldspar and quartz. The two most common feldspars are plagioclase, which tends to be white, gray, or blue; and orthoclase (also called potassium feldspar, or K-feldspar), which tends to be pink (1st figure above d). 

Credits: Stephen Marshak (Essentials of Geology)

How Can You Tell One Mineral From Another?

How Can You Tell One Mineral From Another?

kacimo samy   7:19 ص

How Can You Tell One Mineral From Another? 

Amateur and professional mineralogists get a kick out of recognizing minerals. They might hover around a display case in a museum and name specimens without bothering to look at the labels. How do they do it? The trick lies in learning to recognize the basic physical properties (visual and material characteristics) that distinguish one mineral from another. Some physical properties, such as shape and colour, can be seen from a distance. Others, such as hardness and magnetization, can be determined only by handling the specimen or by performing an identification test on it. Identification tests include scratching the mineral by another object, placing it near a magnet, weighing it, tasting it, or placing a drop of acid on it. Let’s examine some of the physical properties most commonly used in basic mineral identification.

Physical characteristics of minerals.

• Colour: Colour results from the way a mineral interacts with light. Sunlight contains the whole spectrum of colours; each colour has a different wavelength. A mineral absorbs certain wavelengths, so the colour you see when looking at a specimen represents the wavelengths the mineral does not absorb.  Certain minerals always have the same colour, but many display a range of colours (figure above a). Colour variations in a mineral are due to the presence of impurities. For example, trace amounts of iron may give quartz a reddish colour. 
• Streak: The streak of a mineral refers to the colour of a powder produced by pulverizing the mineral. You can obtain a streak by scraping the mineral against an unglazed ceramic plate (figure above b). The colour of a mineral powder tends to be less variable than the colour of a whole crystal, and thus  provides a fairly reliable clue to a mineral’s identity. Calcite, for example, always yields a white streak even though pieces of calcite may be white, pink, or clear. 
• Luster: Luster refers to the way a mineral surface scatters light. Geoscientists describe luster by comparing the appearance of the mineral with the appearance of a familiar substance. For example, minerals that look like metal have a metallic luster, whereas those that do not have a nonmetallic luster the adjectives are self-explanatory  (figure above c, d). Terms used for types of nonmetallic luster include silky, glassy, satiny, resinous, pearly, or earthy. 
• Hardness: Hardness is a measure of the relative ability of a mineral to resist scratching, and it therefore represents the resistance of bonds in the crystal structure to being broken. The atoms or ions in crystals of a hard mineral are more strongly bonded than those in a soft mineral. Hard minerals can scratch soft minerals, but soft minerals cannot scratch hard ones. Diamond, the hardest mineral known, can scratch most anything, which is why it is used to cut glass. In the early 1800s, a mineralogist named Friedrich Mohs listed some minerals in sequence of relative hardness; a mineral with a hardness of 5 can scratch all minerals with a hardness of 5 or less. This list, the Mohs hardness scale, helps in mineral identification. To make the scale easy to use, common items such as your fingernail, a penny, or a glass plate have been added (Table below). 
• Specific gravity: Specific gravity represents the density of a mineral, as represented by the ratio between the weight of a volume of the mineral and the weight of an equal volume of water at 4°C. For example, one cubic centimetre of quartz has a weight of 2.65 grams, whereas one cubic centimetre of water has a weight of 1.00 gram. Thus, the specific gravity of quartz is 2.65. In practice, you can develop a “feel” for specific gravity by hefting minerals in your hands. A piece of galena (lead ore) feels heavier than a similar-sized piece of quartz. 
• Crystal habit: The crystal habit of a mineral refers to the shape of a single crystal with well-formed crystal faces, or to the character of an aggregate of many well-formed crystals that grew together as a group (figure above e). The habit depends on the internal arrangement of atoms in the crystal.  A description of habit generally includes adjectives that highlight the shape of the crystal. For example, crystals that are roughly the same length in all directions are called equant or blocky, those that are much longer in one dimension than in others are columnar or needle-like, those shaped like  sheets of paper are platy, and those shaped like knives are bladed. 
• Special properties: Some minerals have distinctive properties that readily distinguish them from other minerals. For example, calcite (CaCO3) reacts with dilute hydrochloric acid (HCl) to produce carbon dioxide (CO2) gas (figure above f). Dolomite (CaMg[CO3]2) also reacts with acid, but not as strongly. Graphite makes a gray mark on paper, magnetite attracts a magnet (figure above g), halite tastes salty, and plagioclase has striations (thin parallel corrugations or stripes) on its surface.
• Fracture and cleavage: Different minerals fracture (break) in different ways, depending on the internal arrangement of atoms. If a mineral breaks to form distinct planar  surfaces that have a specific orientation in relation to the crystal structure, then we say that the mineral  has cleavage and we refer to each surface as a cleavage plane. Cleavage forms in directions where the bonds holding atoms together in the crystal are the weakest (figure below a–e). Some minerals have one direction of cleavage. For example, mica has very weak bonds in one direction but strong bonds in the other two directions. Thus, it easily splits into parallel sheets; the surface of each sheet is a cleavage plane. Other minerals have two or three directions of cleavage that intersect at a specific angle. For example, halite has three sets of cleavage planes that intersect at right angles, so halite crystals break into little cubes. Materials that have no cleavage at all (because bonding is equally strong in all directions) break either by forming irregular fractures or by forming conchoidal fractures (figure below f). Conchoidal fractures are smoothly curving, clamshell-shaped surfaces; they typically form in glass. Cleavage planes are sometimes hard to distinguish from crystal faces (figure below g).

 Mohs hardness scale. Mohs’ numbers are relative in reality, diamond is 3.5 times harder than corundum, as the graph shows.

The nature of mineral cleavage and fracture.

Credits: Stephen Marshak (Essentials of Geology)

Crystals and Their Structure

Crystals and Their Structure

kacimo samy   6:50 ص

Crystals and Their Structure 

What Is a Crystal? 

Some characteristics of crystals.

