To understand what you’ll observe among the landscapes and rock outcrops along the roadways and hiking trails of central Oregon, the Grand Canyon region, and elsewhere, and how it got that way, a brief introduction to the discipline of geology is in order first. To begin, geology is itself a science devoted to the study of the earth system, and to a geologist the concept of a “system” implies two things: form and process. That is, geologists are interested not only in describing observable features of the earth system at all scales, but they are interested in understanding the physical, chemical, and sometimes biological processes responsible for their creation. The earth system is also very old, so by necessity, the study of geology must incorporate time, lots of time. Many earth processes and the forms they manifest require a great deal of time to develop, and indeed, many processes are cyclical and/or repeated, so that no process ever truly ends (or begins) and no form ever lasts indefinitely.
The Scientific Method
Geologists apply a logical methodology to the study of the earth. This “way of knowing” involves certain ordered steps, which when followed to their logical conclusion, result in a fundamental inference or “truth” that is acceptable to their peers, other geoscientists, as having been “proven beyond reasonable doubt”. This generality does not imply that scientific inferences are absolute and unchangeable; all scientific concepts are continually subjected to the scrutiny of the scientific method and modified as becomes necessary for their truthfulness to remain sound.
In geological study, an initial body of related observations are collected and integrated into a hypothesis, a testable statement or model about a geological feature (or features) and the process (or processes) responsible for its (or their) formation. Further geological observations are systematically obtained in the field and/or laboratory in an attempt to verify the validity of the hypothesis; its verification leads to the formulation of a theory, its falsification leads to modification of the original hypothesis and more testing. Theories are those concepts accepted as being truthful after having gone through the rigorous “rite of passage” that is the scientific method; certain universally accepted concepts may even become scientific principles or laws.
Uniformitarianism is a fundamental inference in geological investigations that has reached the lofty status of a principle. In essence, this concept can be summarized as “the present is the key to the past”. In other words, when a geologist observes a form resulting from a process, it can be surmised that that form has always resulted from that process (or that process always results in that form), and thus, geological features preserved at the earth’s surface or within the earth’s crust can be used to infer past geological events and/or their environments of formation. This concept could be applied to the active processes of plate tectonics such as volcanism or crustal deformation, or to active geomorphic processes such as glaciation or mass wasting that have resulted in the present-day configuration of the landscape. Uniformitarianism is a truth; however, certain changes to the earth system can occur so rapidly, so catastrophically, and/or so infrequently, such as meteorite impacts and mass extinction events that a strict use of the concept breaks down. Certain processes and their resultant forms occurred so long ago that prevailing physical, chemical, and most certainly biological conditions were not the same as today, and in these cases, uniformitarianism also breaks down.
This then begs the dual questions: what is geologically old versus young, and how do geologists determine this? The answer to the first question is that Earth has gotten older, and its age less approximate, the longer geologists have studied it (we now believe the earth is about 4.6 billion years old); the answer to the second question is the same for any science, qualitatively and quantitatively (or more accurately, by use of relative and absolute dating techniques). In relative dating, the goal is to unravel a sequence of geological “events” that may include the formation of a rock body, or its alteration (removal by erosion; deformation by faulting, folding, or intrusion of magma; and/or metamorphism by heat and pressure), but the actual age of the event is unknown. In absolute dating, the objective is to actually date the timing of the event which helps make interpretation of the sequence of events it is related to more accurate.
Determining the relative age of a geological “event” requires the determination of its proper place or position within a chronological sequence that includes all of the events found at a given locality, even though the actual ages of events are unknown. The order of events and the relative passage of time can be deduced from several universal stratigraphic principles related to rock bodies and geological structures (formulated by geologists Nicholaus Steno and William Smith):
a) Principle of Original Horizontality – Sedimentary rocks (and more broadly, the volcanic lava flows and pyroclastic deposits often discussed in this guidebook) are initially deposited in subhorizontal layers conforming to the topography they are laid down upon, therefore tilted strata indicates deformation and the passage of time necessary to complete that deformation.
b) Principle of Superposition – In a sequence of undeformed sedimentary and/or volcanic rocks, the oldest rocks are at the base, becoming progressively younger toward the top.
c) Principle of Lateral Continuity – A layer of sedimentary and/or volcanic rock is initially deposited as a broad, continuous sheet, its pattern disrupted only by obstructions (landforms) that occur at higher positions and/or by subsequent erosion.
d) Principle of Cross-Cutting Relations – A rock body must exist before it can be altered, therefore strata altered by another intruding rock body or disrupting geological structure must be older than that altering event.
e) Principle of Inclusions – A rock body that contains fragments of another rock type must be younger than those fragments.
f) Principle of Faunal Succession – In a sequence of sedimentary rocks, changes in fossil content (fossil assemblages) occur systematically upward even though the rock type may not change; the changes imply the passage of time. This same concept can be applied to plant fossils as well.
Sedimentary (and often volcanic) rocks are initially deposited in subhorizontal layers. When an interruption in the process of deposition occurs, geologists can infer that some amount of time is missing from the rock record, presumably induced by some geological event or events that affected the area where the rocks where forming. This event or events could involve simple nondeposition, or involve much more extensive alteration related to intrusion, uplift and deformation, and removal by erosion. The break in the sequence, and its implied gap in time, is referred to as an unconformity. The time passage and degree of alteration indicated by the interruption in the deposition of sediment can be interpreted from a characterization of the type of unconformity (Figure 1):
a) Disconformity – Layers of subhorizontal sedimentary or volcanic rock separated by a nondeposition or erosion surface; implies the shortest time gap and the least alteration. Figure 1a illustrates the formation of a disconformity.
b) Angular Unconformity – Younger sedimentary or volcanic layers deposited horizontally over older strata that was initially deposited horizontally, then tilted by deformation and beveled by erosion; implies an intermediate amount of time and alteration. Figure 1b illustrates the formation of an angular unconformity.
c) Nonconformity – Younger rock layers deposited horizontally over the eroded surface of older intrusive igneous and/or metamorphic “crystalline” basement rock; implies the greatest time gap and the most significant alternation. Figure 1c illustrates the formation of a nonconformity.
