Introduction
If you were to cut the earth in half directly along the line of the equator, you would see that our planet is not a homogeneous ball of matter. It has internal structure or layering associated with changes in composition, density, and temperature with depth; the concentration of metals, the density of matter, and the temperature all increase toward the earth’s core. Figure 1 displays a simplified model of the earth’s interior or geosphere. This model was developed by geologists over many years from the synthesis and interpretation of a great deal of physical and chemical data. The outermost layers of the geosphere consist of crust and lithosphere, essentially bedrock that has been formed and reformed, assembled and disassembled, as the “stuff” of continents and ocean basins through the ages. Knowledge of how the earth’s crust forms and is deformed, what rock types occur where and why, and of what those rocks consist is a vital aspect of the study of geology.
Figure 1. A simplified model displaying the concentric layers of material with differing thickness, composition, and density that comprise the earth’s interior.
Plate Tectonics
Every scientific discipline has certain fundamental theories that provide a framework on which to hang a multitude of observations, computations, and interpretations. In the discipline of geology, the plate tectonic theory can be used to explain nearly all of earth’s major geologic features; especially those associated with mountain building, crustal composition and deformation, major sea level fluctuations, erosive events, and basin filling events. According to plate tectonic theory, the earth’s rigid, brittle outer rind (comprised of the crust and lithosphere) is broken into a number of large (and small) pieces, or plates, that float on and are dragged about by convective flow within a deeper, hotter, mushy-plastic layer (the upper mantle or asthenosphere) (Figure 2). Tectonic plates may be comprised of either, or both, continental crust and lithosphere, or oceanic crust and lithosphere. Continental material, which is dominantly felsic in composition, is less dense than oceanic material, which is dominantly mafic; thus, continental plates tend to float higher on the mantle to form continents, while oceanic plates tend to sink lower and form ocean basins. The tectonic forces of mantle convection result in three additional drivers for the motions of these plates (Figure 2). Because the asthenosphere is a solid, albeit extremely hot and plastic-like, there is considerable friction between the tectonic plates and convecting mantle; thus, frictional drag by the flowing asthenosphere on the underside of a tectonic plate causes that plate to be rafted along much like a chunk of ice floating on a river. Mantle convection also pushes on a tectonic plate from behind, at the plate’s newly formed, buoyant margin, while gravity pulls on the leading edge of the plate or slab of crust and lithosphere where it is older, colder, and denser.
Figure 2. The earth’s tectonic plates are rafted about on the upper mantle’s asthenosphere, a hot, mushy, plastic material undergoing convective flow.
The earth is essentially spherical, and because it is not expanding, tectonic plates may interact in three geometric arrangements: plates may split and move away from each other; plates may collide; or plates may shear past one another like the blades of a pair of scissors. Each interaction defines a style of plate motion and generates certain geologic processes and features that characterize different plate tectonic boundaries: divergent, convergent, and transform. Divergent boundaries result from tensional forces that rift and pull plates apart (Figure 3). As the plate fragments first begin to move away from each other, the plate material lifts, fractures, and thins, and then sags into the developing gap to form two halves of a mid-oceanic ridge bisected by a rift zone. Over time, the plate fragments separate completely, forming a basin (that naturally floods with seawater). Molten mantle material rises and solidifies along the edges of the diverging plates, forming new oceanic crust and lithosphere within the growing gap, the oceanic crust becoming progressively buried by layers of sediment over time (Figure 3). The mid-ocean ridge separating diverging plates is often sheared and broken by transform faults as the two plates pivot away from each other (the plates cannot move at the same rate everywhere along the boundary because the earth is spherical – plates move faster at the equator than at higher latitudes).
Figure 3. A mature divergent plate boundary exhibiting an ocean basin and well-developed spreading center or rift zone separating oceanic crust and lithosphere of progressively increasing age.
Convergent boundaries result from compressional forces that cause plates to collide. Plate collisions are complicated by the composition and density of the colliding material, resulting in three possible convergent boundaries. When plates of differing density collide, the less dense plate is forced up and over, while the denser plate dives down and under in a process called subduction. During oceanic-to-oceanic plate collisions, the older, colder, less buoyant oceanic plate subducts; while during oceanic-to-continental plate collisions, the oceanic plate, being denser, always subducts (Figure 4). Subduction causes the downgoing crustal slab and overlying sediments to partially melt, generating magma, which rises upward through the overlying plate, while the remainder of the descending plate is reincorporated into the upper mantle. Rising magma may cool and solidify within the crust as plutons or erupt at the surface as volcanoes, the combined features form a magmatic (volcanic) arc (Figure 4), a chain of mountains with an arcuate shape controlled by the geometry of the collision boundary on a spherical earth. As one plate subducts, slivers of crustal rock and ocean floor sediments are scraped up, bulldozer-like, by the leading edge of the overriding plate to form an accretionary wedge, while sediments striped by erosion from the volcanic arc accumulate in adjacent fore-arc and back-arc basins (Figure 4). Alternatively, when continental plates collide, neither subduct because they are of similar composition and density, instead they crumple together and compete for “stacking” rights. Collisions form long, linear mountain belts of deformed crust and/or magmatic arcs of solidified molten rock.