The word crystal brings to mind sparkling chandeliers, elegant wine goblets, and shiny jewels. But, as is the case with the word mineral, geologists have a more precise definition. A crystal is a single, continuous (that is, uninterrupted) piece of a crystalline solid, typically bounded by flat surfaces, called crystal faces, that grow naturally as the mineral forms. The word comes from the Greek krystallos, meaning ice. Many crystals have beautiful shapes that look like they belong in the pages of a geometry book. The angle between two adjacent crystal faces of one specimen is identical to the angle between the corresponding faces of another specimen. For example, a perfectly formed quartz crystal looks like an obelisk  (figure above a, b); the angle between the faces of the columnar part of a quartz crystal is always exactly 120°. This rule, discovered by one of the first geologists, Nicolas Steno (1638– 1686) of Denmark, holds regardless of whether the whole crystal is big or small and regardless of whether all of the faces are the same size. Crystals come in a great variety of shapes, including cubes, trapezoids, pyramids, octahedrons, hexagonal columns, blades, needles, columns, and obelisks (figure above c).

Because crystals have a regular geometric form, people have always considered them to be special, perhaps even a source of magical powers. For example, shamans of some cultures relied on talismans or amulets made of crystals, which supposedly brought power to their wearer or warded off evil spirits. Scientists have concluded, however, that crystals have no effect on health or mood. For millennia, crystals have inspired awe because of the way they sparkle, but such behaviour is simply a consequence of how crystal structures interact with light.

Looking Inside Crystals 

Patterns and symmetry in minerals.

What makes crystals have regular geometric forms? This problem was the focus of study for centuries. An answer finally came from the work of a German physicist, Max von Laue, in 1912. He showed that an X-ray beam passing through a crystal breaks up into many tiny beams to create a pattern of dots on a screen (figure above a). Physicists refer to this phenomenon as diffraction; it occurs when waves interact with regularly spaced objects whose spacing is close to the wavelength of the waves you can see diffraction of ocean waves when they pass through gaps in a seawall. Von Laue concluded that, for a crystal to cause diffraction, atoms within it must be regularly spaced and the spacing must be comparable to the wavelength of X-rays. Eventually, Von Laue and others learned how to use X-ray diffraction patterns as a basis for defining the specific arrangement of atoms in crystals. This arrangement defines the crystal structure of a mineral. 

If you've ever examined wallpaper, you've seen an example of a pattern (figure above b). Crystal structures contain one of nature’s most spectacular examples of such a pattern. In crystals, the 

pattern is defined by the regular spacing of atoms and, if the crystal contains more than one element, by the regular alternation of atoms (figure above c). (Mineralogists refer to a 3-D geometry of points representing this pattern as a lattice.) The pattern of atoms in a crystal may control the shape of a crystal. For example, if atoms in a crystal pack into the shape of a cube, the crystal may have faces that intersect at 90° angles galena  (PbS) and halite (NaCl) have such a cubic shape. Because of the pattern of atoms in a crystal structure, the structure has symmetry, meaning that the shape of one part of the structure is the mirror image of the shape of a neighbouring part. For example, if you were to cut a halite crystal or a water crystal (snowflake) in half, and place the half against a mirror, it would look whole again (figure above d). 

The nature of crystalline structure in minerals. The arrangement of atoms can be portrayed by a ball and stick model, or by a packed ball model.

To illustrate crystal structures, we look at a few examples. Halite (rock salt) consists of oppositely charged ions that stick together because opposite charges attract. In halite, six chloride (Cl–) ions surround each sodium (Na+) ion, producing an overall arrangement of atoms that defines the shape of a cube (figure above a, b). Diamond, by contrast, is a mineral made entirely of carbon. In diamond, each atom bonds to four neighbours arranged in the form of a tetrahedron; some naturally formed diamond crystals have the shape of a double tetrahedron  (figure above c). Graphite, another mineral composed entirely of carbon, behaves very differently from diamond. In contrast to diamond, graphite is so soft that we use it as the “lead” in a pencil; when a pencil moves across paper, tiny flakes of graphite peel off the pencil point and adhere to the paper.  This behaviour occurs because the carbon atoms in graphite  are not arranged in tetrahedra, but rather occur in sheets  (figure above d). The sheets are bonded to each other by weak bonds and thus can separate from each other easily. Of note, two different minerals (such as diamond and graphite) that have the same composition but different crystal structures are polymorphs.

The Formation and Destruction of Minerals 

New mineral crystals can form in five ways. First, they can form by the solidification of a melt, meaning the freezing of a liquid to form a solid. For example, ice crystals, a type of mineral, are made by solidifying water, and many different minerals form by solidifying molten rock. Second, they can form by precipitation from a solution, meaning that atoms, molecules, or ions dissolved in water bond together and separate out of the water. Salt crystals, for example, precipitate when you evaporate salt water. Third, they can form by solid-state diffusion, the movement of atoms or ions through a solid to arrange into a new crystal structure, a process that takes place very slowly. For example, garnets  grow by diffusion in solid rock. Fourth, minerals can form at interfaces between the physical and biological components of the Earth System by a process called biomineralization. This occurs when living organisms cause minerals to precipitate either within or on their bodies, or immediately adjacent to their bodies. For example, clams and other shelled organisms extract ions from water to produce mineral shells. Fifth, minerals can precipitate directly from a gas. This process typically occurs around volcanic vents or around geysers, for at such locations volcanic gases or steam enter the atmosphere and cool abruptly. Some of the bright yellow sulphur deposits found in volcanic regions form in this way.

The growth of crystals.