Figure 1. The formation and appearance of geological unconformities: (A) disconformities; (B) angular unconformities; and (C) nonconformities.
To measure the absolute passage of time and the precise age of a geological event, geologists must have a process that occurs at a constant rate and a means of keeping a cumulative record of that process. All elements found in minerals (minerals make up rocks) have isotopes; atoms with a variable number of neutrons in the nucleus. Some of these isotopes are unstable under the conditions of pressure and temperature where they occur in surface soil or regolith or within the earth’s crust. With passage of time the nuclei of such isotopes break down spontaneously, emitting subatomic particles and energy as radioactive decay, and becoming altered to the isotopes of new elements.
All radioactive isotopes decay at specific rates called the half-life of the element. Half-life is the time it takes for half of the original volume (weight %) of the unstable, radioactive isotope (called the parent isotope) to decompose to a more stable isotope (called the daughter isotope). The time period remains the same for each iteration of volume change (weight % change). The volume change can be measured if the original volume of the parent and daughter isotopes can be assumed with accuracy. Half-life for relatively large samples (it doesn’t take much when considering the size of atoms) is constant for a given element’s isotope and is only affected by geologic processes that subject the original material containing the parent isotopes to enough heat and pressure to result in the chemical alteration of the original material, thereby resetting the radioactive “clock”. Thus, barring chemical alteration, the longer a suitable geological material containing radioactive isotopes exists (since its formation), the less parent isotopes and the more daughter isotopes there will be (that is, the ratio of parent to daughter isotopes will decrease). This process is illustrated in Figure 2.
Figure 2. A graphical illustration of a typical decay curve for a radioactive isotope; the shape of the curve is determined by the isotope’s half-life, or the time it takes for half of the original volume (weight %) of the unstable, radioactive isotope (called the parent isotope) to decompose to a more stable isotope (called the daughter isotope).
Many radiometric dating techniques exist for measuring absolute time, each is applicable to dating materials of varying age depending on the half-life of the radioactive isotope contained in the sample and used in the analysis. In general, parent isotopes with long half-lives can be used to date older geological events, but not younger events; and parent isotopes with short half-lives can be used to date younger geological events, but not older events. This is because the longer the half-life of the isotope, the less the amount of parent isotope has decayed to daughter, and the shorter the half-life, the greater the amount of parent isotope has decayed to daughter. In the former case, the less time that has passed, the more difficult it is to measure the change in the ratio of parent to daughter. In the latter case, the longer the time that has passed, the more difficult it is to measure any parent isotope remaining.
Two absolute dating methods have been routinely applied to determine the age of sediments and volcanic materials (and the events that they represent). The Carbon-14 (14C) dating technique involves the radioactive decay of the unstable 14C isotope into stable Nitrogen-14 (14N). Since carbon is a common constituent of all living things, plants and animals accumulate a small percentage of 14C over their lifetimes. When the organism dies, it ceases to accumulate carbon, and the 14C decays to 14N at a fixed rate known as its half-life (about 5730 years for 14C). The relatively rapid decay rate of 14C and the miniscule amount of the isotope present (about 98.8% of the carbon stored in an organism is 12C), limits the usefulness of this method to dating events less than about 75,000 years in age to as little as a few hundred years. Determining the ratio of 14C to 14N reveals the number of half-lives that have passed since the organism died. For example, assume that a tree was buried in sediments by the advance of a glacier; long after glacier retreat, the tree is recovered from a streamside bank cut in the glacial sediments and subjected to 14C analysis. A sample from the tree’s bark is determined to have an average ratio of 14C to 14N that indicates only 12.5% of the parent (14C) remains, and thus, three half-lives have passed since death and burial of the tree (after one half-life the ratio would be 50%, after two it would be 25%, after three it would be 12.5%, and so on……). The passage of three half-lives indicates that the glacial advance occurred about 17,190 years ago (3 x 5730 years).
The Potassium-40 (40K) to Argon-40 (40Ar) radiometric dating technique can be readily applied to materials that contain minerals such as orthoclase feldspar and mica; fortunately, these minerals are abundant in lava flows and pyroclastic deposits produced by volcanism. 40K simultaneously decays to 40Ar and 40Ca, but this geological dating method only monitors the change to 40Ar because 40Ca is a stable isotope found abundantly in all potassium-bearing minerals even before the decay process begins (and therefore, it would be considerably more difficult to determine how much 40Ca resulted from the decay of 40K). 40K has a much longer half-life than 14C (1.25 billion years in fact!) and is more abundant in volcanic materials, thus, this method can be used to date geological events as young as about 5,000 years and as old as many billions of years. Let’s assume that a volcano erupted at some time in the past and filled an old stream valley that had been previously carved into the landscape with a succession of lava flows. Samples of rock containing the appropriate minerals are collected from the basal lava flow, and subjected to the 40K to 40Ar dating technique. The mineral samples are determined to have an average ratio of 40K to 40Ar that indicates 99.994% of the parent (40K) remains, and thus, only 0.00006 half-lives have passed since the basal lava flow was emplaced. This age indicates that the eruptive event that produced the succession of lava flows began about 75,000 years ago (0.00006 x 1.25 billion years). It also suggests that carving of the stream valley occurred at least that many years ago, a minimum age for this erosive event.
The age of a geological event is therefore estimated by some combination of the following: 1) material contained within the rock or sediment of interest is dated and this date is used to infer the age of the rock or sediment (and the event that produced it); 2) material contained within a rock or sediment underlying or overlying the feature of interest is dated and used to infer a maximum or minimum age, respectively, for that feature (less precise than #1); 3) material contained within a rock or sediment altered by some event can be dated and used to infer a minimum age for the altering event (the material had to exist to become altered); and/or 4) material contained within a rock or sediment underlying or overlying another unit altered by an event can be dated to provide a maximum or minimum age for that altering event (less precise than #3).