Figure 4. A mature convergent plate boundary exhibiting the subduction of a denser oceanic plate beneath an adjacent continental plate; active continental plate margins such as this are often comprised of parallel accretionary wedge and volcanic arc, with adjacent fore-arc and back-arc basins.
Transform boundaries result from shearing forces as plates pass by each other side-by-side (Figure 5). Because transform boundaries are not perfectly straight (they have kinks), shearing can generate localized pull-apart basins or collisional uplift (transtension and transpression, respectively), and relatively broad zones of deformation consisting of multiple en echelon fault segments, but overall motion of the plates is lateral (aside one another). Transform boundaries can occur as stand-alone features, but often these plate margins occur in combination with divergent boundaries (Figure 3); essentially “taking up the slack” as two adjacent, diverging plates pull apart along a curved boundary.
Figure 5. Shearing and the formation of a transform plate boundary.
Mantle plumes are another source of tectonic activity related to plate motions. Although their origin is poorly understood, it is generally assumed that mantle plumes occur at locations where heat sources deep within the earth’s interior cause warming and outward expansion of the overlying mantle (Figure 6). (Heating of a substance causes it to expand in the direction of least resistance, which in this case is upward toward lesser pressure and lower density material). Mantle plumes may be independent of mantle convection, or they may be the initial driver forming new mantle convection cells. In either case, these plumes result in decompression-melting of ascending asthenosphere, melting to form magma that pools in the plume head at the base of the lithosphere. As the plume head expands, mushroom-like, motion of the overlying plate drags the plume head to the side. Pulses of magma periodically burn their way upward through the overlying tectonic plate as it moves, forming a “hot spot” where the magma erupts through vents onto the surface constructing a linear chain of volcanic landforms on the plate’s surface increasing in age away from the hot spot, while the mantle plume remains more or less stationary, fixed to its deeper mantle source.
Figure 6. A mantle plume model, showing the postulated source of magma and associated hot spot volcanism trailing away from the current hot spot in the direction of motion of the overlying plate.
Composition of the Earth’s Crust
Minerals
The materials that form the earth’s crust rarely occur as pure substances composed of the atoms of only one type of element (native metal’s such as gold being an exception); instead, they commonly occur as chemical compounds made up of several elements. These naturally occurring substances are called minerals. Minerals form the fundamental building blocks of the three basic types of rock (igneous, sedimentary, and metamorphic) that make up the earth’s crust.
Any substance with a specific, orderly internal arrangement of atoms is a crystalline solid, but for it to be considered a mineral, it must be naturally occurring, inorganic, and have a definite chemical composition that can be expressed by a formula. Quartz is composed of the elements silicon (Si) and oxygen (O) arranged into complex three-dimensional structures and having a chemical formula SiO2 (a crystalline arrangement resulting in a ratio of one silicon atom for every two oxygen). Quartz is a major constituent of many igneous rocks. In contrast, volcanic glass, or obsidian, common in central Oregon, is also composed of silicon and oxygen, but is amorphous (a substance that is solid but not crystalline, and therefore not a mineral, because its atoms lack an orderly internal arrangement).
Minerals are formed by the process of crystallization (the growth of a solid from a material whose constituent atoms combine in proper chemical proportions and in orderly three-dimensional arrangements). They are often grouped into categories related to such characteristics as the dominant element or element group they contain, how/where they form, and what types of rock they result in or are associated with. Crystallization can occur in several ways: 1) cooling a liquid below its melting point causes crystals to form, a process that produces various minerals during the solidification of magma to form igneous rock; 2) evaporation of a salt solution (such as sea water) concentrates the ions dissolved into it, and the remaining solution becomes saturated (it cannot hold any more ions) so that any further evaporation causes the ions to chemically bond and form crystals that precipitate from the solution, a process that is responsible for the formation of certain chemical sedimentary rocks; 3) crystals may form by rearrangement of ions and/or atoms in solid materials (recrystallization) at high temperatures and/or pressures, a process known as metamorphism that is enhanced as conditions become more extreme, or when in the presence of solutions, and which results in the formation of metamorphic rocks; and 4) crystallization from a saturated hydrothermal (heated-water) solution, a process that often occurs in the proximity of magmatic intrusions and/or is associated with metamorphism and results in the formation of hydrothermal deposits that contain metallic ore minerals either by concentrated precipitation of crystals directly onto the surfaces of fractures and joints and/or disseminated precipitation of crystals within a rock body’s pore spaces.
Silicate minerals dominate the earth’s crustal composition, generated by either solidification of magma or through metamorphic recrystallization of preexisting silicates. The tectonic setting and concomitant volcanism of central Oregon assures that silicates will comprise the dominant constituent minerals in the rocks and sedimentary deposits of this region. Silicate minerals crystallize from a cooling magma in an orderly fashion that is primarily related to its elemental composition and temperature at the time of crystallization (a process characterized by Bowen’s Reaction Series, after N.L. Bowen, its discoverer). Figure 7 displays the resultant minerals and igneous rocks formed by this process; these minerals are generally referred to as the “rock-forming minerals” because of their significance to crustal composition.
