The landscapes generated by plate tectonic interactions and mountain building are not static of course. 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 operating under variable climates 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 erosive agents. Forces of erosion require gravity and energy and include mass wasting, running water, glacial flow, wave action, and wind. Position matters, and uplift driven by plate tectonics and isostatic adjustment of the earth’s crust are as responsible for generating erosion as the processes themselves (in essence, what goes up must come down). Collectively, weathering and erosion are geomorphic processes that result in the deconstruction of older geologic features and the formation of new materials and landscapes.
The key to recognizing landform features developed in terrain underlain by igneous, sedimentary and metamorphic rocks lies in understanding the power of differential erosion. The chief thing to conceptualize is that some rocks are inherently more resistant to weathering and erosion than other rocks; thus, as rocks are eroded, some stand in relief as other rocks adjacent to them wear down at a faster rate. Several good generalities to consider (beware, there are always exceptions to the rule) are as follows: 1) igneous rocks are more resistant than metamorphic rocks which are more resistant than sedimentary rocks; 2) denser, more tightly intergrown crystals are more resistant to weathering and erosion, thus finer-grained igneous rocks are generally more resistant than coarser-grained rocks, and more highly altered metamorphic rocks are more resistant than less altered rocks, although foliated metamorphic rocks are generally less resistant because of inherent weakness created by foliation planes; 3) among clastic sedimentary rocks, gravel to sand sized particles are generally resistant to erosion, while silts and clays are less so, thus conglomerates and sandstones are resistant, while mudrocks are not; 4) crystalline carbonate rocks, such as limestones and dolostones, are resistant to erosion in semiarid to arid environments where they are not prone to chemical dissolution, but often weather more readily in moist environments; and 5) inherent weaknesses in the rock reduce resistance, so foliation in metamorphic rocks (as previous mentioned), joints produced by differential loading and tectonism, and bedding planes in sedimentary rock, all create avenues by which water can penetrate into rock and increasing weathering. Figure 1 displays two idealized cases involving differential erosion of igneous, sedimentary and metamorphic rocks (realize that many variations exist in nature). In Figure 1a, an igneous intrusion stands in relief against adjacent, weaker sedimentary or metamorphic rocks to form a “wall-like” feature. Felsic intrusives are generally more resistant than mafic intrusives, as are felsic lineations in foliated metamorphic rocks and would thus, stand in similar relief. In Figure 1b, stream erosion of a layer-cake of undeformed sedimentary rocks forms a stair-step-like pattern on the landscape with more resistant sandstones and limestones standing as cliffs, while less resistant mudrocks form gentle slopes.
Figure 1. Differential erosion of igneous, sedimentary and metamorphic rocks; (A) shows the relief produced during erosion of a resistant igneous intrusion surrounded by weaker sedimentary or metamorphic rock, while (B) shows the stair-step like pattern that develops through erosion of sedimentary rocks of variable resistance.
In field settings, geologists have recognized a variety of large-scale accumulations, or “bodies” of plutonic igneous and metamorphic rock. Plutonic, or intrusive igneous rocks form where magma has been injected into older, preexisting rock, to later cool and solidify and can form several types of tabular- to bulbous-shaped intrusive bodies. Metamorphic rocks require extreme heat and pressure to form, usually related to tectonic forces deep within the earth’s crust. Intrusive igneous rocks and metamorphic rocks are only exposed at the earth’s surface after prolonged uplift and erosion. Their relatively high density and interlocking crystals make for strong, resistant bodies of rock that tend to stand out in relief among weaker surrounding sedimentary country rock. Intrusive dikes are often wall-like in appearance, sills form tableland, while laccoliths, stocks, and batholiths tend to be dome-like in shape. Thus, they are most readily observed in the ancient core regions of continental interiors called “shields”, or in regions subjected to intense uplift and deformation of the crust related to mountain building. Once exposed at the surface, the crystalline nature of metamorphic rocks makes them resistant to erosion, so they tend to stand in relief (along they their igneous counterparts) on the landscape. Figure 2 describes an idealized setting where intrusive igneous and metamorphic basement rocks have been pushed upward by faulting, causing the overlying cover of sedimentary rocks to be buckled upward into a broad anticlinal fold called an anticline. Erosion proceeds with uplift, eventually exposing the metamorphic core of the mountain range. The large scale exposure of intrusives and metamorphic rocks tends to form domal-shaped protrusions where sedimentary rock has been stripped away by erosion.
