In many high latitude and alpine areas of the world (especially in the Northern Hemisphere) the processes of glaciation have recently greatly modified the landscape. Glaciers are large masses of ice that form on land in areas above what geologists call the snowline, where the winter accumulation of snow is greater than the summer melting of snow and ice. Figure 1 shows a simple schematic of the snowline globally from pole to pole and across the United States from west to east. Glacial systems are essentially giant natural conveyor belts, with snow and ice conveyed from the upper end of the glacier to the terminus, eroding, transporting, and depositing rock and sediment as they move. Globally, two major types of glaciers occur. Alpine (or temperate) glaciers form in mountainous regions where the combination of high winter snowfall and cool summer temperatures at high altitudes permit the annual accumulation of snow and ice (Figure 2a). These glaciers are confined to individual valleys or to small ice caps overlying the drainage divide between several valleys when ice spreads downslope under its own weight. Generally, they are not frozen to their beds and most flow occurs by basal sliding, while accumulation and ablation (melting) are great and mass transfer is high. Continental (or polar) glaciers form at high latitudes and may cover extensive areas, even near sea level, because temperatures are relatively cold throughout the year and very little of the snow that accumulates ever melts (Figure 2b). Ice spreads outward in all directions under its own weight and is not constrained by underlying topography (ice can even flow upslope over significant topographic obstacles). The accumulation area of a continental glacier is under cold polar conditions, but the ablation area may extend into warmer temperate conditions. These glaciers are generally frozen to their beds except near their margins and most movement occurs by internal plastic flow, while accumulation and ablation is sparse and mass transfer is low. Some continental glaciers exhibit zones of high flow rates and mass transfer, correlated with topographically controlled areas where the glacier bed is saturated by water. Note the contrast in scales between alpine and continental glaciers; the size and extent of glacial erosional and depositional features will consequently show a similar variation in.
Figure 1. Global scale and regional scale (United States) snowlines; the global snowline is primarily controlled by air temperature (cooler temperatures toward the poles lower the snowline), while the regional U.S. snowline is more controlled by westerly air flow (moist air masses moving in from the Pacific Ocean lower the snowline along the west coast relative to the continental interior and east coast).
Figure 2. Diagrams depicting alpine (A) and continental (B) glaciers showing generalized zones of accumulation and ablation separated by the ELA and their relationship to zones of erosion and deposition within the glacial system.
Glaciers as Systems
Continental glaciers never occupied more the than the upper fringes of the western United States (the focus of this website), so we will not deal with them here. Instead, we’ll focus on alpine glacier systems which did affect the many of the mountainous areas of the west multiple times during the Quaternary. Isolated remnants of these larger glaciers still cling to many peaks, although global warming is driving many to extinction. Glaciers form in locations where the snow that accumulates during winter doesn’t entirely melt during summer (above the snowline – Figure 1). Figure 3 displays a typical valley glacial in an alpine setting. As snow falls in a zone of accumulation, it becomes compacted and recrystallized in the presence of meltwater and by its own weight to form glacial ice, and begins to flow downslope in response to gravity. Accumulation also occurs via wind drifted snow and by avalanching from surrounding, steep, ice-free slopes. The ice eventually moves into a zone of ablation, where summer melting exceeds not only winter snowfall, but the production or mass flux of ice from the source area (the accumulation zone). Ablation may also occur by calving of icebergs directly into sea water; a common process related to tidewater glaciers such as those of Glacier Bay, Alaska, although not to former or current glaciers in the western conterminous Untied States. The accumulation zone on an individual glacier is separated from the ablation zone by the equilibrium-line altitude (ELA) (Figure 3), regionally equivalent to the snowline, a feature readily observed on modern living glaciers where “fresh” snow changes to “dirty” ice by the end of summer.
Figure 3. The components of a typical valley glacial in an alpine setting.
