Geomorphic Systems of Green Lakes Valley
Geomorphic Systems of Green Lakes Valley
Abstract and Keywords
There are at least three justifications for the examination of the geomorphology of the area in which ecosystem studies are conducted. First, the present landscape and the materials that make it up provide the substrate on which ecosystem development occurs and may impose constraints, such as where soil resources are limited, on that development. Second, the nature of the landscape and the geomorphic processes acting on it often define a large part of the disturbance regime within which ecosystem processes occur (Swanson et al. 1988). Third, the processes of weathering, erosion, sediment transport, and deposition that define geomorphic dynamics within the landscape are themselves ecosystem processes, for example, involving the supply of resources to organisms. In this last context, it is noteworthy that drainage basins (also called watersheds or catchments) were recognized as units of scientific study during a similar time period in both geomorphology and ecology (Slaymaker and Chorley 1964; Bormann and Likens 1967; Chorley 1969). The drainage basin concept, the contention that lakes and streams act to integrate ecological and geomorphic processes, remains important in both sciences and underlies the studies in Green Lakes Valley reviewed here. Over the past 30 years, Niwot Ridge and the adjacent catchment of Green Lakes Valley have been the subject of much research in geomorphology. Building on the studies of Outcalt and MacPhail (1965), White (1968), and Benedict (1970), work has emphasized the study of present-day processes and dynamics, especially of mass wasting in alpine areas. These topics have been reviewed by Caine (1974, 1986), Ives (1980), and Thorn and Loewenherz (1987). Studies of geomorphic processes have been conducted in parallel with work on Pleistocene (3 million to 10,000 yr BP) and Holocene (10,000 yr BP to present) environments in the Colorado Front Range (Madole 1972; Benedict 1973) that have been reviewed by White (1982). This chapter is intended to update those reviews in terms that complement the presentation of ecological phenomena such as nitrogen saturation in the alpine (chapter 5) as well as to refine observations and conclusions of earlier geomorphic studies.
There are at least three justifications for the examination of the geomorphology of the area in which ecosystem studies are conducted. First, the present landscape and the materials that make it up provide the substrate on which ecosystem development occurs and may impose constraints, such as where soil resources are limited, on that development. Second, the nature of the landscape and the geomorphic processes acting on it often define a large part of the disturbance regime within which ecosystem processes occur (Swanson et al. 1988). Third, the processes of weathering, erosion, sediment transport, and deposition that define geomorphic dynamics within the landscape are themselves ecosystem processes, for example, involving the supply of resources to organisms. In this last context, it is noteworthy that drainage basins (also called watersheds or catchments) were recognized as units of scientific study during a similar time period in both geomorphology and ecology (Slaymaker and Chorley 1964); Bormann and Likens 1967); Chorley 1969). The drainage basin concept, the contention that lakes and streams act to integrate ecological and geomorphic processes, remains important in both sciences and underlies the studies in Green Lakes Valley reviewed here.
Over the past 30 years, Niwot Ridge and the adjacent catchment of Green Lakes Valley have been the subject of much research in geomorphology. Building on the studies of Outcalt and MacPhail (1965), White (1968), and Benedict (1970), work has emphasized the study of present-day processes and dynamics, especially of mass wasting in alpine areas. These topics have been reviewed by Caine (1974, 1986), Ives (1980), and Thorn and Loewenherz (1987). Studies of geomorphic processes have been conducted in parallel with work on Pleistocene (3 million to 10,000 yr BP) (p.46) and Holocene (10,000 yr BP to present) environments in the Colorado Front Range (Madole 1972; Benedict 1973) that have been reviewed by White (1982). This chapter is intended to update those reviews in terms that complement the presentation of ecological phenomena such as nitrogen saturation in the alpine (chapter 5) as well as to refine observations and conclusions of earlier geomorphic studies. In sequence it offers: (1) a brief description of the landforms of Green Lakes Valley and Niwot Ridge, (2) a summary of the historical development of that landscape, (3) a review of contemporary landforming processes in the valley, and (4) an evaluation of climatic influences on landscape development (Barsch and Caine 1984).
The Indian Peaks sector of the Colorado Front Range lies immediately south of Rocky Mountain National Park and straddles the Continental Divide at its most easterly point. This part of the range forms a prominent north-south barrier that contrasts the High Plains—30 km to the east and 2300 m lower in elevation. On the Eastern Slope of the range, environmental research has been concentrated on Niwot Ridge and Green Lakes Valley and in the Loch Vale Watershed (Baron 1992). The Green Lakes Valley is generally oriented west-east, with its western boundary defined by the Continental Divide between Navajo and Arikaree Peaks and its northern boundary by the crest of Niwot Ridge (Fig. 4.1). It is representative of the geomorphic processes of alpine catchments of the Southern Rockies in terms of its size, bedrock, relief, and landforms.
The Alpine Landscape
The Colorado Rockies provided the type areas from which Barsch and Caine (1984) defined a “Rocky Mountain Relief” type in a global classification of mountain geomorphology (p.47) (p.48) (p.49) typology. This is a landscape that has lower total relief, more-rounded ridges and summits, and a less-pervasive effect of Pleistocene glaciation than many other high mountain systems, such as those of the Alps, Scandinavia, and the Himalaya. It includes a wide variety of finer scaled structural features that have been characterized by O’Neill and Mark (1990) and demonstrates how the slope angles of these mountains are distinctly different from those found in unglaciated areas such as the Appalachians and Oregon.
Studies in Green Lakes Valley have been conducted in a set of 10 nested catchments with areas that vary from less than 10 hectares to the 710 hectares of the entire drainage above the Albion townsite (Fig. 4.1). The morphometric characteristics of these basins are summarized in Table 4.1 and Fig. 4.2. Catchment shape, estimated as a form ratio (i.e., from the equidimensional basin above Navajo to the trapezoidal form of the entire catchment above Albion) varies from 1.0 to 2.5. Within sections of the valley (e.g., between Navajo and Green Lake 4 or between the Lake Albion Inlet and Albion), there is a tendency for the basin shape to elongate with downvalley distance as valley width remains approximately constant (Fig. 4.2a). Relief in the valley increases in predictable fashion with catchment area, but at a lower rate than the increase in basin area (Figs. 4.2c, 4.3). The floor of Green Lakes Valley has the stepped form of a glaciated mountain valley (e.g., Embleton (p.50) and King 1975) and the spacing of the bedrock steps tends to increase with downvalley distance in the upper valley, to Green Lake 4, and again along the lower valley (Fig. 4.2b). The discontinuity in these trends occurs between Green Lakes 3 and 4 and corresponds to a large step that separates the two different sections of the valley (Caine and Thurman 1990).