Figure 7. A diagram illustrating Bowen’s Reaction Series and the resultant igneous “rock-forming minerals” formed by this process. The colored arrows under “mineral types” loosely correspond to mineral coloration.
Silicate minerals are a major class of minerals which contain silicon and oxygen bonded in the form of a silica tetrahedron (a pyramidal shaped anion group consisting of one silicon and four oxygen atoms – the complex anion SiO4-4) as their basic chemical unit. These silica tetrahedra can bond together in different patterns of arrangement which result in several silicate mineral families: 1) minerals composed of single, independent SiO4 units which are interconnected through their bonds with other elements (examples include olivine and garnet); 2) minerals composed of silica tetrahedra that share two of their oxygen atoms with adjacent SiO4 units to form single, chain-like structures (examples include the pyroxene mineral group, such as the mineral augite); 3) minerals composed of silica tetrahedra in which some SiO4 units share two oxygen atoms and others share three, forming double-chain structures (examples include the amphibole mineral group, such as the mineral hornblende); 4) minerals composed of silica tetrahedra in which the SiO4 units share three oxygen atoms to form sheet-like structures (examples include the mica mineral group, such as biotite and muscovite); and 5) minerals composed of silica tetrahedra in which the SiO4 units share all four oxygen atoms in bonds with adjacent units, forming a framework structure (examples include the feldspar mineral group (plagioclase and orthoclase) and quartz. Exploring the relationships between Bowen’s Reaction Series and the formation of silicate minerals, you may notice that the more complex the chemical bonding arrangement, the more enriched the magma is with respect to silica and the lower the temperature of the magma at which the mineral forms (Figure 7).
Rocks (The Rock Cycle)
Rocks are complex aggregates of minerals that make up the solid portion of the earth’s crust. Three classes of rocks, igneous, sedimentary, and metamorphic, are recognized on the basis of how and where they formed (origin), their appearance (texture), and their composition (mineralogy). In general, the formation of rock is cyclical, and can be best expressed by the Rock Cycle (Figure 8). As the diagram suggests, there is no beginning or ending to rock formation. In the most basic cycle: 1) an igneous rock forms from the solidification of magma; 2) physical and chemical weathering and mechanical abrasion during transport to a new location transforms igneous rock into sediment; 3) sediment is lithified by compaction and cementation to become sedimentary rock; 4) heat and pressure combine to physically and chemically alter the sedimentary rock into metamorphic rock; and 5) with enough heat, the metamorphic rock is melted to form magma, magma cools and crystallizes to form igneous rock, and the cycle is renewed. However, as Figure 8 illustrates, this cycle can be shortened if a sedimentary rock or a preexisting igneous rock were melted to form magma; if a metamorphic rock or a preexisting sedimentary rock were weathered and eroded to form sediment; or if an igneous or a preexisting metamorphic rock were subjected to metamorphism to form a metamorphic rock. These processes are intimately linked with plate tectonic activity. Magma and the pressure-temperature conditions for metamorphism are of course generated by tectonic activity, but the uplift necessary for weathering and erosion (and the production of sediment), as well as subsidence necessary for basin filling (deposition of sediment) is also often a result of tectonics (or sea level change which can result from tectonic activity).
Figure 8. The fundamental processes and rock types comprising Earth’s Rock Cycle. Note that each rock type can result from three process pathways; for example, igneous rocks can form from the initial melting of metamorphic, sedimentary, or even other igneous rocks.
Igneous rocks are those that form from the cooling and crystallization of molten material called magma (Figure 8). As indicated earlier, solidification of magma results in the crystallization of silicate minerals. Silicate minerals crystallize from a cooling magma in an orderly fashion that is primarily related to the molten material’s elemental composition and temperature at the time of crystallization (a process characterized by Bowen’s Reaction Series), and the minerals that chiefly constitute igneous rock can be divided into three groups, mafic, intermediate, and felsic based on the dark, salt-and-peppery, or light color they exhibit (which is more directly related to their elemental content) (Figure 7). Thus, as we shall see, color becomes an important tool for igneous rock identification.
The source for magma is not the earth’s liquid outer core; instead, it is generated at relatively shallow depths within the lithosphere and/or asthenosphere, usually at convergent and divergent plate boundaries and mantle plume hot spots. The newly generated magma is lighter than the surrounding rocky material and it rises toward the surface. The magma generating source(s) generally determines the magma’s composition and the silicate minerals that will crystalize from it, mantle sources generally contain higher concentrations of ferromagnesium ions and tend to crystalize mafic minerals, while lithospheric sources generally contain more silica generating ions and crystallize abundant felsic minerals.