Using these stratigraphic “tools” of relative and absolute dating, geologists have been able to correlate sedimentary and volcanic rock units in isolated locations, often over vast distances. One such correlative tool uniquely suited to volcanically active regions past and present is the use of tephra (volcanic ash) units. Each tephra unit has a unique chemical signature (no two volcanic ashes are exactly alike), and so changes in tephra upward within a volcanic or sedimentary sequence, although not systematic, also indicate the passage of time. When a tephra unit is found at a given location, its chemistry can be determined and tied in to known tephras for that area. Applying Lateral Continuity, when enough locations have been determined, the aerial distribution and fallout pattern of that tephra unit can be mapped and its eruptive source revealed. Once the age of the tephra has been determined radiometrically, it can be tied into the overall regional tephra chronology, and the age of other rocks or sediments associated with it can also be inferred.
The ultimate result of all this age-dating via relative and absolute methods has been the correlation of geological materials and events from one location to another and the construction of a global geological time scale (Table 1). The rock bodies and deformational events contained within the earth’s crust have been dated, correlated, and tied to specific time intervals: eras, periods, and epochs. In central Oregon, the principles of relative dating might be applied to the differentiation of the sequence of events represented by a series of interlayered, overlapping lava flows, pyroclastic deposits, and lacustrine or fluvial sediments. The actual age of these geologic units could be determined by a combination of 14C and 40K-40Ar dating in order to further define the age relationships of these materials.
Table 1: Globally Standardized Geological Time Scale
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 associated with mountain building and crustal deformation, as well as major sea level alterations, erosive events, and basin filling events. Figure 3 displays a simplified model of the earth’s interior. This model was developed by geologists over many years from the synthesis and interpretation of a great deal of physical, chemical, and biological data. Simply note that the earth is composed of concentric layers of material with differing thickness, composition, and density. 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 (Figure 4) within a deeper, hotter, mushy-plastic layer (the upper mantle or asthenosphere).
Figure 3. A simplified model displaying the concentric layers of material with differing thickness, composition, and density that comprise the earth’s interior.
Figure 4. 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 basically spherical, and because it is not expanding, the tectonic forces of mantle convection result in three possible interactions between these plates. 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 5). As the plate fragments first begin to move apart, the plate material lifts, fractures, and thins, and then sags into the developing gap. Over time, the plate fragments rift 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 that is itself progressively buried by layers of sediment (Figure 5).
Figure 5. 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 (Figure 6). Plate collisions are complicated by the composition and density of the colliding material. When plates of differing density collide, the less dense plate (associated with either continental crust or younger, warmer, more oceanic crust) is forced up and over, while the denser plate (associated with older, colder oceanic crust) dives down and under in a process called subduction (Figure 6). The downgoing crustal slab and overlying sediments are partially melted, allowing magma to rise upward through the overlying plate, while the remainder is reincorporated into the upper mantle. When plates are of similar composition and density, neither subduct, instead they crumple together and compete for “stacking” rights. Collisions form long, linear mountain belts of deformed crust and/or solidified molten rock.
Figure 6. 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 side-by-side (Figure 7). 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 association with divergent boundaries (Figure 5); essentially “taking up the slack” as two adjacent, diverging plates pull apart on a plane curved in three dimensions.
Figure 7. Shearing and transform plate motion; note the offset of the cinder cone.
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 8). (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). 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.
Figure 8. A mantle plume model, showing the postulated source of magma and associated volcanism trailing away from the current hot spot in the direction of motion of the overlying plate.
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 crustal rocks. In contrast, volcanic glass, or obsidian 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 9 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 9. 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. 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 9).
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 10). 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 forms 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, and the cycle is renewed. However, 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 tectonics, 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 most often a result of tectonics.
Figure 10. 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. As indicated earlier, solidification of magma results from the crystallization of silicate minerals. The silicate minerals that chiefly form igneous rock can be divided into three groups, mafic, intermediate, and felsic (Table 2) based on the dark, salt-and-peppery, or light color they exhibit (which is more directly related to their elemental content). Thus, as we shall see, color becomes an important tool for igneous rock identification.
Table 2: Igneous Rock Classification
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. The newly generated magma is lighter than the surrounding rocky material and it rises toward the surface. Magma can cool and crystallize at depth or when it is extruded as lava onto the earth’s surface. The texture of a rock sample is determined by the groundmass (background crystals) 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 (Table 2) characterized by a coarse-grained groundmass of crystals that are large enough to be seen with the unaided eye. If the rising magma reaches near the surface or is extruded at the surface it will cool rapidly and form extrusive (volcanic) igneous rock (Table 2). Volcanic igneous rocks typically have either no crystals and a glassy appearance (obsidian) or a fine-grained groundmass of crystals with a dull, earthy appearance, and minerals that can only be identified with the aid of a microscope. The rising magma may partially cool and crystallize at depth before being extruded as lava at the surface, in which case, the 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 andesite porphyry.
Geologists thus classify igneous rocks on the basis of mineralogy and texture, the mineralogy controlled by the magma composition and the texture controlled by the rate of cooling of the magma. Table 2 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.
Volcanic eruptions can release copious amounts of ash and rock fragments 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 source). The material may be loosely consolidated into a tuff (fine grained) or volcanic breccia (coarse grained), or strongly indurated into a banded, welded tuff if the ashy material was sufficiently hot as it landed. The term tuff or welded tuff is tacked on as a suffix to the volcanic igneous rock name, such as rhyolite tuff. The outpouring of lava and pyroclastic material at the surface is often associated with much escaping gas. If gas bubbles become trapped in volcanic rocks, they take on a porous appearance, and the rocks can be described as having a vesicular texture. Extremely porous, or even frothy, volcanic rock forms scoria (mafic and fine grained) or pumice (felsic and glassy). When the volcanic rock is only moderately porous, the term vesicular or scoriaceous is tacked on as a prefix to the rock name, such as vesicular basalt or scoriaceous basalt.
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 has led to the use of the terms intrusive (plutonic) and extrusive (volcanic) to describe the origin of igneous rocks. Both kinds of igneous rocks occur as larger “bodies” of igneous rock with certain unique characteristics (Figure 11). These bodies include:
1) Intrusive igneous bodies (plutons)
1. Batholith – large masses of igneous rock with no known floor that may represent singular or multiple, coalescing chambers of solidified magma.