Table 4.1. Basin Morphometry of the Green Lakes Valley
Mean Slope (°)
Stream Gradient (°)
Green Lake 5
Green Lake 4
Green Lake 3
Green Lake 1
The ridges defining Green Lakes Valley have two contrasting forms. Around the upper valley, especially west of Green Lake 4, they consist of bedrock or blockfieldcovered, narrow ridges (aretes) between glacial cirques (Fig. 4.4) (White 1976, 1982. Farther east, and most obviously on Niwot Ridge, they consist of broadly convex ridges (Benedict 1970). The surfaces of these ridges are extensively patterned by relict sorted polygons, with active solifluction lobes (wedges of stones created by ice), terraces, and patterned ground, where a high groundwater table is maintained in autumn and winter (Figs. 4.5, 4.6) (Benedict 1970, Fahey 1975).
Below these ridges, the alpine slopes of Green Lakes Valley fit the cliff-talussub-talus model that describes a rock-dominated system, the coarse debris system (Caine 1974; Thorn and Loewenherz 1987). These slopes are dominated by rockwalls with steep, though not vertical, profiles that are broken by structurally defined ledges. The alpine cliffs are frequently indented by couloirs (steep avalanche gullies), which channel snow, water, and rockfall debris and so influence the form of the subjacent talus (Rapp 1960; Davinroy 1994). The proportion of the landscape in (p.51) (p.52) the cliff, talus and sub-talus units (almost 100% above Green Lake 5) is reflected in basin relative relief and mean slope (Fig. 4.2d, e).
The talus slopes of the upper valley are predominantly straight and steep, with angles in excess of 32° (Fig. 4.7) and a planar form (Rapp 1960). Those of the lower valley, especially on its south wall below Kiowa Peak, are more varied. Here, talus development has been more influenced by snow avalanche and debris flow activity, and so the slopes have a more-concave profile form (White 1968). Talus cones (Fig, 4.8), reflecting the gullied topography of the overlying cliff, are also more common and better developed here than in the upper valley.
The lower talus slope and the valley floor beneath it carry a distinctive suite of features, particularly on the north-facing side of the valley. This is where the three valley-side rock glaciers of Green Lakes Valley occur (White 1976). Elsewhere, the toe of the talus slopes may be modified by debris flow deposits or small, lobate terraces, suggesting creep or flow in the blocky debris. A fringe of large blocks is also common where the valley floor is open beyond the foot of the talus.
(p.53) Most of the streams in Green Lakes Valley are small and carry flows only during the May-September period of snow melt (chapter 5). The channels are generally less than 2 m wide with a stable but rough bouldery floor and occasionally disappear beneath the surface materials in the upper valley. A great variability in flow depths, widths, and gradients (e.g., Fig. 4.2f) contribute to channel hydraulics that are highly variable in both space and time (Furbish 1993). In places such as the inlet to Green Lake 4, a greater organization is forced on the channel form by the influence of deep winter snow cover that forms sub-nival pavements along the drainage line (White 1972; Hara and Thorn 1982).
Channel length increases in simple predictable fashion with increased catchment area (Fig. 4.2g) but drainage density (channel length/area) remains relatively constant, or at least not varying consistently, throughout the nested basins of the system (Fig. 4.2h). Average channel gradients tend to decrease downstream, as basin area increases (Fig. 4.2f) and is clearly conditioned by the stepped form of the valley floor.
(p.54) The six lakes of Green Lakes Valley all occupy natural bedrock basins but have been modified during the last century. Those of the lower valley have been impounded and their levels raised by up to 3.5 m. In the upper valley, Green Lakes 4 and 5 now have levels about 1 m below their natural level because of channelization of their outlet streams.
The morphometry of the lake basins can be estimated from the surveys of McNeely (1983) and Harbor (personal communication). Despite dams, their volumes are relatively slight and amount to no more than 10 days of summer inflows to them (Table 4.2). The total area of lakes and their importance in terms of dominating biogeochemical processes within the landscape tend to increase downvalley (Fig. 4.2i, j). All of the lakes occupy bedrock basins that become less linear with downvalley distance: The length/width ratio changes from more than 3.5 at Green Lakes 4 and 5 to 1.2 at Lake Albion (Table 4.2). Volume development (Dv), as estimated by the ratio: 3D/Dmax (where D = mean depth and Dmax = is maximum depth) (Hutchinson 1957), reflects the origin of the lakes as glacial rock basins and the pattern of valley steps (Table 4.2). Dv is greatest at Green Lake 1, Green Lake 3, and Green Lake 5, all situations where glacial overdeepening of the valley is most marked.
Table 4.2. Characteristics of the Green Lakes
Volume (*106 m3)
Green Lake 5
Green Lake 4
Green Lake 3
Green Lake 2
Green Lake 1
Sources: Data derived from surveys of McNeely (1983) and Harbor (personal communication).
The Development of the Landscape
The geologic history of Green Lakes Valley is that of the Colorado Front Range, which originated in the regional uplift of the Laramide Orogeny (70-50 Ma) (e.g., Madole et al. 1987). Erosion during uplift subsequently removed between 3.0 and 3.5 km of Paleozoic and Mesozoic rocks from the crest of the range to expose the granites and gneisses of its core (Gregory and Chase 1992), 1994).
The east slope of the Colorado Front Range descends from the Continental Divide (at about 4000 m elevation) to the High Plains (below 2300 m) as a series of summits (p.55) and ridges. The nature and origin of the erosion surface(s) has been the subject of debate for more than 100 years (Bradley, 1987). It is now usual to interpret this landscape as derived from a single surface of late-Eocene age (Epis and Chapin 1975); Gregory and Chase 1992), which may have had considerable relief and may also have been warped and faulted during late-Tertiary uplift (Bradley 1987).
The history of the late-Eocene surface after its formation is also the subject of continuing debate. For a long time, there was broad agreement that it had been produced at low elevation, perhaps close to sea level (Davis 1911), and then uplifted 2000-3000 m during the Pliocene. This post-Laramide uplift initiated fluvial dissection during an episode of valley incision to define the broad outline of the present drainage pattern (Wahlstrom 1947; Scott 1975; Bradley 1987). Such uplift may have been partially driven by isostatic responses to valley erosion (e.g., Small and Anderson 1995); Wernicke et al. 1996). This would also be consistent with the suggestion that erosion surfaces like that of the Front Range may develop at relatively high elevation with later dissection and erosion initiated by climatic change rather than tectonic uplift (Molnar and England 1990); Gregory and Chase 1992), 1994).
Locally, remnants of the late-Eocene erosion surface on Niwot Ridge and in Green Lakes Valley decline eastward from 3700 m at the alpine climate station (Dl) and the saddle between Kiowa Peak and Mt. Albion to 3500 m at Niwot Mountain. It is also represented by the summit elevations of the Indian Peaks themselves (generally between 4000 and 4100 m). Together, these remnants suggest an eastward gradient of ca. 0.06, which is consistent with the elevation of ridges running east from the Continental Divide to the Great Plains.