The first step in forming a crystal is the chance formation of a seed, or an extremely small crystal (figure above a). Once the seed exists, other atoms in the surrounding material attach themselves to the face of the seed. As the crystal grows, crystal faces move outward but maintain the same orientation (figure above b). The youngest part of the crystal is at its outer edge.
In the case of crystals formed by the solidification of a melt, atoms begin to attach to the seed when the melt becomes so cool that thermal vibrations can no longer break apart the attraction between the seed and the atoms in the melt. Crystals formed by precipitation from a solution develop when the solution becomes saturated, meaning the number of dissolved ions per unit volume of solution becomes so great that they can get close enough to each other to bond together.
As crystals grow, they develop their particular crystal shape, based on the geometry of their internal structure. The shape is defined by the relative dimensions of the crystal (needle- like, sheet-like, etc.) and the angles between crystal faces. Typically, the growth of minerals is restricted in one or more directions, because existing crystals act as obstacles. In such cases, minerals grow to fill the space that is available, and their shape is controlled by the shape of their surroundings. Minerals without well-formed crystal faces are anhedral grains (figure above c). If a mineral’s growth is unimpeded so that it displays well-formed crystal faces, then it is a euhedral crystal. The surface crystals of a geode, a mineral-lined cavity in rock, may be euhedral (figure above d).
A mineral can be destroyed by melting, dissolving, or some other chemical reaction. Melting involves heating a mineral to a temperature at which thermal vibration of the atoms or ions in the lattice break the chemical bonds holding them to the lattice. The atoms or ions then separate, either individually or in small groups, to move around again freely. Dissolution occurs when you immerse a mineral in a solvent, such as water. Atoms or ions then separate from the crystal face and are surrounded by solvent molecules. Chemical reactions can destroy a mineral when it comes in contact with reactive materials. For example, iron-bearing minerals react with air and water to form rust. The action of microbes in the environment can also destroy minerals. In effect, some microbes can “eat” certain minerals; the microbes use the energy stored in the chemical bonds that hold the atoms of the mineral together as their source of energy for metabolism.

Credits: Stephen Marshak (Essentials of Geology)

What Is a Mineral?

What Is a Mineral?

kacimo samy   6:16 ص

What Is a Mineral? 

To a geologist, a mineral is a naturally occurring solid, formed by geologic processes, that has a crystalline structure and a definable chemical composition. Almost all minerals are inorganic. Let’s pull apart this mouthful of a definition and examine its meaning in detail.

• Naturally occurring: True minerals are formed in nature, not in factories. We need to emphasize this point because in recent decades, industrial chemists have learned how to synthesize materials that have characteristics virtually identical to those of real minerals. These materials are not minerals in a geologic sense, though they are referred to in the  commercial world as synthetic minerals. 
• Formed by geologic processes: Traditionally, this phrase implied processes, such as solidification of molten rock or direct precipitation from a water solution, that did not involve living organisms. Increasingly, however, geologists recognize that life is an integral part of the Earth System. So, some  geologists consider solid, crystalline materials produced by organisms to be minerals too. To avoid confusion, the term “biogenic mineral” may be used when discussing such  materials. 
• Solid: A solid is a state of matter that can maintain its shape indefinitely, and thus will not conform to the shape of its container. Liquids (such as oil or water) and gases (such as air) are not minerals (Some Basic Concepts from Chemistry). 
• Crystalline structure: The atoms that make up a mineral are not distributed randomly and cannot move around easily. Rather, they are fixed in a specific, orderly pattern. A material in which atoms are fixed in an orderly pattern is called a crystalline solid. 
• Definable chemical composition: This simply means that it is possible to write a chemical formula for a mineral (Some Basic Concepts from Chemistry). Some minerals contain only one element, but most are compounds of two or more elements. For example, diamond and graphite have the formula C, because they consist entirely of carbon. Quartz has the formula SiO2 it contains the elements silicon and oxygen in the proportion of one silicon atom for every two oxygen atoms. Calcite has the formula CaCO3, meaning it consists of a calcium (Ca ) ion and a carbonate (CO3 ) ion. Some formulas are more complicated: for example, the formula for biotite is K(Mg,Fe)3(AlSi3O10)(OH)2. 
• Inorganic: Organic chemicals are molecules containing some carbon-hydrogen bonds. Sugar (C12H22O11), for example, 
• is an organic chemical. Almost all minerals are inorganic. Thus, sugar and protein are not minerals. But, we have to add the qualifier “almost all” because mineralogists do consider about 30 organic substances formed by “the action of geologic processes on organic materials” to be minerals. Examples include the crystals that grow in ancient deposits of bat guano.

The nature of crystalline and noncrystalline materials.

With these definitions in mind, we can make an important distinction between minerals and glass. Both minerals and glass are solids, in that they can retain their shape indefinitely. But a mineral is crystalline, and glass is not. Whereas atoms, ions, or molecules in a mineral are ordered into a crystal lattice, like soldiers standing in formation, those in a glass are arranged in a semi-chaotic way, like people at a party, in small clusters or chains that are neither oriented in the same way nor spaced at regular intervals (figure above a, b). 

If you ever need to figure out whether a substance is a mineral or not, just check it against the criteria listed above. Is motor oil a mineral? No it’s an organic liquid. Is table salt a mineral? Yes it’s a solid crystalline compound with the formula NaCl. Is the hard material making up the shell of an oyster considered to be a mineral? Microscopic examination of  an oyster shell reveals that  it consists of calcite, so it can be called a biogenic mineral. Is rock candy a mineral? No. Even though it is solid and crystalline, it’s made by people and it consists of sugar (an organic chemical). 

Some Basic Concepts from Chemistry 

To describe minerals, we need to use several terms from chemistry. To avoid confusion, terms are listed in an order that permits each successive term to utilize previous terms. 

Examples of states of matter and chemical bonds.