Magma can cool and crystallize at depth within the crust or when it is extruded as lava onto the earth’s surface, but the location where crystallization occurs determines the cooling rate and thus the size of the crystals formed (referred to as the igneous rock’s texture). When identifying an igneous rock, its texture is determined by the groundmass (the most common background crystal size) of the rock. If the rising magma doesn’t reach the surface, but cools slowly and crystallizes within the earth’s’ crust, it forms intrusive (plutonic) igneous rock characterized by a uniformly coarse-grained groundmass of crystals that are large enough to be seen with the unaided eye, such as the granite shown in Figure 9a. If the rising magma reaches near the surface or is extruded at the surface it will cool rapidly and form extrusive (volcanic) igneous rock typically characterized by a fine-grained groundmass of crystals that are too small to be seen with the unaided eye, such as the rhyolite shown in Figure 9b. When volcanic igneous rocks form by extremely rapid cooling, crystals have no time to form and take on a glassy appearance like the obsidian shown in Figure 9c. The rising magma may begin cooling and undergo partial crystallization at depth before being extruded as lava to cool and crystallize rapidly at the surface. The resulting volcanic igneous rock will contain two distinct grain sizes, and is described as having a porphyritic texture; the term porphyry is tacked on as a suffix to the rock name, such as the sample of andesite porphyry shown in Figure 9d.
Figure 9. These photographs display coarsely crystalline (A), finely crystalline (B), glassy (C), and porphyritic (D) igneous rock textures.
Volcanism often releases copious amounts of ash and rock fragments (or lapilli) known as pyroclastic material which may travel through the air for varying lengths of time and distance (smaller particles will stay aloft longer and travel farther from the volcanic source). The ash and lapilli may be loosely consolidated by burial and compaction to form fragmental tuff (Figure 10a) or become rapidly indurated during the release of heat from extremely hot volcanic ash to form a banded, welded tuff (Figure 10b). The outpouring of lava and pyroclastic material at the surface is often associated with much escaping gas and the cooling and solidification of gaseous lava extruded at the surface often traps bubbles called vesicles within the crystallizing lava to form a fine-grained vesicular volcanic rock called scoria (Figure 10c) and a vesicular glassy volcanic rock called pumice (Figure 10d).
Figure 10. These photographs provide examples of fine-grained pyroclastic igneous rocks called fragmental tuff (A), and welded tuff (B), as well as the fine-grained, vesicular igneous rock called scoria (C) and glassy, vesicular igneous rock called pumice (D).
Geologists thus classify igneous rocks on the basis of mineralogy (color) and textural characteristics, the mineralogy controlled by the magma composition and the texture mainly controlled by the rate of cooling of the magma. Figure 11 shows ten basic types of igneous rocks determined by this classification scheme: 1) magma of ultra-mafic composition yields coarse-grained peridotite or much more rarely fine-grained komatiite; 2) mafic composition yields coarse-grained gabbro and fine-grained basalt; 3) magma of intermediate composition yields coarse-grained diorite and fine-grained andesite; and 4) magma of felsic composition yields coarse-grained granodiorite to granite and fine-grained dacite to rhyolite. The central portion of Figure 11 identifies the general composition and color of the igneous rock (igneous rocks containing abundant felsic minerals are predominantly light-colored, whereas igneous rocks containing mostly mafic minerals are dark-colored). When identifying coarse-textured igneous rocks, use the dominant color of the visible crystals; but when identifying fine-textured igneous rocks, the color of the groundmass is used. For example, granites are generally composed of light-colored minerals, while rhyolites are composed of light-colored groundmass (Figure 9a and Figure 9b). The bottom portion of Figure 11 shows the key compositional characteristics of each main category of igneous rocks; read from left to right, one can determine the approximate percentage by volume of the major mineral constituents associated with a given rock type. The colors in this part of the table correspond roughly with mineral color and can be used to gage the color of the igneous rock. This portion of the table is particularly useful for identifying phaneritic igneous rocks because mineral crystals can be observed and identified with practice. For example, granite, a felsic igneous rock, typically contains more than 20% quartz, about 40% orthoclase feldspar, about 30% Na-rich plagioclase feldspar, and roughly 10% mica (biotite and muscovite) and amphibole (hornblende).
Figure 11. This diagram shows ten basic types of igneous rocks classified on the basis of texture and mineral content (or color).
Specialized names are applied to granitic igneous rocks with an abundance of certain minerals. The name syenite is used to describe a granite that contains only about 5% quartz and mostly orthoclase feldspar; whereas the name monzonite is used to describe a granite that contains roughly 5% quartz, but Na-rich plagioclase feldspar tends to dominate. Trachyte and latite are the fine-grained felsic equivalents of syenite and monzonite, respectively.
Igneous rocks are formed either by cooling of viscous magma at depth, or by cooling of less viscous magma that erupts to form lava and/or pyroclastics at the surface. These basic differences in the locus of magma cooling and solidification have led to the use of the terms intrusive (plutonic) and extrusive (volcanic) to describe the origin of igneous rocks. Both types of igneous rock occur as larger “bodies” with certain unique characteristics (Figure 12):
1) Intrusive igneous bodies (plutons)
- Batholith – large masses of igneous rock with no known floor that may represent singular or multiple, coalescing chambers of solidified magma.
- Stock – same as batholith, but smaller than 40 square miles, may represent an isolated portion of a larger batholith.
- Dike – tabular mass or sheet that cuts across pre-existing “country” rocks at an angle.
- Sill – tabular sheet injected between parallel rock layers.
- Laccolith – dome-shaped mass (a mushroom-shaped sill) intruded into layered sedimentary rocks.