The erosion surface remnants of the Indian Peaks are covered by coarse deposits that have been extensively reworked by mass wasting and frost processes during the Pleistocene. Initially, these deposits were interpreted as glacial tills from a Pliocene or early Pleistocene icesheet that developed on the range before the present valley system was established (Wahlstrom 1947). They have since been reexamined by Madole (1982), who found no evidence to support a glacial origin.
The glacial record on the East Slope of the Front Range has been comprehensively reviewed by Madole (1972, 1976, 1986) and Meierding and Birkeland (1980). Glacial deposits and features are generally confined to the valleys above 2500 m elevation and related to three glaciations that are differentiated by moraine morphology, soil development, and surface weathering (Madole 1976). All three glaciations were limited to existing valley systems that they postdate. With the exception of the Lake Devlin deposits (Madole 1976), absolute dating of glacial events is usually restricted to minimum estimates of deglaciation (Fig. 4.9).
Bonnett (1970, quoted in White 1982) reports till of pre-Bull Lake age in the drainage of North Boulder Creek at 2485 m elevation, east of the more obvious moraines between North Boulder Creek and Como Creek. Bull Lake glacial deposits occur more extensively in the North Boulder Creek drainage. These deposits derive from at least two major glacial intervals (stades) that have long been thought to have ended before 70 ka (70,000 years ago) (Richmond 1986; Madole 1976). Recent dating (p.56) (p.57) by cosmogenic beryllium-10 isotope (10Be) of the Bull Lake moraines near Fremont Lake, Wind River Range (138 ± 4 ka, 10Be years) supports this age estimate (Gosse et al. 1995a), leaving a hiatus of more than 50 ka before the subsequent Pinedale glaciation.
The evidence of Pinedale glaciation is a dominant element in the contemporary scenery of the Front Range. Here, as elsewhere in the Rockies, the earliest evidence of Pinedale glaciation dates from about 30 ka (Nelson et al. 1979; Madole 1986; Sturchio et al. 1994). Pinedale glaciation reached its maximum extent by 23 ka, when glacial Lake Devlin was impounded in the valley of a right-bank tributary to North Boulder Creek. The 10,000-yr sequence of proglacial lake sediments that followed this was terminated by catastrophic drainage during deglaciation at 13 ka (Madole 1986). Thereafter, deglaciation of the North Boulder Creek valley, including Green Lakes Valley, seems to have occurred quickly: Harbor (1984) reports a 12-ka basal date from the sediments in Green Lake 5, which matches the age of deglaciation elsewhere in the Rocky Mountains (Davis 1988; Fall et al. 1995; Gosse et al. 1995a).
Within Green Lakes Valley, Pinedale glacial ice accounts for the roches mou tonnees (glacially abraded and streamlined bedrock knob) and smoothed bedrock surfaces on the valley floor above Lake Albion. On a larger scale, the stepped long profile of the valley, including the bedrock hollows now occupied by lakes, record the imprint of glacial erosion, though these forms may be largely inherited from earlier glacials (Caine 1986). In the valley above Albion, at 3300 m, the approximate location of the equilibrium line during Pinedale glaciation, there is little evidence of glacial deposition from the main Pinedale glaciation. At the foot of the present Arikaree Glacier (Fig. 4.10), the 5-m-high moraine of “Triple Lakes” age probably represents the latest Pinedale glacial expansion, which may be correlated with the Younger Dry as (around 11, 000 years BP) event around the North Atlantic (Zielinski and Davis 1987); Davis 1988; Menounos and Reasoner 1997).
Holocene Landform Development
Early work on the Holocene history of Arikaree cirque, at the head of Green Lakes Valley, suggested three stades (intervals) of Neoglaciation (Ives 1953; Benedict 1968; Mahaney 1971; Carroll 1974). These were widely recognized in the Rocky Mountains but have only recently been well dated (e.g., Zielinski and Davis 1987); Gosse et al. 1995b). Davis (1988) provides a comprehensive review of Neoglacial chronologies in the Rocky Mountains, including the Colorado Front Range.
During the Altithermal, from about 9000 to 5000 BP, fossil evidence from Lake Isabelle suggests milder conditions with a higher tree line than at present (Elias 1985; chapter 15), which supports the suggestion that there were no permanent snowpatches in the Front Range during this interval (Benedict 1973). This would have allowed soil development in the hollows now occupied by snowpatches and explain the existence of paleosols at these sites (Burns 1979, 1980).
Cooler and wetter conditions, with persistent snowpatches and a lower tree line, followed the Altithermal (Elias 1985). This is associated with increased rates of geomorphic activity in the Arapaho Pass area between 5 and 3 ka (Benedict 1985) and (p.58) the development of rock glaciers (lobate body of angular debris with a steep front at the angle of repose) in the Front Range (Madole 1976; Benedict 1981; Dowdeswell 1982; Harbor 1985).
The Audubon stade of Neoglaciation (2.4-0.95 ka) involved further rock glacier development and extensive snow cover above 3700-m elevation (Benedict 1973, 1993. In Green Lakes Valley, the lichen-kill zones around and north of Arikaree Glacier are evidence of a more-extensive snow cover at this time (Fig. 4.10; Mahaney 1971; Carroll 1974).
The Little Ice Age is represented in the Indian Peaks by deposits of the Arapaho Peak stade (300-100 BP) (Benedict 1973), including small moraines on the proximal face of earlier moraines at Arikaree Glacier. Glacier recession in the twentieth century coincides with documentary records (Lee 1900; Fenneman 1902; Henderson 1904; Ives 1940, 1953), which show a marked loss of ice mass from Arapaho Glacier between 1900 and the midcentury (Waldrop 1964) and a great reduction in the intensity of glacial processes (Reheis 1975). In Green Lakes Valley, Arikaree Glacier experienced a similar loss of mass in the first half of this century (Ives 1953) but has since maintained a quasi-balanced mass budget (Johnson 1980).
Holocene hillslope development in the alpine of the Front Range has undergone marked changes in intensity during the past 10,000 years (Benedict 1985). Neoglacial intervals of extensive snow cover and glacier expansion correlate with periods of greater activity in the cliff-talus and rock glacier systems (Madole 1976; White 1976; Dowdeswell 1982; Harbor 1985). On other alpine slopes, the same intervals (p.59) of wetter, cooler conditions coincide with intervals of more rapid movement in stonebanked terraces and solifluction lobes (Benedict 1967, 1970, 1973, 1976, 1993).