Figure 2. An idealized model of uplift and erosion of sedimentary cover to expose the metamorphic core of some mountain ranges.
Sedimentary rocks accumulate as horizontal, tabular bodies in three-dimensions. Sedimentary rocks may remain undeformed (that is, unaltered) for long periods of time, remaining essentially as flat sheets on the landscape. They may also be deformed (folded and faulted) by uplift and mountain building processes, in which case they are usually observed tilted on edge. Regardless of their deformational history, sedimentary rocks are not all alike, and thus they erode at different rates. Where sedimentary rocks are flat-lying, weaker bodies of sedimentary rock tend to form extensive slopes on the landscape because they cannot remain stacked up, while stronger, more resistant bodies of sediment tend to form cliffs (Figure 1b). Remember, mudstones tend to be weaker rocks, whereas sandstones and limestones tend to be stronger; faster weathering and erosion of the weaker rocks tends to control the pace of cliff retreat of the stronger, overlying layers.
Where sedimentary rocks have undergone deformation and lay tilted at some angle relative to the earth’s surface (a common feature of fold structures exposed by erosion), the weaker rocks tend to form valleys between ridges of stronger, more resistant rocks (Figure 3). Note the asymmetry of the ridges, often called “hogbacks.” The longer, gentler slope on one side of the ridge is associated with the dip of the bedding making up the rock unit, while the steeper slope on the opposing side of the ridge cuts through and exposes the beds. This asymmetry is how geologists identify anticlines and synclines. Recall that in an anticline, bedding dips downward and away from the fold axis, so when an anticline is exposed by erosion, the hogbacks formed of resistant strata that generally parallel the fold axis are oriented such that the gentle sides of the hogback face outward. Synclines are just the opposite; bedding dips downward and into the fold axis, so when a syncline is exposed by erosion, the hogbacks are oriented such that the gentle sides of the hogback face inward. 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 and the folded rock layers will form a distinctive, rounded “nose,” with the gentle side of hogbacks facing outward on anticlines and inward on synclines. Monoclines, having only one limb, form extended lines of parallel ridges and swales that express the alternating layers of resistant versus weaker sedimentary rock units. Bedding exposed in the hogback ridges dips outward and away from the uplifted block associated with the monocline. Metamorphic rocks may be foliated (exhibit distinct layering of mineral crystals), and zones of weakness related to foliation can be exploited by erosion to produce lineations at the earth’s surface that have a resemblance to parallel ridges and valleys (similar to folded and tilted sedimentary rocks).
Figure 3. Differential erosion of deformed sedimentary rocks; in the diagram, the weaker (gray) layers are eroded faster to form valleys, whereas the stronger (brown) layers form ridges.
The main geomorphic feature produced by faulting is a ground rupture (not all faults rupture through to the surface though). When ground ruptures occur, they result in a fault scarp, an offset of the ground surface. Repeated movement on a fault may produce distinctive scars on the landscape in the form of long, linear cliff- or bench-like features that may extend for miles. Faults cause zones of weakness in the earth’s crust and ancient faults (and faults that do not produce surface rupture) exhumed by erosion are often occupied by streams, forming strangely long, linear stream valleys or series of shallow basins (sometimes occupied by lakes or playas). On standard topographic maps, fault scarps are revealed by their step-like pattern of widely spaced, then closely spaced, then widely spaced contour lines. Shading on Goggle Earth photo-mosaics, color-enhanced satellite imagery, and special shaded-relief topographic maps may reveal these linear fault scarps. Transtension may cause distinctive elongate basins to form along strike-slip faults, while transpression can cause odd, linear uplifts, often with trench-like centers.