• Element: A pure substance that cannot be separated into other materials. 
• Atom: The smallest piece of an element that retains the characteristics of the element. An atom consists of a nucleus surrounded by a cloud of orbiting electrons; the nucleus is made up of protons and neutrons (except in hydrogen, whose nucleus contains only one proton and no neutrons). Electrons have a negative charge, protons have a positive charge, and neutrons have a neutral charge. An atom that has the same number of electrons as protons is said to be neutral, in that it does not have an overall electrical charge. 
• Atomic number: The number of protons in an atom of an element. 
• Atomic weight: Approximately the number of protons plus neutrons in an atom of an element. 
• Ion: An atom that is not neutral. An ion that has an excess negative charge (because it has more electrons than protons) is an anion, whereas an ion that has an excess positive charge (because it has more protons than electrons) is a cation. We indicate the charge with a superscript. For example, Cl has a single excess electron; Fe2 is missing two electrons.
• Chemical bond: An attractive force that holds two or more atoms together (figure above a–c). For example, covalent bonds form when atoms share electrons. Ionic bonds form when a cation and anion (ions with opposite charges) get close together and attract each other. In materials with metallic bonds, some of the electrons can move freely. 
• Molecule: Two or more atoms bonded together. The atoms may be of the same element or of different elements. 
• Compound: A pure substance that can be subdivided into two or more elements. The smallest piece of a compound that retains the characteristics of the compound is a molecule. 
• State of matter: The form of a substance, which reflects the degree to which the atoms or molecules comprising the matter are bonded together. figure above d–f defines three of the states solid, liquid, and gas. There are more bonds in a solid than in a liquid, and more in a liquid than in a gas. Which state exists at a given location depends on pressure and temperature, as indicated by a phase diagram (figure above g). A fourth state, plasma, exists only at very high temperatures. 
• Chemical: A general name used for a pure substance (either an element or a compound). 
• Chemical formula: A shorthand recipe that itemizes the various elements in a chemical and specifies their relative proportions. For example, the formula for water, H2O, indicates that water consists of molecules in which two hydrogens bond to one oxygen.
• Chemical reaction: A process that involves the breaking or forming of chemical bonds. Chemical reactions can break molecules apart or create new molecules and/or isolated atoms. 
• Mixture: A combination of two or more elements or compounds that can be separated without a chemical reaction. For example, a cereal composed of bran flakes and raisins is a mixture you can separate the raisins from the flakes without destroying either. 
• Solution: A type of material in which one chemical (the solute) dissolves in another (the solvent). In solutions, a solute may separate into ions during the process. For example, when salt (NaCl) dissolves in water, it separates into sodium (Na ) and chloride (Cl ) ions. In a solution, atoms or molecules of the solvent surround atoms, ions, or molecules of the solute. 
• Precipitate: A compound that forms when ions in liquid solution join together to create a solid that settles out of the solution; (verb) the process of forming solid grains by separation and settling from a solution. For example, when saltwater evaporates, solid salt crystals precipitate.

Credits: Stephen Marshak (Essentials of Geology)

What Drives Plate Motion, and How Fast Do Plates Move?

What Drives Plate Motion, and How Fast Do Plates Move?

kacimo samy   5:09 ص

What Drives Plate Motion, and How Fast Do Plates Move? 

Forces Acting on Plates 

We've now discussed the many facets of plate tectonics theory but to complete the story, we need to address a major question: “What drives plate motion?” When geoscientists first proposed plate tectonics, they thought the process occurred simply because convective flow in the asthenosphere actively dragged plates along, as if the plates were simply rafts on a flowing river. Thus, early images depicting plate motion showed simple convection cells elliptical  flow paths in the asthenosphere. At first glance, this hypothesis looked pretty good. But, on closer examination it became clear that a model of simple convection cells carrying plates on their backs can’t explain the complex geometry of plate boundaries and the great variety of plate motions that we observe on the Earth. Researchers now prefer a model in which convection, ridge push, and slab pull all contribute to driving plates. Let’s look at each of these phenomena in turn.

Convection is involved in plate motions in two ways. Recall that, at a mid-ocean ridge, hot asthenosphere rises and then cools to form oceanic lithosphere which slowly moves away from the ridge until, eventually, it sinks back into the mantle at a trench. Since the material forming the plate starts out hot, cools, and then sinks, we can view the plate itself as the top of a convection cell and plate motion as a form of convection. But in this view, convection is effectively a consequence of plate motion, not the cause. Can convection actually cause plates to move? The answer may come from studies which demonstrate that the interior of the mantle, beneath the plates, is indeed convecting on a very broad scale. Specifically, geologists have found that there are places where deeper, hotter asthenosphere is rising or upwelling, and places where shallower, colder asthenosphere is sinking or downwelling. Such asthenospheric flow probably does exert a force on the base of plates. But the pattern of upwelling and downwelling on a global scale does not match the pattern of plate boundaries exactly. So, conceivably, asthenosphere-flow may either speed up or slow down plates depending on the orientation of the flow direction relative to the movement direction of the overlying plate. 

Forces driving plate motions. Both ridge push and slab pull make plates move.

Ridge-push force develops simply because the lithosphere of mid-ocean ridges lies at a higher elevation than that of the adjacent abyssal plains (figure above a). To understand ridge-push force, imagine you have a glass containing a layer of water over a layer of honey. By tilting the glass momentarily and then returning it to its upright position, you can create a temporary slope in the boundary between these substances. While the boundary has this slope, gravity causes the weight of elevated honey to push against the glass adjacent to the side where the honey surface lies at lower elevation. The geometry of a midocean ridge resembles this situation, for sea floor of a midocean ridge is higher than sea floor of abyssal plains. Gravity causes the elevated lithosphere at the ridge axis to push on the lithosphere that lies farther from the axis, making it move away. As lithosphere moves away from the ridge axis, new hot asthenosphere rises to fill the gap. Note that the local upward movement of  asthenosphere beneath a mid-ocean ridge is a consequence of sea-floor spreading, not the cause. 

Slab-pull force, the force that subducting, downgoing plates apply to oceanic lithosphere at a convergent margin, arises simply because lithosphere that was formed more than 10 million years ago is denser than asthenosphere, so it can sink into the asthenosphere (figure above b). Thus, once an oceanic plate starts to sink, it gradually pulls the rest of the plate along behind it, like an anchor pulling down the anchor line. This “pull” is the slab-pull force.

The Velocity of Plate Motions 

Relative plate velocities: The blue arrows show the rate and direction at which the plate on one side of the boundary is moving with respect to the plate on the other side. The length of an arrow represents the velocity. Absolute plate velocities: The red arrows show the velocity of the plates with respect to a fixed point in the mantle.