- Lopolith – basin-shaped mass intruded into layered sedimentary rocks.
- Volcanic neck – solidified volcanic vent that remains standing after the surrounding cone has been eroded away.
2) Extrusive igneous bodies
- Volcanic cones – conical structures composed of lava and/or pyroclastic material extruded from a central vent.
- Cinder cone – small, steep-sided cone, principally composed of cinders.
- Shield volcano – large, gradually sloping cone, composed largely of mafic lava flows.
- Strato (composite) volcano – moderately large, graceful cone, composed of alternative layers of ash and lava.
- Lava flows – sheets of molten lava extruded from a vent which run down the flanks of the cone and spread out at its base.
- Pyroclastic deposits – volcanic material ejected from a vent which generally forms a veneer of coarse to fine debris as distance from the source increases; comprised of volcanic bombs, blocks, and rock fragments, lapilli, and ash.
Figure 12. Common intrusive and extrusive igneous rock bodies.
Sedimentary rocks are the product of the weathering and erosion of preexisting rocks to form sediment, the transportation, abrasion, sorting, and eventual deposition of that sediment, and lithification of that material by compaction and cementation. The weathering of minerals in igneous rocks is fundamental to understanding the nature of sediments and sedimentary rocks. When igneous rock-forming minerals undergo chemical decomposition to produce new substances, the general rule of thumb is “that which crystallizes first, weathers first” (that is, the least stable minerals formed at the highest pressure/temperature conditions). Therefore, minerals such as olivine and calcium-rich plagioclase feldspar and rocks enriched in these minerals are the least stable, while quartz and quartz-rich rocks are the most stable (Figure 7 and Figure 11). The ferromagnesian minerals (minerals rich in iron, magnesium, and calcium – olivine, augite, hornblende, and biotite) decompose rapidly to form soluble salts and small amounts of silica and clay. The soluble iron may oxidize to form the iron oxide minerals limonite and hematite – sedimentary iron ore. Muscovite mica is relatively stable, and only small amounts are dissolved. Muscovite fragments cleave and are broken into very small flakes; the smallest flakes sometimes undergo further change and become clay minerals. Feldspars are generally decomposed rapidly to form clay minerals such as kaolinite and soluble cations of calcium, sodium, and potassium that are carried out to sea in solution where they recombine with other anions and may precipitate as new minerals such calcite, dolomite, or gypsum. Quartz is one of the most stable minerals and is only slightly affected by weathering. Quartz mineral fragments are freed from the rock matrix, broken down and rounded, and become the grains of sand and silt of beaches, river valleys, and sand dunes, as well as ocean floor muds.
Sedimentary rocks are those that have formed as aggregates of clastic material either eroded from pre-existing rocks, aggregates of bio-chemically and/or chemically precipitated minerals that accumulate on the floors of oceans, lakes, and playas, or that form from the deposition of organic (mainly plant) detritus. Pyroclastic material ejected during volcanic eruptions to accumulate on the landscape can also be considered “clastic” and is often described as volcaniclastic sediment. These materials are later lithified (buried, compacted, and cemented together) to become sedimentary rock. A more detailed description of sedimentary rocks and their depositional environments can be found in the section of my website entitled THE GEOLOGY OF SEDIMENTARY ROCKS.
Clastic sedimentary rocks are classified by the size of the particles found within the rock and the dominant mineral composition (Figure 13). Rock containing at least 30% gravel sized particles is called a conglomerate if the grains are rounded (Figure 14a) or a breccia if the grains are angular (Figure 14b). Clastic sedimentary rock dominantly composed of sand-sized grains (more than 50%) forms a sandstone (Figure 14c), while clastic rocks containing greater than 50% silt and clay are called mudstone. If a mudstone contains greater than 2/3 silt, it is typically classified as a siltstone (Figure 14d) and if it contains greater than 2/3 clay, it is classified as a claystone (Figure 14e), although claystones that are fissile (show thin, parallel layering) are instead called shale. Clastic sediments can accumulate in nearly any terrestrial, transitional, or marine depositional environments. Relatively coarse-grained clastic rocks typically involve short transport distances and a limited amount of abrasion of the sediment, while fine-grained clastics require considerable transport and abrasion, recycling from one depositional setting to another, and chemical alteration to form clay.
Figure 13. Classification of Clastic Sedimentary Rocks.
Figure 14. Clastic sedimentary rocks are classified primarily on the basis of grain size, from gravel-sized particles forming conglomerate (A) and breccia (if angular) (B), to sand-sized particles forming sandstone (C), to silt-sized particles forming siltstone, and clay-sized particles forming claystone (shale if fissile).