Physical and Chemical Weathering Systems
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 4). 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 with mineral crystals. The process operates best where moisture is readily available (from rain and/or snow melt) and daily cycles of freeze-thaw occur frequently (common when the average air temperature is near freezing); thus, in the Northern Hemisphere, mid- to high latitude regions and mountainous areas are affected most. Exfoliation works on steep terrain to break up exposed rock subparallel to the erosion surface (Figure 5). Gradual removal of overlying material releases the pressure on underlying rock; the rock expands to form sheeted joints subparallel to the weathering front, which eventually break up into smaller slabs and pieces that are transported from the steep erosion surface, primarily by mass wasting. Another process, probably relatively minor volumetrically, involves thermal expansion of rock exposed to wildfires. The fire superheats the exterior of a rock exposed at the surface 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 boulder-strewn surficial deposits such as moraines are likely candidates for the process.
Figure 4. 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).
Figure 5. Physical weathering by exfoliation; removal of overlying soil, sediment, and rock from steep slopes relieves the pressure on underlying rock allowing expansion, the formation of sheeted joints, and slabbing of exposed rock surfaces subparallel to the weathering front.
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, minerals at the top are the least stable (see Figure 7 in the section of my website entitled CRUSTAL COMPOSITION AND DEFORMATION under GEOLOGY BASICS). 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 and wetter. Rates of chemical weathering are therefore generally higher where precipitation and air temperature is greater, although precipitation tends to be the more important factor during chemical weathering. Chemical sedimentary rocks and predominantly mafic to intermediate volcanic rock is especially vulnerable to chemical attack for these reasons and rapidly weathers to rich soils.
Physical and chemical weathering can work in tandem. Hydration of micas and clay minerals causes them to expand internal to a body of rock. This exerts pressure on surrounding mineral crystals which can eventually “pop” them loose from the rock by a process called granular disintegration. In more arid settings, dissolved salts can be carried into a body of rock along grain-to-grain boundaries. The salt crystals grow upon evaporation of the water and expand the fractures between crystals, again “popping” them loose by granular disintegration.
Soil itself is the chief by-product of the weathering process. As rock is broken down by physical and chemical processes, it forms a material called regolith, basically a loose mass of inorganic particles (rock fragments and mineral grains consisting of sand, silt, and clay). Regolith can be derived directly from the residuum of weathered bedrock, or it can occur as the sediment of surficial deposits transported from other locations by some agent of erosion. 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. Soil formation tends to be a function of parent material (the type of regolith), climate (warmer and wetter works best), biota (vegetation provides organic matter and soil organisms decompose it), topography (more sunlight dries out soils on southerly exposures, but relatively steep slopes tend to be more prone to erosion), and time (all thing weather with time). Figure 6 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 6. A typical soil profile developed under a forest cover; note that soils develop horizonation over time, each horizon characterized by specific features and processes.
Mass Wasting Systems
Mass wasting refers to the gravity induced movement of loose, earthy material (soil and regolith) downslope from positions of lesser stability to greater stability. Consider a block of unconsolidated soil and regolith resting on a sloping surface (Figure 7). Many environmental factors influence the possibility that this block will fail and move down slope, but each factor can be linked to two fundamental controls: normal stress, which acts on the block of material perpendicular to the slope, pressing it into the slope, is the resisting force holding the material in place, while shear stress, which acts parallel to the slope, dragging it down slope, is the driving force pulling the material downslope. Notice that as slope angle increases, more shear stress will be exerted on the block, while normal stress simultaneously decreases. Of course, mass wasting is also influenced by the properties of the block of material itself, which can be summarized as the unconsolidated material’s shear strength. The shear strength is related to the cohesion, internal friction, and effective normal stress of the soil and regolith comprising the block, while the shear stress is related to the material’s mass and the 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, when water’s cohesive properties cause particles to stick together, but the lowest effective normal stress occurs under saturated conditions, when water floats particles away from each other, 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).
Figure 7. This block diagram illustrates the two main forces controlling the mass wasting process: normal stress, acting perpendicular to the slope, presses the block into the slope and helps hold it in place, while shear stress, acting parallel to the slope, helps drag the block down the slope.