How fast do plates move? It depends on your frame of reference. To illustrate this concept, imagine two cars speeding in the same direction down the highway. From the viewpoint of a tree along the side of the road, Car A zips by at  100 km an hour, while Car B moves at 80 km an hour. But relative to Car B, Car A moves at only 20 km an hour. Geologists use two different frames of reference for describing plate velocity. If we describe the movement of Plate A with respect to Plate B, then we are speaking about relative plate velocity. But if we describe the movement of both plates relative to a fixed location in the mantle below the plates, then we are speaking of absolute plate velocity (figure above). 

To determine relative plate motions, geoscientists measure the distance of a known magnetic anomaly from the axis of a mid-ocean ridge and then calculate the velocity of a plate relative to the ridge axis by applying this equation: plate velocity distance from the anomaly to the ridge axis divided by the age of the anomaly (velocity, by definition, is distance wtime). The velocity of the plate on one side of the ridge relative to the plate on the other is twice this value. 

To estimate absolute plate motions, we can assume that the position of a mantle plume does not change much for a long time. If this is so, then the track of hot-spot volcanoes on the plate moving over the plume provides a record of the plate’s absolute velocity and indicates the direction of movement. (In reality, plumes are not completely fixed; geologists use other, more complex methods to calculate absolute plate motions.) 

Working from the calculations described above, geologists have determined that plate motions on Earth today occur at rates of about 1 to 15 cm per year about the rate that your fingernails grow. But these rates, though small, can yield  large displacements given the immensity of geologic time. At a  rate of 10 cm per year, a plate can move 100 km in a million years! Can we detect such slow rates? Until the last decade, the answer was no. Now the answer is yes, because of satellites orbiting the Earth with global positioning system (GPS) technology. Automobile drivers use GPS receivers to find their destinations, and geologists use them to monitor plate motions. If we calculate carefully enough, we can detect displacements of millimeters per year. In other words, we can now see the plates move this observation serves as the ultimate proof of plate tectonics. 

Due to plate tectonics, the map of Earth‘s surface slowly changes. Here we see the assembly, and later the breakup, of Pangaea during the past 400 million years.

Taking into account many data sources that define the motion of plates, geologists have greatly refined the image of continental drift that Wegener tried so hard to prove nearly a century ago. We can now see how the map of our planet’s surface has evolved radically during the past 400 million years (figure above), and even before.

Credits: Stephen Marshak (Essentials of Geology)

How Do Plate Boundaries Form and Die?

How Do Plate Boundaries Form and Die?

kacimo samy   4:57 ص

How Do Plate Boundaries Form and Die? 

The configuration of plates and plate boundaries visible on our planet today has not existed for all of geologic history, and will not exist indefinitely into the future. Because of plate motion, oceanic plates form and are later consumed, while continents merge and later split apart. How does a new  divergent boundary come into existence, and how does an existing convergent boundary eventually cease to exist? Most new divergent boundaries form when a continent splits and separates into two continents. We call this process rifting. A convergent boundary ceases to exist when a piece of buoyant lithosphere, such as a continent or an island arc, moves into the subduction zone and, in effect, jams up the system. We call this process collision.

Continental Rifting 

During the process of rifting, lithosphere stretches.

A continental rift is a linear belt in which continental  lithosphere pulls apart (figure above a). During the process, the lithosphere stretches horizontally and thins vertically, much like a piece of taffy you pull between your fingers. Nearer the surface of the continent, where the crust is cold and brittle, stretching causes rock to break and faults to develop. Blocks of rock slip down the fault surfaces, leading to the formation of a low area that gradually becomes buried by sediment. Deeper in the crust, and in the underlying lithospheric mantle, rock is warmer and softer, so stretching takes place in a plastic manner without breaking the rock. The whole region that stretches is the rift, and the process of stretching is called rifting. 

As continental lithosphere thins, hot asthenosphere rises beneath the rift and starts to melt. Eruption of the molten rock produces volcanoes along the rift. If rifting continues for a long enough time, the continent breaks in two, a new midocean ridge forms, and sea-floor spreading begins. The relict of the rift evolves into a passive margin. In some cases, however, rifting stops before the continent splits in two; it becomes a low-lying trough that fills with sediment. Then, the rift remains as a permanent scar in the crust, defined by a belt of faults, volcanic rocks, and a thick layer of sediment. 

A major rift, known as the Basin and Range Province, breaks up the landscape of the western United States  (figure above b). Here, movement on numerous faults tilted blocks of crust to form narrow mountain ranges, while sediment that eroded from the blocks filled the adjacent basins (the low areas between the ranges). Another active rift slices through eastern Africa; geoscientists aptly refer to it as the East African Rift (figure above c, d). To astronauts in orbit, the rift looks like a giant gash in the crust. On the ground, it consists of a deep trough bordered on both sides by high cliffs formed by faulting. Along the length of the rift, several major volcanoes smoke and fume; these include the snow-crested Mt. Kilimanjaro, towering over 6 km above the savannah. At its north end, the rift joins the Red Sea Ridge and the Gulf of Aden Ridge at a triple junction. 

Collision 

India was once a small, separate continent that lay far to the south of Asia. But subduction consumed the ocean between India and Asia, and India moved northward, finally slamming into the southern margin of Asia about 40 to 50 million years ago. Continental crust, unlike oceanic crust, is too buoyant to subduct. So when India collided with Asia, the attached oceanic plate broke off and sank down into the deep mantle while India pushed hard into and partly under Asia, squeezing the rocks and sediment that once lay between the two continents into the 8-km-high welt that we now know as the Himalayan Mountains. During this process, not only did the surface of the Earth rise, but the crust became thicker. The crust beneath a collisional mountain range can be up to 60 to 70 km thick, about twice the thickness of normal continental crust. The boundary between what was once two separate continents is called a suture; slivers of ocean crust may be trapped along a suture.

Continental collision (not to scale).

