Introduction

In many areas (especially in the Northern Hemisphere, and more locally the Cascade Range of Oregon), the processes of glaciation have recently greatly modified the landscape.  Glaciers are large masses of ice that form on land in areas where the winter accumulation of snow is greater than the summer melting of snow and ice.  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 (Figure 11).  Alpine (temperate) glaciers form in mountainous regions where the combination of high winter snowfall and cool summer temperatures at high altitudes permit the accumulation of snow and ice (Figure 11a).  These glaciers are confined to individual valleys or to small ice caps overlying the drainage divide between several valleys.  Generally, they are not frozen to their beds and most flow occurs by basal sliding, while accumulation and ablation are great and mass transfer is high.  Continental (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 11b).  Ice spreads outward in all directions under its own weight and is not constrained by underlying topography.  These glaciers are generally frozen to their beds except near their margins and most movement occurs by internal flow, while accumulation and ablation is sparse and mass transfer is low.  The accumulation area of a continental glacier is under cold polar conditions, but the ablation area may extend into warmer temperate conditions.  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 2.3.1

Figure 11.  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 ice system.

Glaciers as Systems

Continental glaciers never occupied central Oregon, so we will not deal with them here.  Instead, we’ll focus on alpine glacier systems which did affect the Cascade Range multiple times during the Quaternary; isolated remnants of these larger glaciers still cling to many peaks.  Glaciers form in locations where the snow that accumulates during winter doesn’t entirely melt during summer.  Figure 12 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 drift and avalanching from surrounding ice-free slopes.  The ice eventually moves into a zone of ablation, where summer melting exceeds not only snowfall, but the production 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 Cascade glaciers.  The accumulation zone is separated from the ablation zone by the equilibrium-line altitude (ELA), regionally called the snowline, a feature readily observed on modern Cascade glaciers where “fresh” snow changes to “dirty” ice.

Figure 2.3.2

Figure 12.  The components of a typical valley glacier 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 melting, 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 (Ohmura et al.; 1993), primarily related to the winter accumulation of snow and the mean summer temperature which chiefly influences the melting rates of snow and ice (Figure 13).  Healthy glaciers are those that exist in equilibrium with environmental conditions, that is, the glacier’s annual mass balance remains constant.  Yearly mass balance curves compare the input of mass to the glacier by snowfall and other means during the accumulation season (October 1st to April 1st for Northern Hemisphere glaciers), versus the loss of mass via ablation processes (most active from April 1st to October 1st) to determine the net mass balance (Figure 14).  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.  After several repeated years of positive mass, the glacier’s area 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, including those of the Cascade Range, are recording repeated yearly net negative mass balance, are shrinking, and are in retreat.

Figure 2.3.3

Figure 13.  Winter accumulation and summer melting (as determined by mean air temperature) control the mass balance of modern alpine glaciers in the Northern Hemisphere (modified from Ohmura et al.; 1993).

Figure 2.3.4

Figure 14.  A hypothetical annual mass balance curve for an alpine glacier.

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 15).  Figure 15a shows the globally averaged record of oscillations in cooling and warming for about the last million 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 15b); notice that the duration of glaciations have increased considerably in the last million years (the period of record shown in Figure 15a).

Figure 2.3.5

Figure 15.  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 (mass flux).  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 Cascade Range resulting in high rates of flow in both modern and ancient glaciers.  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.

Alpine glaciers move by a combination of internal (plastic) flow and basal sliding.  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 (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 surface slopes and bedrock surface slopes; both factors increase the affect of gravity, increase 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 (surface or bed 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.  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 12).  The spacing and size (width and depth) of crevasses increases proportional to the velocity of flow; generally forming a pattern of convex crevasses (bending up ice at the center) above the ELA and concave crevasses (bending down ice at the center) below the ELA.

The downslope movement and pressure of the overlying ice and the debris it carries cause erosion of rock at the base of the glacier (Figure 11 and 12), generally where mass flux is greatest.  Glacial erosion is primarily induced by the dual processes of plucking and abrasion (Figure 16).  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 meltwater as it 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 it flows.  Rocky debris carried in the glacial ice is eventually conveyed to the ablation zone where it melts out to form a mass of poorly sorted, unconsolidated sediment called till (Figure 11 and 12).  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 2.3.6

Figure 16.  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.

Glacial Landforms

Many features of glacial erosion are preserved in alpine environments.  Whalebacks, or roche moutonnees, asymmetrically-molded, bedrock hillocks with gentle, up-valley sides and steep, down-valley sides are an ubiquitous feature of higher elevations subject to glaciation that form by a combination of plucking and abrasion (Figure 16).  These rock surfaces are often smooth, 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, striation, and chatter-mark orientation to indicate ice flow direction.  Figure 17 depicts larger landforms produced before, during, and after glaciation.  Prior to glaciation, the landscape is cut by a typical V-shaped stream valley.  Glacial ice accumulates in the source area of a stream system, on the upper slopes of mountains, and flows down hill.  The glaciers form 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 cirque 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 2.3.7

Figure 17.  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.

Sediment deposited by glaciers creates several unique landforms (Figure 18).  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 arcuate-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 2.3.8

Figure 18.  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 upvalley; sediment reworked by meltwater streams forms outwash deposits, while fine-grained material may accumulate in proglacial lakes.

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 where glaciers can more readily nucleate, or cause the uplift of mountain ranges that 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 may cause climate change over decadal to millennial scales.  These climatic influences can be reinforced by 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 19).  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 2.3.9

Figure 19.  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.

Of course, astronomical variations in earth’s orbital conditions may have been the main “trigger” inducing glaciation, however, once cooling began, several other negative feedback mechanisms would have accelerated the trend toward glaciation (a negative feedback process is one in which the output of the process becomes an input back into the process – cooling causes more cooling).  Several feedbacks are worth mentioning, such as the albedo effect, conversion and loss of vegetation, increased dust in the atmosphere, and cooling of ocean surface waters.  The albedo effect of snow cover on the land and sea ice covering ocean surfaces would have substantially increased cooling.  Essentially, snow and ice are much more effective at reflecting incoming short-wave solar radiation than vegetation or open water.  As snow and ice coverage becomes more extensive due to cooling of the atmosphere (and expansion of the winter season), less radiation is absorbed by the earth’s surface to be converted to the heat energy that normally warms the atmosphere.  As previously mentioned, vegetative cover is good at absorbing radiation and helps drive the global heat engine, but forest vegetation is better than say grasslands for this purpose.  Therefore, as cooling occurs and forests and converted to cold-tolerant grasslands, less radiation is absorbed to be converted to heat energy.  Cooling causes some areas of the earth’s landscapes (especially those at high altitude and latitude) to lose their vegetative cover altogether which makes them susceptible to wind erosion.  These areas become atmospheric dust generators; dust in the atmosphere causes scattering and reflection of incoming solar radiation, allowing less to be absorbed by the earth’s surface and converted to heat.  As cooling accelerates, ocean surface waters cool, and one affect of this is more efficient absorption of CO2 (cool water can hold more dissolved CO2 than warm water), an important greenhouse gas that is responsible for warming the earth’s atmosphere when it absorbs heat energy released from the earth’s surface.  All of these factors and more combine to accentuate global cooling causing an maintaining earth’s deep freeze during glaciations.