It is not difficult to conclude that many of these factors can overlap to create ideal mass wasting conditions. On naturally steep slopes of mountainous terrain, where wet soil and regolith conditions can prevail for much of the year, shear stresses can be high. Add to that the effect of stream bank erosion, or steepening of slopes by the multitude of roads built mainly for timber harvesting, and shear stresses climb even more. On the other hand, shear strength is lost by natural and human-caused forest fires and timber harvesting which reduces vegetative cover and cohesion, high rates of soil weathering and the ground shaking induced by earthquakes which reduces 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 8). Creep occurs universally, even on slopes of very low angle; the process is a result of freeze-thaw of wet soil and regolith 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, rests directly 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 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 the steep slopes, under saturated conditions, and on highly weathered soils. Lahars are a unique form of flow incorporating pyroclastic material, formed during or subsequent to a volcanic eruption, and are commonly found in volcanic terrains.
Figure 8. A classification of mass-wasting processes based on moisture content and rate of movement.
While mass wasting is a common geomorphic process acting on steep terrain, 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. The most ubiquitous feature of erosion produced by running water is the stream’s drainage basin or watershed which typically consists of four basic components (Figure 9): 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 narrow, sinuous, downslope-directed depression within the valley. Figure 9 indicates that 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 or tributary systems that coalesce in the downstream direction.
Figure 9. 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 9). 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 (known as discharge) at any moment is a product of the runoff (overland flow) to the channel and groundwater seepage into the channel (Figure 9). 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 as the water table rises during individual precipitation and snowmelt events and during wet seasons, and shrink between storms and during dry seasons as the water table falls. Streams that are dominated by runoff tend to be flashy, readily receiving and discharging their water (their discharge quickly rises to a pronounced peak after storms); whereas streams dominated by groundwater flow are more sluggish as water more gradually fills and empties from the channel (their discharge rises more slowly to a subdued peak).
From source to mouth, the gradient or slope of the stream channel decreases gradually, resulting in a concave upward longitudinal profile (Figure 10a). The geometry of the stream channel changes too, becoming wider and deeper as its gradient (slope) decreases (Figure 10b). 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. Stream profiles are affected by obstructions to channel downcutting that may be caused by influx of sediment from tributaries, a more resistant bedrock substrate, or recent faulting. Sharp changes in gradient can produce rapids or waterfalls in extreme situations (Figure 11). Waterfalls develop where a stream channel has cut through a resistant layer and is able to undercut that layer by scooping out the weaker material below through backwasting and headward erosion.
Figure 10. 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.
Figure 11. Waterfalls and rapids mark locations along a stream channel where obstructions to flow create a rapid increase in gradient; rapids often form at channel constrictions, while waterfalls are typically associated with undercutting of resistant rock units.
A stream’s longitudinal profile, and zones dominated by erosion versus deposition are ultimately controlled by its base level (Figure 11), or the lowest level to which the stream can erode. All streams become depositional upon reaching the ocean, so sea level determines ultimate base level. Sea level fluctuations brought on by climate change and tectonic uplift or subsidence can affect ultimate base level (and the longitudinal profile) (Figure 12a); a rise in sea level (or subsidence) raises the stream’s base level in the landward direction, while a drop in sea level (or uplift) lowers the stream’s base level in the seaward direction. A local base level (such as a lake or larger stream) can temporarily change a stream’s erosional regime to depositional as water makes its way to the world ocean (Figure 12b).
Figure 12. A stream’s base level, the lowest elevation to which it can erode, is determined by sea level (A), although local base level’s can exist within a watershed, such as a lake or larger stream (B).
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.
The stream’s capacity to erode and transport sediment is related to its velocity, discharge, and turbulence. A stream can erode its channel and entrain larger particles with greater velocity, and thus, velocity is essentially a measure of the stream’s erosive energy (known as stream competence). As depicted in Figure 13, channel downcutting and entrainment processes include: A) corrasion – Flowing water and air bubbles are forced into cracks, joints, and along bedding planes, acting as a jackhammer; B) hydraulic action – flowing water acts as a lift force on loosened chunks of rock and stream bed particles; C) abrasion – particles suspended in flowing water collide with the bed or bed particles abrading them like sand paper; and D) Attrition – suspended particles can also collide with each other for the same abrasive affects. Stream discharge determines the amount of sediment 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 increases. 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.