Figure 15 shows a simple classification for biological, chemical, and biochemical sedimentary rocks. Terrestrial and transitional marine biological detritus (plant fragments) that accumulates in oxygen-poor environments, often settings where poorly circulated or stagnant water occurs, can form peat or coal. Peat is distinguished from coal by the loose, porous, poorly cemented nature of its constituent organic material. Coal forms through varying degrees of heat and pressure from burial and compaction, and can become so indurated that it is considered a metamorphic rock. Chemical sedimentary rocks can be formed by direct precipitation processes. Ions in saltwater are concentrated by evaporation. When enough water has been removed, the ions will bond to form mineral crystals which settle to the floor of a playa (a terrestrial desert basin often surrounded by mountainous terrain), or a shallow marine shelf or coastal sabkha, often with restricted inflow from the main ocean, such as coastal lagoons and tidal flats. Massive limestones (Figure 16a) composed of tiny calcite crystals and evaporite rocks such as rock gypsum (gypsum crystals) and rock salt (halite crystals) form in this way. Crystals can also be indirectly precipitated from saltwater as the hard parts (fossils) of formerly living marine organisms through biomineralization, and preserved as fossiliferous limestone (Figure 16b), formed as an aggregate of many fossil fragments which may not be recognizable due to erosion, or chalk, a fine-grained earthy material formed from the skeletal remains of marine microorganisms. Oolitic limestone (Figure 16c) is formed as an aggregate of many tiny spherical balls of lime mud called oolites, each oolite is formed by accumulation of fine concentric layers of lime mud as they roll around on the sea floor, while intraclastic limestone (Figure 16d) forms from the recemented fragments of a previously deposited limestone ripped up from the sea floor by storm wave activity. Chemical or biological silica in the form of very fine grained (cryptocrystalline) quartz can also precipitate as a concentrated siliceous ooze on the deep ocean floor which is incorporated within calcareous sediment to eventually form chert nodules or bedded chert. Silica can also form the hard parts of shelled organisms such as freshwater diatoms in playas and shallow lakes.
Figure 15. Classification of Biological, Biochemical, and Chemical Sedimentary Rocks.
Figure 16. The classification of limestones into its common forms: massive (A), fossiliferous (B), oolitic (C), and intraclastic (D) varieties.
Metamorphic rocks are formed from preexisting igneous, sedimentary, or even other metamorphic rocks that have been subjected to heat, pressure, and chemical reactions involving hydrothermal fluids and vapors. Metamorphism is associated with tectonic forces and confining pressures that cause intense strain and crustal deformation and it often accompanies folding and faulting. These processes generate changes, collectively known as metamorphism, that usually obliterate the textures, bedding, fossils, and other features of pre-existing rocks and produces new textures and minerals. The effects of heat, pressure, and intragranular fluids during metamorphism cause recrystallization of smaller mineral crystals into larger crystals, reorientation of mineral grains perpendicular to pressure fields, and/or growth of completely new minerals from old minerals through chemical ionic diffusion and/or exchange reactions that do not involve the melting of the pre-existing rock. Metamorphic rocks containing minerals with a preferred orientation are described as exhibiting foliation. New minerals grown during metamorphism are referred to as index minerals (Figure 17). Metamorphic rocks are formed when sufficient heat and pressure are applied to pre-existing rocks to alter their original textural properties and/or mineral composition and produce a suite or assemblage of unique metamorphic minerals. Figure 18 indicates some of the more common parent rocks and their metamorphic equivalents; note that the resulting metamorphic rock is dependent on the composition of the original rock and metamorphic grade to which it was subjected.
Figure 17. Index minerals and their relationship to the intensity of metamorphism.
Figure 18. Parent rocks and their relationship to metamorphic rocks.
Taken together, recrystallization and rearrangement of existing minerals and growth of new mineral assemblages are all important features that define a metamorphic facies (Figure 19); each facies is associated with a range of temperatures and pressures known as the metamorphic grade to which the original “parent” rock was subjected and can be determined from the size of mineral crystals, the type of foliation, and the index minerals observed in the metamorphic rock. Ultimately, the metamorphic grade determines what the metamorphic facies characteristics of the new metamorphic rock will be, and these characteristics are used to identify and classify various types of metamorphic rock. Two main categories of metamorphic rock are recognized, foliated and nonfoliated (Figure 20). Foliation is produced by the adjustment and recrystallization of minerals in the direction of minimum stress and generally results from the metamorphism of pre-existing rocks with a mixed mineralogy such as shale or granite. The layering can result from the alignment of platy minerals such as clay, muscovite mica, or chlorite. Increasing metamorphic grade changes the original rock from a slate (Figure 21a) to a phyllite (Figure 21b) or to a schist (Figure 21b). High grade metamorphism can generate alternating bands of felsic and mafic minerals to form a gneiss (Figure 21d). Foliation may be regular and parallel, but becomes increasingly distorted (wavy or contorted) under higher temperature and/or pressure conditions (higher metamorphic grade). Nonfoliated metamorphic rocks, on the other hand, generally result from the metamorphism of pre-existing rocks with a mineralogy dominated by a singular mineral such as pure limestone to form marble (Figure 22a), quartz-rich sandstone to form quartzite (Figure 22b), peridotite to form serpentinite (Figure 22c), and basalt to form amphibolite (Figure 22d). Metamorphism of these rocks does not produce foliation; instead, the texture is granular, resulting primarily from recrystallization of smaller mineral grains into fewer, larger crystals.
Figure 19. The relationships between depth, pressure, temperature (which control metamorphic grade), and metamorphic facies (the characteristics of the resulting metamorphic rock).