In general, Holocene lake sediments reflect geomorphic activity in the terrestrial environment only poorly (Caine 1974; Andrews et al. 1985; Harbor 1984, 1985). Variations in the organic content of the sediments in the small lake below Arapaho rock glacier show intervals of increased geomorphic activity during the late-Holocene (Davis 1988). However, most of the Holocene sediments in the larger alpine lakes are homogeneous with few textural changes that would reflect changes in hillslope activity (Caine 1986).
Alpine Geomorphic Processes
Contemporary geomorphic activity in Green Lakes Valley has been described in terms of three overlapping systems that involve different materials and dynamics (Caine 1974, 1986; Thorn and Loewenherz 1987). These are the coarse debris system, the fine sediment system, and the geochemical system, which are treated here separately.
The Coarse Debris System
As in many mountain areas, the movement and storage of coarse debris (>8-mm size) dominates the sediment budget of Green Lakes Valley (Table 4.3). This system (p.60) involves the bedrock cliffs, their associated talus, and the accumulated blocky debris in the talus slopes and rock glaciers below them (Fig. 4.11). This system has been effectively closed during the Holocene but was presumably breached during glacial periods when the accumulated debris in it would be incorporated into the tills deposited at lower elevations (Figs. 4.7, 4.8).
Table 4.3. Sediment Budgets for the Upper Part of the Green Lakes Valley Above Green Lake 4
Volume (m3 yr−1)
Work (106 J yr−1)
Coarse debris system
Rock glacier flow
Fine sediment system
Soil creep and solifluction
Fluvial sediment export
Source: Caine (1986) with selected updates and additions.
A 25-year record in upper Green Lakes suggests that rockfall from the alpine cliffs yields about 12 m3 yr−1 of debris, corresponding to a rate of cliff retreat of only 0.02 mm yr−1 (Caine 1986). By comparison to the existing volume of talus in the valley, this mass is negligible and has little effect on the landscape. Much of the rockfall material is finer than the existing debris mantle into which it is incorporated, without modifying the slope form. Even large rockfalls (involving up to 50 m3 each), which occur with a decadal frequency, have little effect on the talus for the large blocks in them usually move over the talus to its toe.
(p.61) Movement of the Green Lakes talus was noted by Ives (1940) and has since been measured by White (1981), Caine (1986), and Davinroy (1994). On average, it amounts to less than 3 cm yr−1 at the surface, tends to decrease downslope from the head of the talus, and is highly variable (e.g., Gardner 1979). Where flows of water, snow, and debris are concentrated through couloirs, rates of deposition and movement may be two orders of magnitude greater than those on open talus slopes (Gardner 1982, 1986; Davinroy 1994). White (1968) reported the influence of debris flows (Figs. 4.12, 4.13) and wet snow avalanches in forming concave talus profiles. The influence on sediment transport of avalanches on Arikaree Peak has been reported in Caine (1986, 1995a). Dry snow avalanches from Kiowa Peak occasionally transport coarse debris onto the ice cover of Green Lake 3, contributing to the sediments of that lake.
In the talus foot zone, the valley-side rock glacier at Green Lake 5 is the only obviously active one in Green Lakes Valley (White 1981). It has a surface velocity of less than 2.0 cm yr−1, an order of magnitude less than the flow rates of valley-floor rock glaciers in the Front Range (White 1981; Benedict et al. 1986). The other two valley-side rock glaciers, at Green Lake 2 and in the Arikaree cirque, appear to have been inactive for millenia. Lichens and rock weathering on the former suggest that it stabilized more than 2000 years ago. The latter appears even older and was ascribed a late-Pinedale age by Carroll (1974).
The fine sediment cascade involves the slow processes of soil creep, frost creep, gelifluction (downward movement of soil due to freeze-thaw), and surficial erosion as well as the more catastrophic ones of landsliding and debris flow. Unlike the coarse debris of talus and blockfields, the finer material of this system may be entrained by water and air flows and incorporated into more-extensive exchanges, potentially extending beyond the alpine environment (Thorn and Darmody 1980), 1985; Burns and Tonkin 1982); Caine 1986; Litaor 1987). However, in the upper Green Lakes Valley, the fine sediment system involves only 15-20% of the mass in the coarse debris system (Table 4.3).
On Niwot Ridge, rates of solifluction and frost creep in the debris mantle are maximized in solifluction lobes and terraces that remain wet in autumn and early winter (Benedict 1970; Table 4.4). In better drained sites, rates of solifluction and creep are an order of magnitude lower (less than 0.1 cm yr−1), and these rates are probably more typical of the alpine tundra (Caine 1986).
Table 4.4. Landforms in Mountain Areas
Active glaciers, perennial snow, ice
Bedrock, blockfield, talus
Sorted patterned ground, “unbound” solifluction
Vegetated solifluction forms, micro-terracettes
(p.63) Estimates of soil erodibility suggest that surface erosion by rainsplash, surface runoff, frost action, and wind may be high where the tundra vegetation is disturbed (Summer 1982). Empirical studies of surface erosion support this and show that rates of soil loss increase by orders of magnitude where the vegetative cover is reduced to less than 25% in snowpatch sites (Caine 1976; Bovis 1978; Bovis and Thorn 1981). These sites become sources of sediments when they are exposed to summer rainstorms, although they may be protected from erosion by a long-lasting snow cover (Caine 1992; Caine and Swanson 1989). Changes in the winter snow cover during the late-Holocene should have modulated these influences over the alpine tundra (Benedict 1993).
Further sites of surficial erosion by wind and rain are the soil mounds and tunnel casts of pocket gophers (Thorn 1982). In fact, Thorn (1978) suggests that this is the most effective mechanism of soil erosion in the tundra of the Colorado Rockies; turning over up to 1500 kg ha 1 of soil each year (Stoecker 1976; chapters 7, 8).
Eolian deposition of dust into Green Lakes Valley may exceed 1 g m−2 yr−1 of silt and clay and appears to be an important source of buffering in the geochemical system (Litaor 1987). The source of the eolian deposits remains unstudied but is likely from the southwestern United States (chapter 3). On tundra surfaces, these fine particles are probably retained to become involved in pedogenesis (Thorn and Darmody 1980), 1985; Burns and Tonkin 1982); on bedrock and boulder surfaces, they may be transported through the fluvial system into lake sediments (Caine 1986.
Little of this activity involving fine sediment is reflected in suspended sediment transport through the stream systems of Green Lakes Valley (Table 4.3). Suspended sediment concentrations in the streams are generally low and yields only exceed 1.0 g m−2 yr−1 in years of high stream discharge (Caine 1995a). Snowpatch sites appear to be important sources of this sediment, but even there, sediment concentration-discharge relationships are not simple but reflect complex influences of the snow cover (Caine 1989, 1992).