2. Stock – same as batholith, but smaller than 40 square miles, may represent an isolated portion of a larger batholith.
3. Dike – tabular mass or sheet that cuts across pre-existing “country” rocks at an angle.
4. Sill – tabular sheet injected between parallel rock layers.
5. Laccolith – dome-shaped mass (a mushroom-shaped sill) intruded into layered sedimentary rocks.
6. Lopolith – basin-shaped mass intruded into layered sedimentary rocks.
7. Volcanic neck – solidified volcanic vent that remains standing after the surrounding cone has been eroded away.
2) Extrusive igneous bodies
1. Volcanic cones – conical structures composed of lava and/or pyroclastic material extruded from a central vent.
a. Cinder cone – small, steep-sided cone, principally composed of cinders.
b. Shield volcano – large, gradually sloping cone, composed largely of mafic lava flows.
c. Strato (composite) volcano – moderately large, graceful cone, composed of alternative layers of ash and lava.
2. Lava flows – sheets of molten lava extruded from a vent which run down the flanks of the cone and spread out at its base.
3. 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 11. Some examples of 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” (is the least stable at the pressure/temperature conditions of the earth’s surface). 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 (Figures 9 and 10). 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 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. Sediments are later lithified (buried, compacted, and cemented together) to become sedimentary rock. Clastic sedimentary rocks are classified by the size of the particles found within the rock and the dominant mineral composition (Table 3). Rock containing gravel sized particles is called a conglomerate if the grains are rounded and a breccia if the grains are angular. If a clastic rock containing abundant clay is fissile (shows close parallel layering) it is called shale instead of claystone. The terrestrial settings of the western U.S. provide abundant depositional environments for clastic materials which tend to be relatively coarse-grained because of short transport distances and the limited amount of abrasion the sediment has undergone. They are also commonly interlayered with lava flows and/or pyroclastic deposits. Clastic rocks can form primarily by the deposition of terrestrial biological debris (plant fragments) such as peat and coal, in which case they are usually classified with the biochemical sedimentary rocks. Peat is distinguished from coal by the loose, porous, poorly cemented nature of its constituent material.
Table 3: Clastic Sedimentary Rock Classification
Table 4 shows a simple classification for biological, chemical, and biochemical sedimentary rocks. Biological detritus that accumulates in oxygen- poor environments (often poorly circulated water) can form peat or coal. Chemical sedimentary rock 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 then settle to the floor of a shallow marine shelf, coastal lagoons and tidal flats, or playas. Massive limestones composed of tiny calcite crystals and evaporite rocks such as rock salt (halite crystals) and rock gypsum (gypsum crystals) form in this way. Crystals can also be indirectly precipitated from saltwater as the hard parts (fossils) of formerly living marine organisms, such as fossiliferous limestone which is formed as an aggregate of many fossil fragments which may not be recognizable due to erosion. Chemical or biological silica in the form of very fine grained (cryptocrystalline) quartz can also precipitate. It often collects as siliceous ooze on the ocean floor which is incorporated within calcareous sediment to eventually form chert nodules or beds. These primarily marine-derived sedimentary rocks are not found in central Oregon. Silica can also form the hard parts of shelled organisms such as diatoms; several playas and shallow lakes in central Oregon contain diatom-rich deposits.
Table 4: Biological, Biochemical, and Chemical Sedimentary Rock Classification
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. These processes generate changes, collectively known as metamorphism, that usually obliterate the textures, bedding, fossils, and other features of pre-existing rocks and produce new textures and minerals. The affects 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 reactions that do not involve the melting of the pre-existing rock. The types of minerals found depend on the pre-existing rock type that has undergone metamorphism as well as the degree of metamorphism known as metamorphic grade (low, medium, or high).
Metamorphic rocks are classified on the basis of whether or not they exhibit a strong preferred orientation of mineral grains into planar layering, or foliation, by the size of their constituent crystals, and by the types of minerals present (Table 5). Foliation is produced by the adjustment and recrystallization of minerals in the direction of minimum stress. The layering can result from the alignment of platy minerals such as mica or chlorite (schistose rocks), or from alternating bands of felsic and mafic minerals (gneissic rocks). Layering may be regular and parallel, but becomes increasingly distorted (wavy or contorted) under higher temperature and/or pressure conditions (higher metamorphic grade). Foliated metamorphic rocks generally result from the metamorphism of pre-existing rocks with a mixed mineralogy such as shale or granite. Nonfoliated metamorphic rocks, on the other hand, generally result from the metamorphism of pre-existing rocks with a predominantly singular mineralogy such as sandstone, limestone, or basalt. 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.
Table 5: Metamorphic Rock Classification
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 permanently deformed; it is either bent (subject to plastic or ductile strain) by folding, or it is fractured (subject to brittle strain) by faulting. Usually, rock is more likely to behave in a brittle fashion (rather than ductile) if it is relatively cool, occurs at shallower depths (is under less pressure), or is more densely lithified or crystalline (composed of closely packed grains or interlocking crystals).
Faults are fractures in bedrock along which the earth’s crust has moved (Figure 12). They 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 (hanging wall block) relative to the bottom block (foot wall block) (Figure 12a). If the hanging wall appears to have moved downward relative to the foot wall, the fault is called a normal fault. If the hanging wall appears to have moved up against gravity, the fault is called a reverse fault. Low angle reverse faults (those with fault planes dipping at less the 15°) are called thrust faults. Faults showing lateral motion are called strike-slip faults and are described with respect to the direction of motion of theopposite block relative to an observer standing on one side of the fault (Figure 12b). Thus, a right lateral fault indicates the rocks on the opposite side of the fault moved to the right and a left lateral fault indicates the rocks on the opposite side of the fault moved to the left.
Figure 12. The three principal types of faults; normal and reverse faults result from vertical offsets of the earth’s crust (A), while strike-slip faults result from lateral offsets (B).
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 (Figure 13). 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 13a); and when strike-slip or oblique-slip fault segments are laterally offset along a linear fault zone, they form an en echelon pattern (Figure 13b).