Figure 20. Classification of Metamorphic Rocks.
Figure 21. The foliated metamorphic rocks slate (A), phyllite (B), schist (C), and gneiss (D) show an increase in crystal size and ultimately banding of felsic and mafic minerals formed under increasing metamorphic grade.
Figure 22. The nonfoliated metamorphic rock marble forms during metamorphism of fairy pure limestone (A), quartzite forms during metamorphism of quartz-rich sandstone (B), serpentinite forms during metamorphism of peridotite (c), and amphibolite forms during metamorphism of basalt (d).
Metamorphism occurs under a variety of temperature/pressure conditions known as metamorphic environments. Four basic sets of conditions can prevail: 1) low T/P associated with diagenesis; 2) low T/high P associated with cataclastic metamorphism; 3) high T/low P associated with contact metamorphism; and 4) high T/P associated with regional metamorphism. Diagenesis is the result of higher temperatures and pressures related to deep burial of sediment. Cataclastic metamorphism is derived from intense shearing along fault zones. Contact metamorphism results from contact of a sediment or rock with magma and may occur as a baked zone surrounding intrusive igneous rock or just beneath an extrusive igneous rock. Regional metamorphism is a pervasive alteration of pre-existing rock related to deep burial and/or deformational mountain building, usually related to subduction at convergent plate boundaries.
Crustal Deformation
Rocks comprising the earth’s crust are continually subjected to differential (nonuniform) stresses that cause deformation and metamorphism. Deformed crustal rocks are subjected to tension (extension), compression, and shear and exhibit rupture and bending, manifested as faults and folds, the by-products of strain (changes in the size and shape of a body of rock), rotation and reorientation, and displacement. Deformation alters the physical characteristics of a body of rock, but not its composition. Taken to its logical extreme, deformation leads to metamorphism, and as we shall see, evidence of both can be exhibited within crustal rocks.
The earth’s crust is relatively rigid and brittle, and when subjected to plate tectonic motions, differential stresses applied to rock cause the accumulation of strain which can lead to bending or breaking of the rock. Differential stress occurs when rock is subject to greater tensional (extensional), compressional, or shear (scissor-like) forces in at least one dimension relative to any other. In general, tension is caused by divergent motion, compression is caused by convergent motion, and shearing is caused by transform motion. Rock can accumulate some strain without being permanently affected; the rock is said to be elastic and will rebound to into original shape if stress is relieved (stress is equivalent to strain). However, once the rate of strain becomes greater than the applied stress, the rock is permanentlydeformed; it is either bent (subject to plastic or ductile strain) by folding, or it is fractured (subject to brittle strain) by faulting. Rocks respond to stress differently, depending on their composition, density, depth of burial (and confining pressure), temperature, and the presence of intergranular fluids. Rigid, crystalline, and coarse, clastic sedimentary rocks tend to behave brittlely, they break (are subject to faulting) under stress, whereas weaker, fine, clastic or evaporite sedimentary rocks and those with intergranular fluids tend to behave more ductilely, they bend (are subject to folding) or even flow under stress. Rocks found at shallow depths (under low confining pressure) tend to fracture brittlely, minerals may be subject to diagenetic alteration, and minerals may grow to fill fractures with veins. On the other hand, deeply buried rocks (under high confining pressure) tend to bend ductilely, minerals commonly become realigned with a preferred orientation perpendicular to maximum stress, recrystallize as fewer, larger grains or form new minerals altogether.
Two geologic structures commonly form in association with crustal rocks that have undergone brittle or ductile deformation. Faults form when vertical, lateral, or oblique stresses related to tension or compression create larger-scale, through-going, brittle fractures in rock, distinctive planes of failure where adjacent blocks of the earth’s crust have been displaced. Folds develop in ductile rocks due to compressional stresses that shorten a block of the earth’s crust; the axis of a fold may exhibit fractures called joints that form due to differential strain within a layer of more brittle rock. Folds may also form in association with faulting.
Faults generally occur as one of two main types; those having predominantly vertical motion, and those having predominantly lateral motion. A fault produced by vertical motion is described with respect to the sense of offset of the top block (called the hanging wall block) relative to the bottom block (called the foot wall block). Image you are standing on the plane of a fault produced by vertical offset of two adjacent blocks (faults of this type usually form at some angle to the earth’s surface). When describing faults with a vertical component of displacement, the block below the fault plane is called the foot-wall block (the block your imaginary feet are resting on), and the block above the fault plane (hanging above your imaginary head) is described as the hanging-wall block. If the hanging wall block is observed to have moved down and away from the foot wall block, the fault is called a normal fault (Figure 23a). If the hanging wall block appears to have moved up against gravity, that is, up and over the footwall block, the fault is called a reverse fault (Figure 23b). Low angle reverse faults (those with fault planes dipping at less the 15°) are called thrust faults. Strike-slip faults result from horizontal movements of the earth’s crust. When describing the sense of displacement on a strike-slip fault, imagine that you are standing on one side of the fault plane and looking across the fault plane to the opposite block. Fault motion is described as a right-lateral (a right-lateral strike-slip fault – Figure 24a) when the block on the far side has moved to the right of your position, while fault motion is described as left-lateral (a left-lateral strike-slip fault – Figure 24b) when the block has moved to the left of your position. Note: it does not matter which side of the fault you stand on, just be sure you are looking across the fault plane to the opposite block.