The stream channels of Green Lakes Valley are effectively covered by coarse sediments and remarkably stable. The only visible changes effected by channelized flows occur where high meltwater flows released by the failure of snow and ice dams are superimposed from the snowcover onto the valley floor, usually outside the stream channels (Caine 1995a). The lake sediments are invariably fine grained, except (p.64) where high-energy slopes impinge directly on the lake shore (as at Green Lake 3), and are predominantly of silt and clay (Caine 1986). Over the last 5000 years, sedimentation into Green Lakes 4 and 5 has averaged only 0.15 mm yr−1 (Harbor 1984), which is consistent with the low suspended sediment concentrations of the streamflows.
The Geochemical System
Geochemical weathering rates in Green Lakes Valley were first estimated for the Martinelli catchment by Thorn (1975, 1976) and have been expanded during LTER studies of the entire valley. Solute concentrations and water discharges have been monitored in nine basins of varying size for 15 years. Solute concentrations in the stream waters are generally low but with regular temporal and spatial patterns (Caine and Thurman 1990).
Caine and Thurman (1990) found no consistent temporal trends in major ion concentrations and yields but more recent analyses of a longer record suggest a tendency toward acidification (Williams et al. 1996). This is most marked in the higher parts of Green Lakes Valley (Caine 1995b), corroborating the forecasts of Lewis (1982) and Kling and Grant (1984). The pattern of acidification has not yet been detected in increased base cation concentrations, which would reflect changing rates of geochemical denudation.
On an annual time scale, patterns of solute removal are dominated by the seasonal cycle, which explains about 65% of the variability in long-term records of specific conductance at Albion and Green Lake 4 (Caine and Thurman 1990). This pattern reflects the influence of snowmelt (Williams et al. 1993; Campbell et al. 1995). In May and June early meltwater flows produce relatively high concentrations in solutes at all sites in Green Lakes Valley (Caine and Thurman 1990); chapter 5). Concentrations then decline to low levels during the summer but rise gradually from late summer into winter as the relative proportions of ground and soil water contributions to streamflow increase. In small headwater basins, an equivalent out-of-phase periodicity in solute concentrations and water flows occurs on a daily basis during the summer (Caine 1989).
Solute concentrations increase consistently downstream (Table 4.5), reflecting a greater mean residence time of water; greater contributions to streamflow from groundwater, and increased biological activity. These influences are also evident in a comparison of the Green Lake 1 and Martinelli catchment records with those from basins of similar area at the head of the main valley (Caine and Thurman 1990).
Total solute yields from Green Lakes Valley average 10 g m−2 yr−1 (Table 4.5). Corrected for precipitation loading, this represents a bedrock lowering rate of less than 0.003 mm yr−1, appreciably lower than rates in mountain catchments elsewhere (e.g., Dethier 1986; Souchez and Lemmens 1987); Pape et al. 1995) but still greater than the rate of sediment yield (Table 4.5).
Spatial patterns of solute yields and concentrations reflect patterns of snow accumulation and meltwater drainage, modulated by soil- or groundwater flows (Dixon 1986; Caine 1992; Litaor 1993). These patterns are also reflected in variations (p.65) in the thickness of the chemically altered surface layers of rocks (Thorn 1975), although these may be affected by differential removal rates (Ballantyne et al. 1989).
Table 4.5. Geochemical Denudation in the Green Lakes Valley
Green Lake 5
Green Lake 4
Green Lake 1
Sediment budgets for Green Lakes Valley above Green Lake 4 were summarized by Caine (1986) and have been updated with more recent data (Table 4.3). The coarse debris system accounts for over 95% of all geomorphic work accomplished at the present time in the valley. Within that system, rockfall dominates because of the relatively great (vertical) distance of transport that it entails, especially in the larger falls, which often traverse the entire talus slope. The geomorphic work done in talus shift is less than this by a factor of three or four, despite the large mass of material involved. By comparison, other processes of coarse debris movement are relatively slight contributors to work done in the basin, although they may be important on individual slopes.
As in other coarse debris systems (e.g„ Rapp 1960), contemporary rates of talus accumulation and shift in Green Lakes Valley are incapable of accounting for the volume of coarse debris accumulated on the valley walls since deglaciation (Caine 1986). About 0.8 × 106 m3 of talus and a further 1.0 × 106 m3 of other debris have accumulated in the upper valley over 12 ka, at an average rate of almost 150 m3 yr−1, an order of magnitude greater than the rate shown by contemporary measurement (Table 4.3). This disparity argues for coarse debris production at rates up to two orders of magnitude greater than the present during the late Pleistocene and in Neoglacial stades (Harbor 1985). As Wahrhaftig (1987) points out with respect to rock glacier systems, this coarse material is likely to remain stored within the alpine until redistributed or removed by another major glaciation.
The fine sediment system is dominated by the large mass of material involved in soil creep and solifluction, which accounts for almost half of the work done in the fine sediment system. The fluvial transport of silt and clay to lake basins involves a (p.66) smaller mass but a greater vertical transport, which allows it to contribute more than 50% to the work done in the system. The major uncertainty in this budget concerns eolian transport, for which the volume of sediment involved and distances of transport are still only poorly estimated.
In contrast to the systems involving sediment and debris, the geochemical system is one in which the coupling of slope and stream systems is the closest. For that reason, it is estimated with greater precision as an average for entire drainage basins. In terms of total mass and work estimates, it has relatively low magnitude, accounting for only 16% of the total geomorphic work done in the basin and a much smaller proportion of the mass involved in geomorphic exchanges (Table 4.3). Nevertheless, it dominates the estimate of basin denudation, accounting for more than 92% of the work done in sediment and solute export and lake sedimentation.
In mountain environments, the concept of a climatic influence on landscape development has usually been defined in terms of altitudinal limits and elevational zones of landforms and processes (e.g., Budel 1954; Bik 1967; Rudberg 1972).
Above treeline, the periglacial landforms of Niwot Ridge and Green Lakes Valley may be classified into the blockfield and solifluction zones of Rapp and Rudberg (1960) and Rudberg (1972) (Table 4.4). Within these zones, landform patterns like those of other mid-latitude mountains (Chardon 1984; Kotarba 1984) are found. However, they are not defined by elevation alone, requiring an orthogonal criterion based on topographic location (e.g., Caine 1978). This separates the landforms of ridges and valleys and so reflects a historical influence, distinguishing landforms in areas of late-Pleistocene glaciation from those on terrain that has remained extraglacial at that time (Table 4.6).
In the solifluction zone, mapping of small-scale forms on Niwot Ridge by Benedict (1970) and surveys of forms in Green Lakes Valley allow the limits of active periglacial forms to be defined (Table 4.6). The variety of periglacial forms found in the zone suggest elevation limits that vary between 3350 m and 3520 m with some difference between the valley floor environment and that of the interfluve. Active sorted patterned ground suggests a 70-m difference in elevational limits between Niwot Ridge and the neighboring valley (Warburton 1991).