Figure 13. 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.
Folds are geologic structures in which the rock units have been deformed (folded), but not broken (Figure 14). 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. Folds are described with respect to the orientation of the fold axis and limbs as labeled on Figure 14. 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.
Figure 14. The three principal types of folds; (A) is an anticline, (B) is a syncline, and (C) is a monocline.
Agents of Landform Modification (Geomorphic Change)
The landscape generated by plate tectonic interactions and mountain building is dynamic. Many forces of weathering and erosion combine to tear down what has been made, in effect, beginning the rock cycle anew. Weathering refers to a host of physical and chemical processes that mechanically break down rock into finer and finer particles, as well as altering its chemical composition to form minerals more stable under earth’s surface conditions. These processes occur in situ, that is, they involve little or no motion of the material. On the other hand, processes of erosion demand motion, picking up material in one place, transporting it, and redepositing it elsewhere, often over multiple cycles involving different agents. Erosive forces require gravity and energy and include mass wasting, running water, wave action, glacial flow, and wind. Collectively, weathering and erosion are geomorphic processes that result in the demolition of older geologic features and the formation of temporary new ones.
Physical (mechanical) weathering uses the application of naturally induced differential stresses to break apart rock. The chief source of stress is the freezing and thawing of water that seeps into fractures and pore spaces within the rock (Figure 15). Water expands upon freezing which can create a jackhammer-like effect at the macro-scale of bedrock joints and bedding planes, and micro-scale of grain-to-grain boundaries and within-grain fractures, twinning planes, and other zones of weakness. The process operates best were moisture is readily available (from rain and/or snow melt) and daily cycles of freeze-thaw occur frequently (where average temperature is near freezing); thus, in the Northern Hemisphere, the upper-mid-latitude region and mountainous areas are affected most. This process is prominent in central Oregon’s Cascade Range and higher portions of the High Lava Plains. Another process, relatively minor volumetrically, involves thermal expansion of rock exposed to forest-fires. The fire superheats the rock exterior causing minerals to expand, but the poor thermal conductivity of rock limits the penetration of the fire’s heat inward, setting up differential stresses that spall layers from the rock surface like peeling an onion. Bedrock exposures and the boulder-strewn moraines of the drier forest country east of the Cascade Crest are especially prone to this process.
Figure 15. Physical weathering of rock by frost wedging; cyclical freeze-thaw gradually widens fractures and pore spaces, prying rocks apart (small arrows); eventually allowing gravity to take over and pull loose blocks downslope (large arrows).
Chemical weathering of rock is primarily a function of two factors: mineral instability and climatic conditions. As previously discussed, the stability of a silicate mineral is related to its crystallization temperature and pressure; the higher the conditions, the more unstable is the mineral at earth’s surface. Therefore, when looking at Bowen’s Reaction Series, mafic minerals at the top are the least stable (Figure 9). Water plays a significant role in chemical weathering because it provides a medium in which chemical reactions occur more readily. This is mainly because water is never a pure substance in nature; it contains dissolved ions which carry charge. In this way, the ions in the water react with adjacent minerals in a process generally known as ion exchange, resulting in the chemical alteration of the preexisting mineral to more stable clay minerals. Chemical reactions also occur more rapidly when the chemical medium is warmer, although plentiful moisture is the predominant control. Rates of chemical weathering in the western U.S. are therefore greater in mountainous areas and along the Pacific Coast where moisture levels area highest.
Soil itself is the chief product of the weathering process. As rock is broken down by physical and chemical processes, it forms regolith, a loose mass of inorganic particles (rock fragments, sand, silt, and clay). Regolith becomes soil only after the addition of organic matter. Plants and animals colonize regolith, die, and decay to provide organic matter. Over time, organic materials build within the “top soil” to provide vital nutrients for further colonization, and the variety and density of organisms grows. Weathering and decomposition continue to add inorganic and organic components to the developing soil as it thickens and becomes more pronounced. Figure 16 displays a typical well-developed soil profile formed under a cover of forest vegetation. The O, A, and E horizons are generally considered the “top soil” where maximum weathering and downward translocation of material occurs, the B horizon represents the zone of accumulation (of the weather products from above), while the C and R horizons represent parent material, which may be unconsolidated sediments and/or a bedrock residuum.
Figure 16. A typical soil profile developed under a forest cover.
Mass wasting refers to the gravity induced movement of loose, earthy material (soil and regolith) downslope from positions of lesser stability to greater stability. Many environmental factors influence the mass wasting process, but each can be linked to two fundamental controls: the shear strength of the material (the resisting force holding the material in place), and shear stress (the driving force pulling the material downslope). The shear strength of an unconsolidated material is related to its cohesion, internal friction, and effective normal stress, while the shear stress is related to the material’s mass and slope angle on which it lies; decreasing shear strength and/or increasing shear stress will cause greater instability. Cohesion is principally a function of binding agents such as living roots and decayed organic matter present in the soil, so loss of cohesion occurs when vegetation is removed (by forest fires and/or timber harvesting). Internal friction is related to the abundance of coarse, angular grains within an earthy material, characteristics that enhance their potential to interlock; loss of friction occurs with greater weathering and an increase in clay content, and by shaking the material (such as may occur during an earthquake or volcanic eruption). Effective normal stress results from the water found in the material, the highest effective normal stresses occurring under moist conditions, but the lowest under saturated conditions, such as may occur after intense rainfall or during seasonally wet periods. Mass is most commonly enhanced by the addition of water from rain and/or snowmelt, while slope angle can be increased when slopes are undermined by streambank erosion or the formation of gullies (or by the building of roadways).
In the western U.S., the naturally steep slopes of mountainous terrain and the naturally wet conditions that prevail for much of the year, combine to enhance shear stress. Add to that the multitude of roads built mainly for timber harvesting, and shear stresses climb even more. Shear strength is lost by natural and human-caused forest fires and timber harvesting which reduces vegetative cover and soil cohesion, by high rates of soil weathering in moist climates, by the significant potential for earthquakes and volcanic eruptions which reduce internal friction, and the naturally wet conditions that saturate earthy material and reduce effective normal stress.