The ELA represents the position on the glacier’s surface where the amount of winter accumulation is equal to the amount of summer ablation (Figure 3), both factors determining the “health” or mass balance (accumulation versus ablation) of the glacier system. The mass balance of modern alpine glaciers worldwide exists within a narrow envelope of climatic conditions, primarily related to the winter accumulation of snow and the mean summer air temperature at the ELA which chiefly influences the melting rates of snow and ice (Figure 4). Healthy glaciers are those that exist in equilibrium with environmental conditions, that is, the glacier’s annual mass balance remains constant (annual accumulation and ablation are matched). Yearly mass balance curves compare the input of mass to the glacier by snowfall and other means during the accumulation season (generally October 1st to April 1st for Northern Hemisphere alpine glaciers), versus the loss of mass via ablation processes (most active from April 1st to October 1st) to determine the net mass balance (Figure 5). Note that ablation occurs throughout the year and above the ELA. Not all of the snow that falls in the accumulation zone will be converted to glacial ice. The conversion process is diagramed in Figure 6 and involves several steps, or phases, including the accumulation of fresh snow, which is compressed by the weight additional snow from above, and then recrystallized into granular shapes (the points of snowflakes under higher pressure are melted and the water redistributed to reform as ice closer to the “core” of the snowflake), while meltwater from above percolates into the pores between ice granules and refreezes, gradually filling in the empty spaces between grains of firn to form solid ice. The density increases during the conversion process from about 0.1 g/cm3 to about 0.9 g/cm3 (the density of water is 1.0 g/cm3 by comparison) (Figure 6). Once enough mass has accumulated, the glacial ice will deform plastically and flow downslope much like Silly Puddy, transferring mass into the ablation zone.
Figure 4. Winter accumulation and summer melting (as determined by mean air temperature) control the mass balance of modern alpine glaciers in the Northern Hemisphere (data from Ohmura et al., 1993).
Figure 5. A hypothetical annual mass balance curve for an alpine glacier; the negative mass balance is a product of greater annual ablation than accumulation.
Figure 6. A generalized description of the ice crystal shape, consistency, and density changes associated with the conversion of snowflakes into glacial ice.
The ELA of a glacier with an annual positive mass balance will drop in elevation, while a negative mass balance would correspond to a rise of the ELA. After several repeated years of positive mass balance, the glacier’s thickness and areal coverage will grow and it will advance downslope. Conversely, years of negative mass balance would lead to shrinkage and upslope retreat. At the present, most alpine glaciers worldwide are recording repeated yearly net negative mass balance, are shrinking, and are in retreat, due to global warming.
Applying the concept of uniformitarianism, it is not difficult to see how a global (or regional) cooling trend could induce glacier growth as the conditions supporting glaciation expand into lower latitudes and altitudes (annual mass balance becomes dominantly positive during expansion). Figure 7a shows the globally averaged record of oscillations in cooling and warming for about the last 900,000 years as determined from oxygen-isotope ratios collected from foraminifera and radiolarians found in marine sediment cores. A similar pattern of oscillations has been inferred from greenhouse gases trapped in ice cores collected from the Greenland and Antarctic ice sheets. By convention, oxygen-isotope stages associated with glaciations are given even numbers. The most recent glacial maximum occurred about 18,000 years ago, whereas the glacial maximum prior to that occurred about 130,000 years ago, during oxygen-isotope stages 2 and 6. Examination of longer records of benthic marine oxygen-isotopes indicates that the world transitioned slowly into its most recent “Ice Age” over a period of about one and half million years, beginning in the late Pliocene about three million years ago (Figure 7b); notice that the duration of glaciations have increased considerably in the last million years (the period of record shown in Figure 7a).
Figure 7. Global paleo-oscillations in glacial cooling and interglacial warming based on oxygen-isotope ratios collected from marine fossils; (A) depicts the record for roughly the last million years, while (B) shows the gradual onset of the earth’s most recent Ice Age over the last five million years.
The rate of flow of the ice downvalley is dominantly controlled by the slope at the glacier base, the ice surface slope, and the amount of accumulation and ablation (which controls the mass flux through the ELA). In alpine valley glaciers, movement is dominantly controlled by the slope at the glacier base and a high rate of mass flux; the steep slopes and high volumes of accumulation and ablation in the Oregon Cascade Range resulting in high rates of flow in both modern and ancient glaciers; by contrast, modern and ancient Rocky Mountain glaciers have (had) lower rates of mass flux and ice flow. Alpine ice caps form where several valley glaciers coalesce at a low divide or when glaciation is severe and persistent; in which case, ice surface slopes may be steep enough to allow flow over areas of relatively flat lying topography, and even uphill across intervening topography (evidence of such flow can be found in the Yellowstone region of Wyoming and Montana where a massive ice cap extended over much of the Yellowstone Plateau and surrounding mountain ranges).