Figure 13. Processes of stream channel erosion and downcutting include: (A) Corrasion; (B) Hydraulic action; (C) Abrasion; and (D) Attrition.
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, with narrow valleys found in steep terrain generally have a straight channel pattern (Figure 14) in which the thalweg (zone of maximum velocity) 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 15). 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 14. A typical straight channel pattern formed on steep gradient streams.
Figure 15. 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.
A third form of channel pattern, braided streams (Figure 16), 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. 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 17) 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 channel is repeatedly blocked downstream and is forced to subdivide into smaller and smaller channels. Alluvial fans (Figure 17a) form where a significant break in slope occurs as a stream valley exits mountainous terrain onto adjacent lowlands, while deltas (Figure 17b) form where streams enter a slow moving body of water (such as a lake or the ocean). Both features tend to have a convex longitudinal profile and are steeper at their upper ends (although alluvial fans are much more convex and steep compared to deltas).
Figure 16. A typical braided stream channel pattern; braiding occurs during the waning stage of flood discharge when the loss of transport energy causes deposition within the channel and growth of mid-channel bars that obstruct and deflect flowing water.
Figure 17. Alluvial fans (A) and deltas (B) form where rapid loss of transport energy causes much of the stream’s sediment to be quickly deposited.
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 18). 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. Floodwaters carry sediment eroded from the channel and from gullies on surrounding slopes which is then deposited on the valley floor away from the channel as floodwaters recede and energy decreases. Coarser sediment is deposited nearer the channel, creating natural levees close to the stream and low-lying backswamps further from the stream. The floodplain is constructed over time by a combination of vertical accretion of sediment from flooding and lateral accretion of sediment from channel migration related to meander development. The floor of many older stream valleys displays raised benches of one or more levels adjacent to the modern floodplain. These features are called stream terraces (Figure 19); 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 18. The characteristics of a typical floodplain associated with a low gradient, meandering stream.
Figure 19. A cross-sectional view of 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 side-slopes.
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 20). Formation of the cliff involves wave energy that is maximized at water level or where a zone of weakness exists, such as joints, bedding planes 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 or sea level for considerable periods of time in order to develop fully (depending in part on the induration of the material subject to wave action and the wave energy available for erosion and transport of sediment).
Figure 20. A cross-sectional view of a typical wave-cut cliff and its adjoining platform.
Waves also generate depositional shoreline features (Figure 21). As a wave propagates toward shore, its base begins to drag on the lake or sea 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 offshore to form well-sorted, sandy beaches (Figure 21a). 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 21b). Beaches and spits comprised of pebbles and sand are common around the margins of abandoned proglacial and pluvial lakes in the Great Lakes region and interior west, marking stillstands (periods of relatively constant water level) at higher levels than current conditions allow. Some of these features are quite impressive accumulations of coarse, rounded sediment that seem oddly out of place in their current settings.
Figure 21. Features of a depositional shoreline: (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 22). 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. There are several types of dunes (Figure 23); 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 22. Formation and propagation of a sand dune.
Figure 23. The most common types of sand dunes include Barchan dunes (A); Parabolic dunes (B); Transverse dunes (C); and Seif, or Longitudinal dunes (D); black arrows indicate dominant wind direction(s).
If the features of wave action are often unusually located in the Great Lakes region and interior west, then those produced by wind are even more so. However, eolian (wind-generated) landforms do occasionally occur within former pluvial lake basins, some active and some fossil. Rapidly deteriorating glacial climatic conditions at the close of the Pleistocene not only caused the disappearance of most of the glaciers from the mountainous west (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 complete desiccation of most pluvial lakes as well (although several lakes in particularly deep basins or with considerable inflow of water still exist). The drying of these lakes left considerable loose, unvegetated sediment exposed to mobilization by the wind. Generally parabolic-shaped dunes commonly formed at the down-wind end of these basins, probably at the time lake-desiccation was ongoing and are now preserved under a vegetative cover. Where dunes are still active, barchans, transverse, and parabolic dunes are prevalent, depending on groundwater level and vegetation.