Mass wasting occurs most commonly by gentle, persistent creep, and as infrequent, but more catastrophic rock falls, debris slides and flows, generally categorized on the basis of water content and rate of movement (Figure 17). Creep occurs universally, even on slopes of very low angle; the process is a result of freeze-thaw of wet soil during cool periods and/or hydration-dehydration of clay minerals in soil. Rock falls literally involve blocks of falling rock that have been loosened from cliffs and other steeply sloping exposures. The process usually involves physically prying the rocks loose by frost action and often results in the construction of a talus cone, a conical mound of debris stacked against the base of the cliff. Slides occur as a coherent mass that slips along a distinct failure surface such as a bedding plane oriented downslope, or where a mass of soil and regolith, such as glacial debris, rests on bedrock. Flows involve loose, earthy material that deforms and moves downslope as an incoherent, plastic mush. Flows can be slow to fast, depending on water content; earth flows containing the less water and being the slowest, debris flows with more water and faster still, and mud flows containing the greatest amount of water and flowing most rapidly. Flows often begin as coherent slide blocks, begin to break up as they move down slope and incorporate water, and as they enter streams, become mud flows in the end. Flows are abundant features on steep slopes and highly weathered soils. Lahars are a unique form of flow incorporating pyroclastic material, formed during or subsequent to a volcanic eruption, and commonly found in volcanic terrains.
Figure 17. A classification of mass-wasting processes based on moisture content and rate of movement.
While mass wasting is a common geomorphic process of the mountainous western U.S., the influence of stream erosion is all-pervasive. In fact, stream erosion is the most significant process of landform change in any terrestrial environmental setting, regardless of climate or geology. Stream channels cannot erode downward (to a lower topographic position) indefinitely. Erosion is limited by base level (Figure 18), or the minimum energy surface of the stream system (a surface below which deposition of sediment carried by the stream ensues). Commonly, base level is equated to the mean sea level position (Figure 18a), but it can be related to local base level surfaces such as lakes (Figure 18b) and/or local equilibrium surfaces associated with river systems (such as smaller streams emptying into larger ones). Ultimate base level is a dynamic surface controlled by isostatic and tectonic adjustments of the earth’s crust and to eustatic changes in sea level. Thus base level can be lowered (and stream erosion enhanced) by crustal uplift or sea level fall and base level can be raised (and stream deposition enhanced) by crustal subsidence or sea level rise. Streams will continue to erode a landscape as long as their channel remains above base level.
Figure 18. The base level of a stream system; sea level is the ultimate base level of all streams (A), although temporary base levels, such as a lake, can exist within a stream system (B).
The most common feature etched into earth’s landscapes is the stream drainage basin or watershed (Figure 19). A watershed consists of four basic components: 1) the stream channel; 2) the adjoining banks, floodplain and valley slopes; 3) the hydrologically connecting groundwater system or aquifer; 4) and the surrounding drainage divide. The stream channel is any permanently flowing water exposed at the earth’s surface and confined to a distinctly sinuous, downslope-directed pathway within the valley. The stream channel is merely the surface expression of the water table, that is, the position within the watershed where the groundwater system intersects with the land surface. The size of a given watershed is determined by the position of the drainage divide, the highest elevation on the terrain, such that all water entering the stream valley flows downslope to the channel and down channel to a singular exit point (either a larger stream, a lake basin, or ultimately the world ocean). Thus, the most fundamental watershed is one that contains a singular path or channel of flowing water conjoined with its surrounding slopes that also serve as the drainage divide. Larger drainages of greater complexity contain multiple smaller watersheds that coalesce in the downstream direction.
Figure 19. A typical watershed; consisting of the stream channel, the adjoining banks, floodplain and valley slopes, the hydrologically connecting aquifer, and the surrounding drainage divide. Flow of water within the watershed is depicted by arrows.
The stream channel consists of a permanently submerged bed, even at low water, and banks that are submerged during variable high water conditions (Figure 19). Duration of submergence is readily observed by examining the type and density of riparian vegetation (plants that thrive in or tolerate periodically wet conditions) growing on stream banks. The quantity of water within the channel at any moment is a product of the runoff (overland flow) to the channel and groundwater seepage into the channel (Figure 19). Runoff occurs when the ground’s infiltration capacity has been exceeded, infiltration being controlled by the size and interconnectedness of pores within soil and regolith and by antecedent moisture conditions (the amount of water in the pores prior to a precipitation or snowmelt event). During a storm, infiltration is quickly reduced by raindrop splash as soil particles plug surface pores that had been formed by earthworms and other larger or smaller burrowing organisms during interceding dry weather. Runoff fills small depressions on the slope to form a thin film of sheetwash that enters rills and gullies (nonpermanent channels) downslope, eventually flowing into streams. Water that is infiltrated makes its way by gravity drainage to the water table and enters the groundwater system. The water table separates unsaturated soil pores above from saturated pores below and naturally rises toward the ground surface during storms and/or wet seasons. Groundwater drains downslope through saturated soil and regolith and seeps into streams from beneath the channel’s surface. Where the groundwater table intersects with the earth’s surface, it becomes the stream channel surface. Thus, the stream channel is not a static feature of the landscape; its dimensions swell during individual precipitation and snowmelt events and during wet seasons, and shrink between storms and during dry seasons. Streams that are dominated by runoff tend to be flashy, readily receiving and discharging their water; whereas streams dominated by groundwater flow are more sluggish as water more gradually fills and empties from the channel.
The geometry of the stream channel changes downslope, becoming wider and deeper as its gradient (slope) decreases (Figure 20). The composition of the channel bed and banks also changes downstream as bedrock exposures are replaced by unconsolidated sediment and the size of the particles making up the channel bed and banks decreases. Water flowing within the channel is quantified by its velocity (flow distance divided by time), discharge (the volume of water passing a given point divided by time or the cross-sectional area times velocity), and turbulence (the degree of mixing within the water column, generally exhibited by its frothiness). In a downstream direction, the stream channel constantly adjusts its shape in an attempt to maintain velocity even though the channel slope is decreasing. Discharge increases downstream as the width and depth of the channel increases (width increasing faster than depth); while turbulence decreases downstream in response to less and less water in contact with the channel bed and banks.