Alpine glaciers move by a combination of internal (plastic) flow and basal sliding (Figure 8). Burial of older ice by younger ice and fresh snowfall causes it to deform and flow downhill under the force of gravity. Individual ice crystals deform as internal planes of shear slide past each other. Differential stress between ice crystals causes recrystallization and realignment of grains parallel to each other and to the minimum stress plane (the bedrock surface slope) which increases the ease of plastic deformation and sliding of grains past one another. The pressure of overlying ice, frictional drag, and geothermal heating from below may cause melting at the base of a glacier. This water acts as a lubricant between the ice and the glacier bed which causes basal sliding of the glacier. Glaciers flow relatively slowly (as compared to streams); however, the maximum velocity occurs in the central portion of a glacier due to frictional drag along the bed and valley walls. Ice velocities are higher in glaciers with steep ice surface slopes and bedrock surface slopes; both factors increase the effect of gravity, increase basal shear stress, and decrease friction between ice crystals (increasing internal flow) and between the glacier and its bed (increasing basal sliding). Glaciers with a high rate of mass flux through the ELA (high accumulation and ablation) have greater flow velocities as well. The dominant direction of flow is down slope (ice surface and/or bedrock slope), although flow also generally occurs downward into the glacier in the accumulation zone and upward through the glacier in the ablation zone as ice is transferred past the ELA. Bedrock “stairs” created by resistant rock units underlying the glacier, create zones of localized extending (faster) and compressing (slower) flow within a glacier (Figure 3). The surface layer of a glacier is brittle because it is often cooler, under less pressure, and has less liquid water content than the subsurface. When ice flows, this upper layer forms cracks perpendicular to the direction of movement called crevasses (Figure 3). The headwall crevasse, called the bergschrund, indicates that the glacier is pulling away from the bedrock headwall. The spacing and size (width and depth) of crevasses down ice increases proportional to the velocity of flow; and crevasses tend to concentrate above zones of extending flow within a glacier. Generally, a pattern of convex crevasses (bending up ice at the center) occurs above the ELA and concave crevasses (bending down ice at the center) occurs below the ELA. The presence of a bergschrund and crevasses are an excellent “quick and dirty” way of judging whether or not an ice mass is still flowing, and can thus still be considered a glacier. Figure 9 presents a photograph of the Bend Glacier perched on the northern flank of Broken Top Volcano in the central Oregon Cascades. The glacier exhibits crevasses perpendicular to its slope which indicates that at the time, this ice mass was still flowing and could still be considered a glacier.
Figure 8. The movement of ice within an alpine glacier is primarily determined by the internal plastic flow of ice and a meltwater lubricant at the glacier bed that causes basal sliding.
Figure 9. The Bend Glacier as photographed in 2009; at the time, the ice mass could still be considered an active glacier based on the presence of crevasses on its surface.
Many features of glacial erosion are preserved in alpine environments. The downslope movement and pressure of the overlying ice and the debris it carries causes erosion of rock at the base of the glacier (Figure 1), generally where mass flux is greatest. Glacial erosion is primarily induced by the dual processes of plucking and abrasion (Figure 10). Plucking occurs in response to regelation flow caused by changes in ice pressure above and below a rocky “bump” at the base of the glacier. Pressure-induced melting of flowing ice upstream of the bump provides meltwater that travels over and around the bump to accumulate in fractures on its downslope side where it refreezes. Expansion of the ice as the meltwater refreezes causes breakup of the rock into small chunks that are subsequently “plucked” up into the flowing ice. Rocky debris eroded in this manner, in addition to other material which falls onto the surface of the alpine glacier from the weathering and mass wasting of surrounding valley walls, becomes incorporated into, and transported by, the glacial ice. This material provides the glacier with “tools” of abrasion as the rock fragments impinge upon each other and the rock surface over which the ice flows. Rocky debris carried in the glacial ice is eventually conveyed to the ablation zone (Figure 1) where it melts out to form a mass of poorly sorted, unconsolidated sediment called till (Figure 3). Thus, glaciers represent open systems involving the transfer of mass (ice and sediment) and energy (the kinetic energy of the glacier’s motion, and the solar, geothermal, and frictional energy that melts the ice from above and below) from place to place.
Figure 10. Diagrams displaying plucking and abrasion, the dual processes primarily responsible for glacial erosion: (A) describes regelation flow; (B) describes the relationships between plucking, abrasion, and regelation flow; and (C) indicates small-scale landforms of glacial erosion related to plucking and abrasion.