Figure 23. Block diagrams illustrating normal (A) and reverse (B) faults; large arrows indicate the differential stress acting on the block, extension (A) and compression (B), while the small arrows indicate the sense of motion on the fault.
Figure 24. Block diagrams illustrating left-lateral (A) and right-lateral (B) strike-slip faults; large arrows indicate the differential stress acting on the block, right-lateral shear (A), and left-lateral shear (B), while the small arrows indicate the sense of motion on the fault.
Faults are described with respect to the orientation of the fracture plane; the strike of the fault is its compass bearing (from north), and the dip of the fault is the angle (from the horizontal) that the fault plane descends into the ground. The earth’s crust often undergoes complex motions that combine multiple orientations, resulting in oblique slip on a fault; and faults rarely occur as singular geologic features. More often, they occur as multiple fault segments within a fault zone, or a region of deformation in the earth’s crust. For example, when a series of normal faults formed side by side in a broad zone, they form horst and graben structures (Figure 25a); or, when oblique motion occurs, they can form an en echelon pattern of isolated fault segments (Figure 25b). Strike-slip faults often have kinks or bends along the length of the fault zone that generates different segments of oblique motion along the fault (Figure 26). Movement past such kinks causes portions of the fault above and below the kink to undergo extension, forming what geologist call a transtensional fault segment and local subsidence at pull-apart subsidence basins (Figure 26a), while the portion of the fault at the kink itself undergoes compression, forming a transpressional fault segment and localized uplifts (Figure 26b). Figure 27 shows a complex imbricate thrust fault zone common to the Front Range of the northern Rocky Mountains, where multiple slivers of older sedimentary rock layers have been pushed eastward over adjacent younger rock layers by compression and crustal shortening.
Figure 25. Fault zones: (A) shows horst and graben structures produced by extension and normal faulting, while (B) shows en echelon structures produced by shearing and strike-slip or oblique-slip faulting.
Figure 26. Strike-slip faults having a component of extension are called transtensional (A), while strike-slip faults with a component of compression are called transpressional (B).
Figure 27. An example of imbricate thrust faulting from the Front Range of the northern Rocky Mountains, where older sedimentary rocks have been pushed eastward by compressional shortening in a series of slices up and over younger sedimentary rocks (thrust faults are denoted by arrows).
Folds are geologic structures in which the rock units have been deformed (folded), but not broken; the fold occurs in response to ductile bending of crustal rocks subjected to compression. Imaging squeezing a floor rug from either end. The compression you exert on the rug causes it to telescope into a series of waves with alternating crests and troughs. The crests and troughs of your crumpled rug illustrate two types of folds that can form by deformation of the earth’s crust: anticlines (with a convex crest) and synclines (with a concave trough). Figure 28 describes the components of a typical fold (anticlines and synclines share many of the same features, although orientations may differ). The axial plane bisects the fold and runs parallel to the crest or hinge line. The strike of the fold axis is its compass bearing (from north), whereas its limbs dip (at angles from the horizontal) either outward from the fold axis in the case of anticlines, inward toward the fold axis in the case of synclines, or downward and away from the fold axis in the case of monoclines. The fold axis and hinge line can form more or less parallel to the earth’s surface, or they can dip into the ground at some angle to the surface (Figure 28). Folds that dip into the ground are said to plunge, and folds can be doubly-plunging (plunge at both ends).
Figure 28. The descriptive components of a typical fold: the axis, the limbs, the hinge line, and the plunge.
Anticlines are convex (upward oriented) folds in which the oldest rock is exposed in the center. Synclines are concave (downward oriented) folds in which the youngest rock lies in the center. If the axis of the fold is inclined with respect to the ground (ideally a horizontal plane), it is said to be a plunging anticline or syncline. Monoclines are simple folds with only one limb, usually developed where they overlie a reverse fault deeper in the crust. The uplifted hanging-wall block causes warping of the overlying rock layers which dip down across the buried fault on one side of the fold axis. Figure 29 summarizes the general appearance of the three main types of folds.
Figure 29. The three principal types of folds; (A) is an anticline, (B) is a syncline, and (C) is a monocline.
Very broad crustal upwarps (antiforms) that plunge in all directions from a central point are known as domes. Folds can also be described in terms of their geometry; how their limbs compare on either side of the fold axis. Fold limbs dip away from or into the fold axis and can be described as symmetric, in which both limbs dip at similar angles relative the earth’s surface, or asymmetric, in which one limb of the fold dips more steeply than the other. Folds can also be described as overturned when one fold limb is tucked back underneath the opposite limb (the crest of the anticlinal fold overlies the trough of an adjacent syncline). Recumbent folds are so overturned as to be sideways (literally laying on their sides, and in which case we don’t identify them as anticlines or synclines). Over a broad area subjected to compression, anticlines and synclines often formed in sets (not unlike the rumpled rug analogy).