Figure 20. Changes in stream channel geometry downstream within the stream system; (A) describes a decrease in gradient, and (B) describes a change in channel cross-sectional area.
The stream’s capacity to erode and transport sediment is related to its velocity, discharge, and turbulence. A stream can entrain larger particles with greater velocity, and thus, velocity is essentially a measure of the stream’s erosive energy (known as stream competence). Stream discharge determines the amount of material that can be transported (known as stream capacity). The stream’s turbulence influences the size of particles that can be suspended within the flowing water. As turbulence decreases downstream, the size of suspended particles also decreases; however, because discharge increases downstream, the total volume of sediment transported by the stream actually goes up. Streams carry sediment by three transport mechanisms: bed load, suspended load, and dissolved load. Bed load refers to sediment transported by a combination of rolling and skipping of particles along the bed of the stream channel. Sediment transported by suspended load refers to particles that remain floating within the water column by turbulence. Dissolved load is simply ions dissolved into the water directly, micro-“particles” that can be carried indefinitely in streams. The amount of material transported by bed load decreases downstream, while the suspended load increases downstream in response to loss of turbulence.
The pattern exhibited by the stream channel and the shape of the valley floor, form in response to these factors. Relatively small, higher gradient streams found in steep terrain typical of mountainous areas have a straight channel pattern (Figure 21) in which the thalweg (zone of maximum velocity) tends to zigzag back and forth from bank to bank. Where the thalweg impinges on the bank, erosion produces a scour pool and coarse, gravelly lag on the bed and a cutbank sloping into the channel. The opposite stream bank displays a point bar related to accumulation of fine sediment on the channel floor where velocity is low. Downstream of the pool and cutbank, within the channel, a riffle forms where gravelly lag accumulates to obstruct stream flow, especially during low-flow conditions. Relatively large, lower gradient streams found in gentle terrain typically exhibit a meandering channel pattern (Figure 22). The movement of the thalweg from bank to bank defines the channel, as accentuated pool and cutbank erosion alternates with point bar deposition to form large, meander loops. Meanders can become so accentuated by lateral and down-gradient migration that they eventually erode into each other forming cutoffs. Segments of the channel are lopped off; the ends closest to the new channel becoming filled with sediment, such that the abandoned channel forms a crescent-shaped lake separated from the active channel known as an oxbow.
Figure 21. A typical straight channel pattern formed on steep gradient streams.
Figure 22. A typical meandering channel pattern formed on a gentle gradient stream. Meander loops tend to migrate laterally and down-gradient within the valley (as indicated by the arrows and dashed lines), and eventually, adjacent loops may migrate into one another creating cut-offs.
As stream gradient decreases, the stream-valley relief becomes increasingly subdued. Valley slopes retreat from the channel margin and become lower, while the valley floor progressively exhibits a larger and more developed floodplain (Figure 23). A floodplain forms adjacent to the stream channel in response to periodic flooding of the valley floor when ample runoff causes discharge to exceed the channel’s capacity to store water. Meandering of the main stream channel generates cutoffs and oxbow lakes. Floodwaters carry sediment eroded from the channel and surrounding banks onto the floodplain and deposit it as energy decreases away from the stream margin and as floodwaters recede. Coarser sediment is deposited as a channel lag and nearer the channel as streambank deposits, creating natural levees close to the stream and low-lying backswamps further from the stream. Backswamps often contain older oxbow lakes and Yazoo streams that flow parallel to the main channel because they are found in this swaley topography between the natural levees and the valley side-slopes. The floodplain is constructed over time by a combination of vertical accretion of overbank deposits from flooding and lateral accretion of channel deposits from channel migration related to meandering of the channel. Slope colluvium washed from valley side-slopes often intertongues with overbank deposits along valley margins.
Figure 23. A model of a meandering stream system’s floodplain and valley sedimentary fill.
A third form of channel pattern, braided streams (Figure 24), occurs where stream discharge and the sediment load it carries fluctuates widely, where the stream banks are poorly vegetated and/or composed of coarse, unconsolidated material, and/or where the stream channel’s gradient (and energy) rapidly decreases; these features are common elements of steep terrain and arid to semiarid conditions pervasive over much of the western U.S. In such a stream, the water flows in a braided pattern around gravelly islands and mid-channel bars, dividing and reuniting as it flows downstream. During periods of high discharge, the entire stream channel may contain water, with the islands covered to become submerged bars; some of the islands could erode, but the sediment would be re-deposited as the discharge decreases, forming new islands or bars. Islands may become resistant to erosion if they become inhabited by vegetation. Alluvial fans and deltas (Figure 25) are a variation of braided stream pattern that forms where rapid loss of transport energy causes sediment to be deposited in a broad apron as the stream subdivides into several smaller channels. Alluvial fans (Figure 25a) form where a significant break in slope occurs as a stream valley exits mountainous terrain onto adjacent lowlands, while deltas (Figure 25b) form where streams enter a slow moving body of water (such as a lake or the ocean). The floor of many older stream valley systems displays raised benches of one or more levels adjacent to the modern floodplain. These features are called stream terraces (Figure 26); they represent alternate periods of cool, moist and warm, dry climate. Cool, moist climatic conditions enhance erosion in upland areas and the capacity of a stream to transport sediment to lowland areas where it is deposited as an aggrading floodplain. When climatic conditions become warmer and drier, erosion in the uplands decreases, sediment supplied to the lowlands decreases, and a period of stream channel incision ensues, cutting a trench into the former floodplain. A return to cool, moist conditions results in construction of a new floodplain at a lower position within the valley, leaving remnants of the old floodplain as terraces along the valley margins.
Figure 24. A typical braided stream pattern forms where channel gradient is steep, stream discharge rapidly fluctuates, and/or stream bank materials are unconsolidated and poorly anchored.
Figure 25. Alluvial fans (A) and deltas (B) form where rapid loss of transport energy causes much of the stream’s sediment to be quickly deposited.