The features of the blockfield zone are separated spatially from those of the solifluction zone, with the single exception of talus that is active (if not intensely so) along the south valley wall to the elevation of Lake Albion (Table 4.6). However, this is in response to bedrock instability and high slope angles rather than to a direct climatic influence. In contrast, blockfield and rock glacier forms are found only above 3650 m, and the latter are restricted to the south wall of the valley, as elsewhere in the Front Range (White 1971).
Above the blockfield zone in the Indian Peaks is a glacial zone, at 3800 m and above, represented in Green Lakes Valley by Arikaree Glacier (Table 4.6). Glacial action here is presently minimal.
Table 4.6. Elevational Limits of Active Periglacial Forms in the Green Lakes Valley and on Niwot Ridge
Tongue-shaped rock glacier
Allochthonous (?) blockfields
Lobate rock glacier
Sorted Patterned ground
Sorted patterned ground
Turf-banked lobes and terraces
(p.67) (p.68) Attempts to define morphoclimatic zones on the basis of contemporary erosion on small plots and denudation rates in stream basins have not usually been successful. For example, Caine (1984) found only slight differences between the “high alpine” and the “alpine tundra” zones. These minor differences are particularly related to (1) the presence of weak glacial effects in the higher zone, (2) the greater activity in the talus slopes of the coarse debris system, and (3) slight increases in geochemical concentrations at lower elevations (Caine 1984).
On the broader scale of the entire Front Range, general contrasts of geomorphic levels are more evident. Although it includes areas of high activity, the subalpine forest is the least active in geomorphic terms today, with the exception, perhaps, of geochemical processing (Bovis 1978; Caine 1984). However, even this relative stability may be catastrophically interrupted by fire (Morris 1986; Morris and Moses 1987), wind storms, or anthropogenic disturbance of the ground cover (Bovis 1978). The stream channels of the subalpine belt often suggest higher rates of sediment transport and channel modification than those of the alpine, but this, too, could be a reflection of more than a century of channel and catchment disturbance.
The geomorphic systems of the Green Lakes Valley and Niwot Ridge are representative of the alpine areas of the southern Rocky Mountains, including the variety of landforms found in that broader geographic context. The effects of Pleistocene environments remain dominant in the present alpine landscape: in the glacial landforms of the valleys and on remnants of the Eocene erosion surface where periglacial modification of the surface materials occurred at the time of valley glaciation. Indirectly, glacial erosion has greatly affected the valley walls by steepening them and so allowing the Holocene development of talus and cliff systems.
In contrast to the lasting imprint of late Pleistocene and Neoglacial conditions, present-day processes of landscape change appear to be relatively modest. Presentday rates of geomorphic activity are slight when compared with those of other high mountain areas, as they are when compared with those of the past. This is especially true if a denudation rate, that is, the export of sediment and solutes, is used to estimate geomorphic activity for the entire system or for parts of it. Mass loss through the stream channels represents no more than 15% of the geomorphic work done in the basin. The remaining work is done in a closed system from which the accumulating coarse debris will not be discharged to lower elevations until the advent of renewed glacial conditions. This relative calm in geomorphic activities allows for the expression of biotic activities to assume a relatively large role in the biogeochemistry of the alpine.
All of this contributes to the high variability in alpine landforms, which is an important influence on other environmental processes and on biotic patterns. The resulting variety also accounts for much of the visual attractiveness of this mountain landscape.
Andrews, J. T., Birkeland, P. W. Harbor, J. Dellamonte, N. Litaor, M. and Kihl. R. 1985. Holocene sediment record, Blue Lake, Colorado Front Range. Zeitschrift fur Gletscherkunde und Glazialgeologie 21: 25–34.
Ballantyne, C. K., Black, N. M. and Finlay. D. P. 1989. Enhanced boulder weathering underlate-lying snowpatches. Earth Surface Processes and Landforms 14: 745–750.
Baron, J. S. 1992. Biogeochemistry of a subalpine ecosystem: Loch Vale watershed. Ecological Studies 90. New York: Springer-Verlag.
Barsch, D., and Caine. N. 1984. The nature of mountain geomorphology. Mountain Research and Development 4: 287–298.
Benedict, J. B. 1967. Recent glacial history of an alpine area in the Colorado Front Range, U. S. A. I. Establishing a lichen-growth curve. Journal of Glaciology 6: 817–832.
Benedict, J. B. 1968. Recent glacial history of an alpine area in the Colorado Front Range, U. S. A. II. Dating the glacial deposits. Journal of Glaciology 7: 77–87.
Benedict, J. B. 1970. Downslope soil movement in a Colorado alpine region: rates, processes and climatic significance. Arctic and Alpine Research 2: 165–226.
Benedict, J. B. 1973. Chronology of cirque glaciation, Colorado Front Range, U. S. A. Quaternary Research 3: 584–599.
Benedict, J. B. 1976. Frost creep and gelifluction features: A review. Quaternary Research 6: 55–76.
Benedict, J. B. 1981. The Fourth of July Valley: Glacial geology and archeology of the timberline ecotone. Ward, CO: Center for Mountain Archeology Research Report 2: 139 pp.
Benedict, J. B. 1985. Arapaho Pass: Glacial geology and archeology at the crest of the Colorado Front Range. Ward, CO: Center for Mountain Archeology Research Report 3: 197 pp.
Benedict, J. B. 1993. A 2000-year lichen-snowkill chronology for the Colorado Front Range, U. S. A. The Holocene 3: 27–33.
Benedict, J. B., Benedict, R. J. and Sanville. D. 1986. Arapaho Rock Glacier, Front Range, Colorado, USA: a 25-year resurvey. Arctic and Alpine Research, 18: 349–352.
Bik, M. J. J. 1967. Structural gomorphology and morphoclimatic zonation in the Central Highlands, Australian New Guinea. In Landform studies from Australia and New Guinea, edited by Jennings J. N. and Mabbutt. J. A. Canberra, Australia: A. N. U. Press.
Bonnett, R. B. 1970. The glacial sequence of upper Boulder Creek drainage basin in the Colorado Front Range. Ph.D. diss., Ohio State University.
Bormann, F. H., and Likens. G. E. 1967. Nutrient cycling. Science 155: 424–429.
Bovis, M. J. 1978. Soil loss in the Colorado Front Range: Sampling design and areal variation. Zeitschrift fur Geomorphologie Supplementband 29: 10–21.
Bovis, M. J., and Thorn. C. E. 1981. Soil loss variation within a Colorado alpine area. Earth Surface Processes and Landforms 6: 151–163.
Bradley, W. C. 1987. Erosion surfaces of the Colorado Front Range: A review. In Geomorphicsystems of North America, edited by Graf. W. L. Boulder, CO: Geological Society of America Centennial Special Vol. 2.