Geoscientists refer to the process during which two buoyant pieces of lithosphere converge and squeeze together as collision (figure above a,  b). Some collisions involve two continents, whereas some involve continents and an island arc. When a collision is complete, the convergent plate boundary that once existed between the two colliding pieces ceases to exist. Collisions yield some of the most spectacular mountains on the planet, such as the Himalayas and the Alps. They also yielded major mountain ranges in the past, which subsequently eroded away so that today we see only their relicts.  For example, the Appalachian Mountains in the eastern United States formed as a consequence of three collisions. After the last one, a collision between Africa and North America around 300 Ma, North America  became part of the Pangaea supercontinent.

Credits: Stephen Marshak (Essentials of Geology)

Special Locations in the Plate Mosaic

Special Locations in the Plate Mosaic

kacimo samy   4:41 ص

Special Locations in the Plate Mosaic 

Triple Junctions 

Examples of triple junction. The triple junction are marked by dots.

Geologists refer to a place where three plate boundaries intersect as a triple junction, and name them after the types of boundaries that intersect. For example, the triple junction formed where the Southwest Indian Ocean Ridge intersects two arms of the Mid–Indian Ocean Ridge (this is the triple junction of the African, Antarctic, and Australian Plates) is a ridge-ridge-ridge triple junction (figure above a). The triple junction north of San  Francisco is a trench-transform-transform triple junction (figure above b).

Hot Spots 

 The dots represent the locations of selected hot-spot volcanoes. The red lines represent hot-spot tracks. The most recent volcano (dot) is at one end of this track. Some of these volcanoes are extinct, indicating that the mantle plume no longer exists. Some hot spots are fairly recent and do not have tracks. Dashed tracks were broken by sea-floor spreading.

Most subaerial (above sea level) volcanoes are situated in the volcanic arcs that border trenches. Volcanoes also lie along mid-ocean ridges, but ocean water hides most of them. The volcanoes of volcanic arcs and mid-ocean ridges are plate boundary volcanoes, in that they formed as a consequence of movement along the boundary. Not all volcanoes on Earth are plate-boundary volcanoes, however. Worldwide, geoscientists have identified about 100 volcanoes that exist as isolated points and are not a consequence of movement at a plate boundary. These are called hotspot volcanoes, or simply hot spots (figure above). Most hot spots are located in the interiors of plates, away from the boundaries, but a few lie along mid-ocean ridges. What causes hot-spot volcanoes? In the early 1960s, J. Tuzo Wilson noted that active hot-spot volcanoes (examples that are erupting or may erupt in the future) occur at the end of a chain of dead volcanic islands and seamounts (formerly active volcanoes that will never erupt again). This configuration is different from that of volcanic arcs along convergent plate boundaries at volcanic arcs, all of the volcanoes are active. With this image in mind, Wilson suggested that the position of the heat source causing a hotspot volcano is fixed, relative to the moving plate. In Wilson’s model, the active volcano represents the present-day location of the heat source, whereas the chain of dead volcanic islands represents locations on the plate that were once over the heat source but progressively moved off.

The deep mantle plume hypothesis for the formation of hot-spot tracks.

A few years later, researchers suggested that the heat source for hot spots is a mantle plume, a column of very hot rock rising up through the mantle to the base of the lithosphere (figure above a–d). In this model, plumes originate deep in the mantle. Rock in the plume, though solid, is soft enough to flow, and rises buoyantly because it is less dense  than surrounding cooler rock. When the hot rock of the plume reaches the base of the lithosphere, it partially melts and produces magma that seeps up through the lithosphere to the Earth’s surface. The chain of extinct volcanoes, or hot-spot track, forms when the overlying plate moves over a fixed plume. This movement slowly carries the volcano off the top of the plume, so that it becomes extinct. A new, younger volcano grows over the plume. 

The Hawaiian chain provides an example of the volcanism associated with a hot-spot track. Volcanic  eruptions occur today only on the big island of Hawaii. Other islands to the northwest are remnants of dead volcanoes, the oldest of which is Kauai. To the northwest of Kauai, still older volcanic remnants are found. About 1,750 km northwest of Midway Island, the track bends in a more northerly direction, and the volcanic remnants no longer poke above sea level; we refer to this northerly trending segment as the Emperor seamount chain. Geologists suggest that the bend is due to a change in the direction of Pacific Plate motion at about 40 Ma. 

Some hot spots lie within continents. For example, several have been active in the interior of Africa, and one now underlies Yellowstone National Park. The famous geysers (natural steam and hot-water fountains) of Yellowstone exist because hot magma, formed above the Yellowstone hot spot, lies not far below the surface of the park. While most hot spots, such as Hawaii and Yellowstone, occur in the interior of plates, away from plate boundaries, a few are positioned at points on mid-ocean ridges. The additional magma production associated with such hot spots causes a portion of the ridge to grow into a mound that can rise significantly above normal ridgeaxis depths and protrude above the sea surface. Iceland, for example, is the product of hot-spot volcanism on the axis of the Mid-Atlantic Ridge.
Credits: Stephen Marshak (Essentials of Geology)

Transform Plate Boundaries

Transform Plate Boundaries

kacimo samy   3:02 ص

Transform Plate Boundaries

The concept of transform faulting.

When researchers began to explore the bathymetry of midocean ridges in detail, they discovered that mid-ocean ridges are not long, uninterrupted lines, but rather consist of short segments that appear to be offset laterally from each other (figure above a) by narrow belts of broken and irregular sea floor. These belts, or fracture zones, lie roughly at right angles to the ridge segments, intersect the ends of the segments, and extend beyond the ends of the segments. Originally, researchers incorrectly assumed that the entire length of each fracture zone was a fault, and that slip on a fracture zone had displaced segments of the mid-ocean ridge sideways, relative to each other. In other words, they imagined that a mid-ocean ridge initiated as a continuous, fence-like line that only later was broken up by faulting. But when information about the distribution of earthquakes along mid-ocean ridges became available, it was clear that this model could not be correct. Earthquakes, and therefore active fault slip, occur only on the segment of a fracture zone that lies between two ridge segments. The portions of fracture zones that extend beyond the edges of ridge segments, out into the abyssal plain, are not seismically active.