Figure 26. A typical low-gradient stream valley displaying the active channel and adjacent flood plain with successively older, bench-like stream terraces perched on the valley slopes.
Processes and landforms associated with wave action is certainly significant to landscape development along the Pacific margin of the western U.S. During the recent Ice Age, several large pluvial lakes formed in the block-faulted basins of the interior Great Basin region, so interestingly enough evidence of their former is still prominently displayed in many places where wave action produced both erosional and depositional features in the form of wave-cut cliffs and shoreline-related beaches and spits. Waves are generated by the transfer of kinetic energy from moving particles of air to moving particles of water. As wind blows across an open reach of water, such as a lake surface, its energy is transferred to the water by frictional drag, forming waves. The size and spacing of the waves is related to the strength of the wind and the distance of open water. Waves that impinge on a shoreline in turn transfer much of their energy to the shore, ably pile-driving into rock and regolith, prying apart fractures and loosening individual clasts and grains, breaking them down into smaller fragments, and entraining and transporting loose particles along the shore, abrading them further as they go.
Waves can effectively erode rock or consolidated sediment, forming a wave-cut cliff and adjacent platform or terrace (Figure 27). Formation of the cliff involves wave energy that is maximized at water level or where a zone of weakness exists, such as a bedding plane or less-indurated material. Wave action creates a notch in the eroding material and an overlying, unstable block that eventually breaks loose from the cliff face into the crashing waves, where it is quickly disaggregated and its particles carried away. The wave-cut platform develops in the low-energy zone just below wave base; sediment removed from the cliff is transported off-shore and deposited as a blanket on the older, deeper portion of the terrace. Wave-cut cliff and platform features form a distinctive step and riser gouged into an otherwise uniform slope. They require a stable lake level for considerable periods of time (depending in part on the induration of the material subject to wave action) in order to develop fully.
Figure 27. A cross-sectional view of a typical wave-cut cliff and its adjoining platform.
Waves also generate depositional shoreline features (Figure 28). As a wave propagates toward shore, its base begins to drag on the lake bottom, and the wave crest rises and gradually outpaces the wave base. Eventually, the wave collapses of its own weight, dissipating its energy on the shoreline, and moving particles up and back in its swash and backwash. This process abrades clasts and smaller grains alike, and has a winnowing effect, leaving coarser particles behind and moving finer grains off-shore to form well-sorted, sandy beaches (Figure 28a). Wave action also generates nearshore currents that run parallel to the shoreline, dragging particles along as longshore drift, depositing grains in quieter water of the nearshore as sand bars, and foreshore, modifying beaches and forming spits that grow laterally with time across embayments (Figure 28b). Beaches and spits comprised of pebbles and sand are common around the margins of pluvial lakes in south-central Oregon; marking pluvial-lake stillstands (periods of relatively constant water level). Some are quite impressive accumulations of coarse, rounded sediment that seem oddly out of place in their current water-barren, semi-arid setting.
Figure 28. (A) displays a cross-sectional view of a typical beach; while (B) shows a spit formed of beach material transported laterally along a shoreline by longshore drift.
The energy associated with blowing wind is strong enough to entrain fine sand, silt, and clay. The sand fraction is too coarse to be blown far and accumulates as dunes at the downwind end of a basin, while silt and clay is suspended for much longer distances and is usually removed from its basin of origin. The typical dune has an asymmetric form, with a gentle upwind (stoss) side and a steep downwind (lee) side (Figure 29). Sand grains are rarely picked up and suspended far; instead, they are blown up the stoss side of the dune by saltation in a series of rolls, hops, and bounces. When sand grains reach the dune crest, depending on wind strength and sand grain size, they saltate and/or cascade down the slip face (the dune’s lee slope) forming cross-bedded sand or are carried a short distance beyond the dune to accumulate as subhorizontal layers of interdune sediment. High energy conditions erode into the lee side of the dune, while the calm air on the lee side of the dune results in deposition of particles which are buried by the advancing slip face material.
Figure 29. Formation and propagation of a sand dune.
There are several types of dunes (Figure 30); their formation is determined by several factors, including wind speed, consistency of wind direction, sand supply, and the anchoring affects of vegetation and shallow groundwater. Three dune types occur where winds are generally unidirectional (from one consistent direction). If wind strength is high and/or sand supply is low (often related to a high water table), barchans dunes form with their horns pointing in the direction of the wind. Moderate winds and sand supplies usually generate transverse dunes, their crests perpendicular to wind direction; while the anchoring affects of vegetation cause parabolic dunes, with their horns pointing upwind. When winds are particularly strong from two directions with the same general orientation (say, northwest and southwest), seif or longitudinal dunes form with their crests parallel to wind direction. If winds are multidirectional or diverted by numerous obstructions, often a phenomenon of smaller, enclosed basins or against mountain barriers, star dunes may form with several prominent, radiating crests.
Figure 30. Common types of sand dunes: A – Barchan dunes; B – Parabolic dunes; C – Transverse dunes; and D – Seif, or Longitudinal dunes. Black arrows indicate dominant wind direction(s), sand grain, and dune migration.
Eolian (wind-generated) landforms do occasionally occur within the former pluvial lake basins of the Great Basin, as well as in several locations in the arid Southwest where ample sandy sediment is available from eroding sandstone bedrock. Rapidly deteriorating glacial climatic conditions at the close of the Pleistocene not only caused the disappearance of most of the mountain glaciers of the western U.S. (the ones currently found at higher elevations probably formed during one or more neoglacial episodes such as the Little Ice Age of the late Holocene), but resulted in the desiccation of many pluvial lakes and ephemeral streams as well (although several lakes in particularly deep basins or with considerable inflow of water perennial still exist, such as the Great Salt Lake). The drying of these lakes and stream channels left considerable loose, unvegetated sediment exposed to mobilization by the wind. Generally, parabolic-shaped dunes formed at the down-wind end of lake basins, braided streams, and alluvial fans, probably at the time desiccation was ongoing and are now preserved under a vegetative cover of sagebrush. Where dunes are still active, barchans, transverse, and parabolic dunes are prevalent, depending on groundwater level and abundance of vegetation.