Whalebacks, or roche moutonnée, asymmetrically molded, bedrock hillocks with gentle, up-valley sides and steep, down-valley sides are a ubiquitous feature of higher elevations subject to glaciation that form by a combination of plucking and abrasion (Figure 10). The up-ice bedrock surface of the roche moutonnée is often smoothed by abrasion, exhibiting the very fine polish of sand and silt grains carried in ice, linear scratches called striations that are grooved into rock by larger abrasive tools, and more rarely chatter-marks (crescentic-shaped “chips” with horns pointing up ice) formed by the chisel-affect of some abrasive tools. Geologists use a combination of roche moutonnee, striations, and chatter-mark orientation to indicate ice flow direction; Figure 11 shows a recently deglaciated surface near the Collier glacier in the central Oregon Cascades that exhibits polish and striations indicating ice flow to the northwest.
Figure 11. A recently deglaciated bedrock surface near the Collier Glacier in the central Oregon Cascades exhibits glacial polish and striations; the striations indicate an ice flow direction downslope to the northwest.
Figure 12 depicts larger landforms produced before, during, and after glaciation. Prior to glaciation, the landscape is cut by a typical V-shaped stream valley (Figure 12a). Glacial ice accumulates in the source area of a stream system, on the upper slopes of mountains, and flows downhill (Figure 12b). After deglaciation (Figure 12c), the glaciers have formed amphitheater shaped depressions at the head of drainages called cirques, while the former V-shaped stream valleys have been eroded into characteristic U-shaped cross-sections. Hanging valleys form where small U-shaped valleys are tributary to a larger U-shaped valley. The tributary glacier in the smaller valley had less erosive power than the larger valley glacier (often called the trunk glacier), and thus, it could not carve its floor as deeply during glaciation, leaving the tributary valley “high and dry” after deglaciation. The topographic high from which several cirques and/or valley glaciers emanate is called a horn. The knife-edge ridges that separate two parallel glacial valleys are called arêtes, whereas a col is a low point or saddle in a ridge separating two glacial valleys that face in opposite directions. Hanging valleys may occur where tributary valley glaciers once joined larger valleys. These tributary valleys are left hanging above the main valley because the volume of ice, and hence glacial erosion, was less significant (not as deep). Tarns are small lake-filled bedrock basins, often found in cirques, that formed by glacial erosion. Interconnected tarns in a stream valley are referred to as paternoster lakes (resembling a string of prayer-beads). Figure 13 offers two views of the cirque basin scoured from the northeast face of Three Fingered Jack in the central Oregon Cascades; the cirque forms a compound amphitheater-shaped depression tucked up against the peak proper, and extending along the Cascade Crest to the north, with an inner and outer set of moraines. The cirque itself feeds into the upper end of a longer U-shaped valley.
Figure 12. Glaciation of alpine terrain results in a unique suite of erosional landforms; (A) depicts a typical alpine watershed developed by stream erosion prior to glaciation, (B) that same watershed during the height of glacial activity, and (C) post-glacial alteration of that watershed.
Figure 13. A compound cirque basin on the northeast side of Three Fingered Jack in the central Oregon Cascades; (A) shows a view up valley into the cirque, and (B) shows a view downvalley transitioning from cirque into U-shaped valley.
Sediment deposited by glaciers creates several unique landforms (Figure 14). Till lodged along the glacier’s bed forms hummocky or stream-lined ground moraine, while till dumped along the margins of the glacier may accumulate in arccuate-shaped ridges called end moraines. Moraines deposited along the valley slopes are referred to as lateral moraines, while moraines formed at the lower end of the glacier are called terminal moraines. Medial moraines form where the lateral moraines of two tributary valley glaciers merge into the same valley; these are rarely preserved after deglaciation. Recessional moraines may be located upvalley from terminal zones. These features represent an end moraine deposited during a stillstand or slight readvance of a glacier when climatic conditions temporarily stabilized during overall deglaciation. Till deposited by a glacier may later be transported and sorted by braided, meltwater streams only to be redeposited as relatively flat lying layers of coarse, gravely to sandy sediment called outwash. Valley floors are often filled with flat-lying deposits of glacial outwash called valley trains which accentuate the U-shaped cross-section of these formerly glaciated valleys. During low-flow conditions, exposed sediment may be picked up by wind and carried down into lowland areas to accumulate as silty loess. Outwash occurring within mountain stream valleys may form multiple stepped benches indicating periods of meltwater input associated with glaciation, alternating with periods of erosion associated with warmer, drier interglacial conditions. In certain locations, water may be trapped behind moraines in small proglacial lakes, where fine sediment accumulates in banded layers called varves. Kettles are small, rounded depressions in till or outwash formed by the melting of isolated blocks of ice that became detached from the glacier margin as the ice mass retreated, some of which are now filled with groundwater, forming ponds or small lakes. These features are only rarely found on in alpine settings. Moraines are generally not well preserved in alpine glacial environments because of the steep terrain and rapid rates of erosion; this is particularly true of the western slope of the Cascade Range where a relatively warm and very wet climate prevails. However, the eastern slope of the Cascades lies in a climatic rainshadow, and the much drier conditions have allowed the preservation of a much more detailed record of glaciation. Figure 12 shows two sets of end moraines, the “inner,” sharp-crested, generally uneroded moraines depicted in the photographs were deposited during the Late Holocene Little Ice Age, while the “outer,” moraines, more subdued by erosion, and blanketed by conifers, are associated with the older, Late Pleistocene Cabot Creek advance of the Suttle Lake glaciation.