Budel, J. 1954. Klima-morphologische Arbeiten in Athiopien im Fruhjahr 1953. Erdkunde 8: 139–156.
Bums, S. F. 1979. Buried soils beneath alpine snowbanks may date the end of the Altithermal. Geological Society of America, Abstracts with Programs 11: 267–268.
Bums, S. F. 1980. Alpine soil distribution and development, Indian Peaks, Colorado Front Range. Ph.D. diss., University of Colorado, Boulder.
Bums, S. F., and Tonkin. P. J. 1982. Soil-geomorphic models and the spatial distribution and development of alpine soils. In Space and time in geomorphology, edited by Thorn. C. E. London: Allen and Unwin.
Caine, N. 1976. The influence of snow and increased snowfall on contemporary geomorphic processes in alpine areas. In Ecological impacts of snowpack augmentation in the San Juan Mountains, Colorado, edited by Steinhoff H. W. and Ives. J. D. CSU-DWS 7052–4, Fort Collins, CO: Colorado State University.
Caine, N. 1978. Climatic geomorphology in mid-latitude mountains. In Landform evolution in Australasia, edited by Davies J. L. and Williams. M. A. J. Canberra, Australia: A. N. U Press.
Caine, N. 1984. Elevational contrasts in contemporary geomorphic activity in the Colorado Front Range. Studia Geomorphologica Carpatho-Balcanica 18: 5–31.
Caine, N. 1986. Sediment movement and storage on alpine hillslopes in the Colorado Rocky Mountains. In Hillslope processes, edited by Abrahans. A. D. Winchester, MA: Allen and Unwin.
Caine, N. 1989. Diurnal variations in the inorganic solute content of water draining from an alpine snowpatch. Catena 16: 153–162.
Caine, N. 1992. Sediment transfer on the floor of the Martinelli Snowpatch, Colorado Front Range, U. S. A. Geografiska Annaler 74A: 133–144.
Caine, N. 1995a. Snowpack influences on geomorphic processes in Green Lakes Valley, Colorado Front Range. Geographical Journal 161: 55–68.
Caine, N. 1995b. Temporal trends in the quality of streamwater in an alpine environment: Green Lakes Valley, Colorado Front Range, U. S. A. Geografiska Annaler A: 207–220.
Caine, N., and Swanson. F. J. 1989. Geomorphic coupling of hillslope and channel systemsin two small mountain basins. Zeitschrift für Geomorphologie 33: 189–203.
Caine, N., and Thurman. E. M. 1990. Temporal and spatial variations in the solute content ofan alpine stream, Colorado Front Range. Geomorphology 4: 55–72.
Campbell, D. H., Clow, D. W. Ingersoll, G. P. Mast, M. A. Spahr, N. E. and Turk. J. T. 1995. Processes controlling the chemistry of two snowmelt-dominated streams in the Rocky Mountains. Water Resources Research 3: 2811–2821.
Carroll, T. 1974. Relative age-dating techniques and a late Quaternary chronology, Arikaree Cirque, Colorado. Geology 2: 321–325.
Chardon, M. 1984: L’etagement des paysages et les processus geomorphologiques actuelsdans les Alpes occidentales. Studia Geomorphologica Carpatho-Balcanica 18: 33–43.
Chorley, R. J. 1969. The drainage basin as the fundamental geomorphic unti. In Water, earth and man, edited by Chorley. R. J. London: Methuen.
Davinroy, T. C. 1994. Rates and controls of rock movement through alpine couloirs, Colorado Front Range. Master’s thesis, University of Colorado, Boulder.
Davis, W. M. 1911. The Colorado Front Range. Annals Association of American Geographers 1:21–83.
Davis, P. T. 1988. Holocene glacier fluctuations in the America Cordillera. Quaternary Science Reviews 7: 129–157.
Dethier, D. P. 1986. Weathering rates and the geochemical flux from catchments in the Pacific Northwest, U. S. A. In Rates of chemical weathering of rocks and minerals, edited by Colman S. M. and Dethier. D. P. Orlando, FL: Academic Press.
Dixon, J. C. 1986. Solute movement on hillslopes in the alpine environment of the Colorado Front Range. In Hillslope processes, edited by Abrahams. A. D. Winchester, MA: Allen and Unwin.
Dowdeswell, J. A. 1982. Relative dating of late Quaternary deposits using cluster and dis-criminant analysis, Audubon Cirque, Mount Audubon, Colorado Front Range. Boreas 11: 151–161.
Embleton, C. and M. King. C. A. 1975. Glacial and periglacial geomorphology. Vol. 1, Glacial geomorphology. London: Edward Arnold.
Epis, R. C., and Chapin. C. E. 1975. Geomorphic and tectonic implications of the post-Laramide, late Eocene erosion surface in the southern Rocky Mountains. Geol. Society of America Memoir 144: 45–74.
Fahey, B. D. 1975. Non-sorted circle development in a Colorado alpine location. Geografiska Annaler 57A: 153–164.
Fall, P. L., Davis, P. T. and Zielinski. G. A. 1995. Late Quaternary vegetation and climate ofthe Wind River range, Wyoming. Quaternary Research 43: 393–404.
Fenneman, N. N. 1902. The Arapahoe Glacier in 1902. Journal of Geology 10: 839–851.
Furbish, D. J. 1993. Flow structure in a bouldery mountain stream with complex bed topog-raphy. Water Resources Research 29: 2249–2263.
Gardner, J. S. 1979. The movement of material on debris slopes in the Canadian Rockies. Zeitschrift für Geomorphologie Supplementband 23: 45–57.
Gardner, J. S. 1982. Alpine mass wasting in contemporary time: Some examples from the Canadian Rocky Mountains. In Space and time in geomorphology, edited by Thorn. C. E. London: Allen and Unwin.
Gardner, J. S. 1986. Sediment movement in ephemeral streams on mountain slopes, Canadian Rocky Mountains. In Hillslope processes, edited by Abrahams. A. D. Winchester, MA: Allen and Unwin.
Gosse, J. C., Klein, J. Evenson, E. B. Lawson, B. and Middleton. R. 1995a. Beryllium-10 dating of the duration and retreat of the last Pinedale glacial sequence. Science 268: 1329–1333.
Gosse, J. C., Klein, J. Evenson, E. B. Lawson, B. and Middleton. R. 1995b. Precise cosmogenic 10Be measurements in western North America: Support for a global Younger Dryas cooling event. Geology 23: 877–880.
Gregory, K. M., and Chase. C. G. 1992. Tectonic significance of paleobotanically estimatedclimate and altitude of the late Eocene erosion surface, Colorado. Geology 20: 581–585.