The distribution of movement along fracture zones remained a mystery until a Canadian researcher, J. Tuzo Wilson, began to think about fracture zones in the context of the sea-floor-spreading concept. Wilson proposed that fracture zones formed at the same time as the ridge axis itself, and thus the ridge consisted of separate segments to start with. These segments were linked (not offset) by fracture zones. With this idea in mind, he drew a sketch map showing two ridge-axis segments linked by a fracture zone, and he drew arrows to indicate the direction that ocean floor was moving, relative to the ridge axis, as a result of sea-floor-spreading (figure above b). Look at the arrows in figure above b. Clearly, the movement direction on the active portion of the fracture zone must be opposite to the movement direction that researchers originally thought occurred on the structure. Further, in Wilson’s model, slip occurs only along the segment of the fracture zone between the two ridge segments (figure above c). Plates on opposite sides of the inactive part of a fracture zone move together, as one plate. 

Wilson introduced the term transform boundary, or transform fault, for the actively slipping segment of a fracture zone between two ridge segments, and he pointed out that these are a third type of plate boundary. At a transform boundary, one plate slides sideways past another, but no new plate forms and no old plate is consumed. Transform boundaries are, therefore, defined by a vertical fault on which the slip direction parallels the Earth’s surface. The slip breaks up the crust and forms a set of steep fractures. 

So far we've discussed only transforms along mid-ocean ridges. Not all transforms link ridge segments. Some, such as the Alpine Fault of New Zealand, link trenches, while others link a trench to a ridge segment. Further, not all transform faults occur in oceanic lithosphere; a few cut across continental lithosphere. The San Andreas Fault, for example, which cuts across California, defines part of the plate boundary between the North American Plate and the Pacific  Plate the portion of California that lies to the west of the fault (including Los Angeles) is part of the Pacific Plate, while the portion that lies to the east of the fault is part of the North American Plate (figure above d, e).

Credits: Stephen Marshak (Essentials of Geology)

Convergent Plate Boundaries and Subduction

Convergent Plate Boundaries and Subduction

kacimo samy   2:56 ص

Convergent Plate Boundaries and Subduction 

At convergent plate boundaries, two plates, at least one of which is oceanic, move toward one another. But rather than butting each other like angry rams, one oceanic plate bends and sinks down into the asthenosphere beneath the other plate. Geologists refer to the sinking process as subduction, so convergent boundaries are also known as subduction zones. Because subduction at a convergent boundary consumes old ocean lithosphere and thus ‘‘consumes’’ oceanic basins, geologists also refer to convergent boundaries as consuming boundaries, and because they are delineated by deep-ocean trenches, they are sometimes simply called trenches. The amount of oceanic plate consumption worldwide, averaged over time, equals the amount of sea-floor spreading worldwide, so the surface area of the Earth remains constant through time. 

During the process of subduction, oceanic lithosphere sinks back into the deeper mantle.

Subduction occurs for a simple reason: oceanic lithosphere, once it has aged at least 10 million years, is denser than the underlying asthenosphere and thus can sink through the asthenosphere if given an opportunity. Where it lies flat on the surface of the asthenosphere, oceanic lithosphere can’t sink. However, once the end of the convergent plate bends down and slips into the mantle, it continues downward like an anchor falling to the bottom of a lake (figure above a). As the lithosphere sinks, asthenosphere flows out of its way, just as water flows out of the way of a sinking anchor. But unlike water, the asthenosphere can flow only very slowly, so oceanic lithosphere can sink only very slowly, at a rate of less than about 15 cm per year. To visualize the difference,  imagine how much faster a coin can sink through water than it can through honey. 

Note that the “downgoing plate,” the plate that has been subducted, must be composed of oceanic lithosphere. The overriding plate, which does not sink, can consist of either oceanic or continental lithosphere. Continental crust cannot be subducted because it is too buoyant; the low-density rocks of continental crust act like a life preserver keeping the continent afloat. If continental crust moves into a convergent margin, subduction eventually stops. Because of subduction, all ocean floor on the planet is less than about 200 million years old. Because continental crust cannot subduct, some continental crust has persisted at the surface of the Earth for over 3.8 billion years.

Earthquakes and the Fate  of Subducted Plates 

At convergent plate boundaries, the downgoing plate grinds along the base of the overriding plate, a process that generates large earthquakes. These earthquakes occur fairly close to the Earth’s surface, so some of them cause massive destruction in coastal cities. But earthquakes also happen in downgoing plates at greater depths. In fact, geologists have detected earthquakes within downgoing plates to a depth of 660 km. The band of earthquakes in a downgoing plate is called a WadatiBenioff zone, after its two discoverers (figure above b).

At depths greater than 660 km, conditions leading to earthquakes in subducted lithosphere evidently do not occur. Recent observations, however, indicate that some downgoing plates do continue to sink below a depth of 660 km they just do so without generating earthquakes. In fact, the lower mantle may be a graveyard for old subducted plates.

Geologic Features of a Convergent Boundary 

To become familiar with the various geologic features that occur along a convergent plate boundary, let’s look at an example, the boundary between the western coast of the South American Plate and the eastern edge of the Nazca Plate (a portion of the Pacific Ocean floor). A deep-ocean trench, the Peru-Chile Trench, delineates this boundary (figure above b). Such trenches form where the plate bends as it starts to sink into the asthenosphere. 

In the Peru-Chile Trench, as the downgoing plate slides under the overriding plate, sediment (clay and plankton) that had settled on the surface of the downgoing plate, as well as sand that fell into the trench from the shores of South America, gets scraped up and incorporated in a wedge-shaped mass known as an accretionary prism (figure above c). An accretionary prism forms in basically the same way as a pile of snow or sand in front of a plow, and like snow, the sediment tends to be squashed and contorted. 