Figure 14. The sediment deposited by glacier systems generates several unique landforms: sediment deposited directly from melting ice forms till that accumulates as lateral and terminal moraines along the glacier margins, as well as recessional moraines and ground moraine as the glacier retreats up valley; sediment deposited from meltwater forms outwash, either directly in contact with ice to form kames and eskers, or beyond the ice margin by meltwater streams to form coarse-grained valley train and/or in proglacial lakes to form fine-grained deposits.
Causes of Glaciation
As we have seen, cyclical fluctuations in climate from warm to cool and back again cause glaciations and interglaciations. Therefore, the question really becomes: what causes climate change? Several factors influence climate over different temporal and spatial scales, including: 1) plate tectonic motions that may concentrate continental masses at high latitudes or cause the uplift of mountain ranges that create terrain where glaciers can more readily nucleate and/or alter patterns of air circulation, thus causing climate fluctuations over millions of years; 2) astronomical variations in the path of the Earth in its orbit around the sun and in the orientation of the earth’s axis cause slight variations in the amount of radiant energy reaching the earth and climatic fluctuations over tens of thousands of years; and 3) various short-term climatic cycles such as sunspot activity and pulses of volcanic activity that affect the amount of radiant energy the earth receives from the sun) and may cause climate change over decadal to millennial scales. These climatic influences can be reinforced by concurrent changes in greenhouse gas concentrations and dust in the earth’s atmosphere that result from climate’s influence on the amount and distribution of vegetation, or by changes in earth’s albedo caused by expansion or contraction of ice and snow cover; factors known as “positive feedbacks” that accelerate pre-existing trends in climate change.
Astronomical variations, otherwise known as the Milankovitch Cycles, after their discoverer, Milton Milankovitch, are considered by most geologists to play the chief role in causing the glacial and interglacial periods of the current Ice Age (Figure 15). The shortest of these variations, the precession of the equinoxes (or simply precession), involves changes in the earth’s axial tilt with respect to its orbit around the sun and has a cycle ranging from of 19,000 to 24,000 years. Essentially, the earth is wobbling like a top over time as in orbits through space; currently the North Pole points toward the North Star, but about half way through the cycle it will point toward Vega. The earth’s obliquity, or the tilt of its axis with respect to the plane of the ecliptic (the orbital plane of the planets in the solar system), changes on an intermediate, 41,000-year cycle. Currently, earth’s axial tilt is about 23.5°, but it is decreasing toward its minimum of 21.1° (which will take about 8,000 years). The total range of axial tilt is between 24.5° and 21.1°. The longest of these variations is the earth’s eccentricity, or the change in the shape of earth’s orbital path about the sun. The orbital path changes by about 6% on a cycle of 100,000 years, fluctuating between a slightly more and slightly less elliptical shape. In general, when each of these cycles is maximized (maximum wobble away from the sun, maximum tilt, and a maximum elliptical shape to its orbit), the earth’s seasonality increases which enhances the likelihood of glaciation. Therefore, if two or even all three of these maximum phases occur simultaneously, the possibility of earth’s decent into another glacial period is all the more probable.
Figure 15. A diagram describing the astronomical variations known as the Milankovitch Cycles that have played a chief role in causing the glacial and interglacial periods of the current Ice Age: (P) the precession of the equinoxes, caused by changes in the earth’s axial tilt direction with respect to its orbit around the sun on a cycle of 19,000 to 24,000 years; (T) the earth’s obliquity, caused by changes in the tilt of the earth’s axis with respect to the plane of the ecliptic on a cycle of 41,000-year cycle; and (E) eccentricity, caused by the change in the shape of earth’s orbital path about the sun on a cycle of 100,000 years.