A chain of volcanoes known as a volcanic arc d evelops behind the accretionary prism. The magma that feeds these volcanoes forms just above the surface of the downgoing plate where the plate reaches a depth of about 150 km below the Earth’s surface. If the volcanic arc forms where an oceanic plate subducts beneath continental lithosphere, the resulting chain of volcanoes grows on the continent and forms a continental volcanic arc. (In some cases, the plates squeeze  together across a continental arc, causing a belt of faults to form behind the arc.) If, however, the volcanic arc grows where one oceanic plate subducts beneath another oceanic plate, the resulting volcanoes form a chain of islands known as a volcanic island arc (figure above d). A back-arc basin exists either where subduction happens to begin offshore, trapping ocean lithosphere behind the arc, or where stretching of the lithosphere behind the arc leads to the formation of a small spreading ridge behind the arc (figure above e).

Credits: Stephen Marshak (Essentials of Geology)

Divergent Plate Boundaries and Sea-Floor Spreading

Divergent Plate Boundaries and Sea-Floor Spreading

kacimo samy   2:49 ص

Divergent Plate Boundaries and Sea-Floor Spreading 

The process of sea-floor spreading.

At a divergent boundary, or spreading boundary, two oceanic plates move apart by the process of sea-floor spreading. Note that an open space does not develop between diverging plates. Rather, as the plates move apart, new oceanic lithosphere forms continually along the divergent boundary (figure above a). This process takes place at a submarine mountain range called a mid-ocean ridge that rises 2 km above the adjacent abyssal plains of the ocean. Thus, geologists commonly refer to a divergent boundary as a mid-ocean ridge, or simply a ridge. Water depth above ridges averages about 2.5  km. 

To characterize a divergent boundary more completely, let’s look at one mid-ocean ridge in more detail (figure above b). The Mid-Atlantic Ridge extends from the waters between northern  Greenland and northern Scandinavia southward across the equator to the latitude of the southern tip of South America. Geologists have found that the formation of new sea floor takes place only along the axis (centerline) of the ridge, which is marked by an elongate valley. The sea floor slopes away, reaching the depth of the abyssal plain (4 to 5 km) at a distance of about 500 to 800 km from the ridge axis. Roughly speaking, the Mid-Atlantic Ridge is symmetrical its eastern half looks like a mirror image of its western half. The ridge consists, along its length, of short segments (tens to hundreds of km long) that step over at breaks that, as we noted earlier, are called fracture zones.

How Does Oceanic Crust Form  at a Mid-Ocean Ridge? 

As sea-floor spreading takes place, hot asthenosphere rises beneath the ridge and begins to melt, and molten rock, or magma, forms (figure above c). Magma has a lower density than solid rock, so it behaves buoyantly and rises, as oil rises above vinegar in salad dressing. Molten rock eventually accumulates in the crust below the ridge axis, filling a region called a magma chamber. As the magma cools, it turns into a mush of crystals. Some of the magma solidifies completely along the side of the chamber to make the coarse-grained, mafic igneous rock called gabbro. The rest rises still higher to fill vertical cracks, where it solidifies and forms wall-like sheets, or dikes, of basalt. Some magma rises all the way to the surface of the sea floor at the ridge axis and spills out of small submarine volcanoes. The resulting lava cools to form a layer of basalt blobs called pillows. Observers in research submarines have detected chimneys spewing hot, mineralized water rising from cracks in the sea floor along the ridge axis. These chimneys are called black smokers because the water they emit looks like a cloud of dark smoke; the colour comes from a suspension of tiny mineral grains that precipitate in the water the instant that the water cools (figure below). 

 A column of superhot water gushing from a vent known as a black smoker along the mid-ocean ridge. A local ecosystem of bacteria, shrimp, and worms lives around the vent.

As soon as it forms, new oceanic crust moves away from the ridge axis, and when this happens, more magma rises from below, so still more crust forms. In other words, like a vast, continuously moving conveyor belt, magma from the mantle rises to the Earth’s surface at the ridge, solidifies to form oceanic crust, and then moves laterally away from the ridge. Because all sea floor forms at mid-ocean ridges, the youngest sea floor occurs on either side of the ridge axis, and sea floor becomes progressively older away from the ridge. In the Atlantic Ocean, the oldest sea floor, therefore, lies adjacent to the passive continental margins on either side of the ocean (figure below). The oldest ocean floor on our planet underlies the western Pacific Ocean; this crust formed about 200 million years ago.

This map of the world shows the age of the sea floor. Note how the sea floor grows older with increasing distance from the ridge axis. (Ma = million years ago.)

The tension (stretching force) applied to newly formed solid crust as spreading takes place breaks the crust, resulting in the formation of faults. Slip on the faults causes divergent boundary earthquakes and produces numerous cliffs, or scarps, that lie parallel to the ridge axis.

How Does the Lithospheric Mantle  Form at a Mid-Ocean Ridge? 

So far, we've seen how oceanic crust forms at mid-ocean ridges. How does the mantle part of the oceanic lithosphere form? This part consists of the cooler uppermost layer of the mantle, in which temperatures are less than about 1,280°C. At the ridge axis, such temperatures occur almost at the base of the crust, because of the presence of rising hot asthenosphere and hot magma, so lithospheric mantle beneath the ridge axis effectively doesn’t exist. But as the newly formed oceanic crust moves away from the ridge axis, the crust and the uppermost mantle directly beneath it gradually cool by losing heat to the ocean above. As soon as mantle rock cools below 1,280°C, it becomes, by definition, part of the lithosphere.

Changes accompanying the aging of lithosphere.

As oceanic lithosphere continues to move away from the ridge axis, it continues to cool, so the lithospheric mantle, and therefore the oceanic lithosphere as a whole, grows progressively thicker (figure above a, b). Note that this process doesn’t change the thickness of the oceanic crust, for the crust formed entirely at the ridge axis. The rate at which cooling and lithospheric thickening occur decreases progressively with increasing  distance from the ridge axis. In fact, by the time the lithosphere is about 80 million years old, it has just about reached its maximum thickness. As lithosphere thickens and gets cooler and denser, it sinks down into the asthenosphere, like a ship taking on ballast. Thus, the ocean is deeper over older ocean floor than over younger ocean floor.

Credits: Stephen Marshak (Essentials of Geology)