Gregory, K. M., and Chase. C. G. 1994. Tectonic and climatic significance of a late Eocene low-relief, high level geomorphic surface, Colorado. Journal of Geophysical Research 99B: 20141–20160.
Hara, Y. and Thorn. C. E. 1982. Preliminary quantitative study of alpine subnival boulder pavements, Colorado Front Range, U. S. A. Arctic and Alpine Research 14: 361–367.
Harbor, J. 1984. Terrestrial and lacustrine evidence for Holocene climatic/geomorphicchange in the Blue Lake and Green Lakes Valleys of the Colorado Front Range. Master’sthesis, University of Colorado, Boulder.
Harbor, J. 1985. Problems with the interpretation and comparison of Holocene terrestrial and lacustrine deposits: An example from the Colorado Front Range, U. S. A. Zeitschrift fur Gletscherkunde und Glazialgeologie 21: 17–24.
Hastenrath, S. 1974. Glaziale und periglaziale Formbildung in Hoch-Semyen, Nord-Athio-pien. Erdkunde 28: 176–185.
Henderson, J. 1904. Arapaho Glacier in 1903. Journal of Geology 12: 30–33.
Hollerman, P. W. 1967. Zur Verbreitung rezenter periglazialer Kleinformen in den Pyrenaenund Ostalpen. Gottinger Geographische Abhandlungen 40: 198 pp.
Hutchinson, J. E. 1957. A treatise on limnology. Vol. 1: Geography, physics and chemistry. New York: Wiley.
Ives, R. L. 1940. Rock glaciers in the Colorado Front Range. Geological Society of America Bulletin 51: 1271–1294.
Ives, J. D. 1980. Geomorphology overview. In Geoecology of the Colorado Front Range: Astudy of alpine and subalpine environments, edited by Ives. J. D. Boulder, CO: Westview Press.
Johnson, J. B. 1980. Mass balance studies on the Arikaree Glacier. In Geoecology of the Colorado Front Range: A study of alpine and subalpine environments, edited by Ives. J. D. Boulder, CO: Westview Press.
Kling, G. W., and Grant. M. C. 1984. Acid precipitation in the Colorado Front Range: Anoverview with time predictions for significant effects. Arctic and Alpine Research 16: 321–329.
Kotarba, A. 1984. Elevational differentiation of slope geomorphic processes in the Polishtatra Mountains. Studia Geomorphologica Carpatho-Balcanica 18: 117–133.
Lee, W. T. 1900. The glacier of Mt Arapahoe, Colorado. Journal of Geology 8: 647–654.
Lewis, W. M., Jr. 1982. Changes in p H and buffering capacity of lakes in the Colorado Rockies. Limnology and Oceanography 27: 167–172.
Litaor, M. I. 1987. The influence of eolian dust on the genesis of alpine soils in the Front Range, Colorado. Soil Science Society of America Journal 51: 142–147.
Litaor, M. I. 1993. The influence of soil interstitial waters on the physicochemistry of major, minor and trace metalsin stream waters of the Green Lakes Valley, Front range, Colorado. Earth Surface Processes and Land forms 18: 489–504.
Madole, R. F. 1972. Neoglacial facies in the Colorado Front Range. Arctic and Alpine Research 4: 119–130.
Madole, R. F. 1976. Glacial geology of the Colorado Front Range. In Quaternary stratigraphy of North America, edited by Mahaney. W. C. Stroudsburg, PA: Dowden, Hutchinson and Ross.
Madole, R. F. 1982. Possible origins of till-like deposits near the summits of the Front Range in north-central Colorado. U. S. Geological Survey Professional Paper 1243: 31 pp.
Madole, R. F. 1986. Lake Devlin and the Pinedale glacial history of the Colorado Front Range. Quaternary Research 25: 43–54.
Madole, R. F., Bradley, W. C. Loewenherz, D. S. Ritter, D. F. Rutter, N. W. and Thorn. C. E. 1987. In Geomorphic systems of North America, edited by Graf. W. L. Boulder, CO: Geological Society of America, Centennial Special Vol. 2.
Mahaney, W. C. 1971. Note on the “Arikaree Stade” of the Rocky Mountains neoglacial. Journal of Glaciology 10: 143–144.
McNeely, R. 1983. Preliminary physical data on Green Lakes 1-5, Front Range, Colorado. University of Colorado Long-Term Ecological Research Working Paper 83/4.
Meierding, T. C., and Birkeland. P. W. 1980. Quaternary glaciation of Colorado. In Coloradogeology, edited by Kent H. C. and Porter. K. W. Denver, CO: Rocky Mountain Associ-ation of Geologists.
Menounos, B., and Reasoner. M. 1997. Evidence for cirque glaciation in the Colorado Front Range during the Younger Dryas chronozone. Quaternary Research 48: 38–47.
Molnar, P., and England. P. 1990. Late Cenozoic uplift of mountain ranges and global climatechange: Chicken or egg. Nature 346: 29–34.
Morris, S. E. 1986. The significance of rainsplash in the surficial debris cascade of the Colorado Front Range foothills. Earth Surfaces Processes and Landforms 11: 11–22.
Morris, S. E., and Moses. T. A. 1987. Forest fire and the natural soil erosion regime in the Colorado Front Range. Annals of the Association American Geographers 77: 245–254.
O’Neill, P. M. and Mark. D. M. 1990. On the frequency distribution of l and slope. Earth Surface Processes and Landforms 12: 127–136.
Outcalt, S. I., and Mac Phail. D. D. 1965. A survey of neoglaciation in the Front Range of Colorado. University of Colorado Studies, Series in Earth Science 4: 124 pp.
Pape, G. A., Dorn, R. I. and Dixon. J. C. 1995. A new conceptual model for understandinggeographical variations in weathering. Annals of the Association of American Geogra-phers 85: 38–64.
Rapp, A. 1960. Talus slopes and mountain walls at Tempelfjorden, Spitsbergen. Norsk Polarinstitut Skrifter 119: 96 pp.
Rapp, A., and Rudberg. S. 1960. Recent periglacial phenomena in Sweden. Biuletyn Periglacjalny 8: 143–154.
Reheis, M. J. 1975. Source, transportation and deposition of debris on Arapaho Glacier, Front Range, Colorado. Journal of Glaciology 14: 407–420.
Richmond, G. M. 1986. Stratigraphy and correlation of glacial deposits of the Rocky Mountains, the Colorado Plateau and the ranges of the Great Basin. Quaternary Science Reviews 5: 99–127.
Rudberg, S. 1972. Periglacial zonation—a discussion. Gottinger Geographische. Abhand-lung 60: 221–233.
Scott, G. R. 1975. Cenozoic surfaces and deposits in the Southern Rocky Mountains. Geological Society of America Memoir 144: 227–248.
Slaymaker, O., and Chorley. R. J. 1964. The Vigil network system. Journal of Hydrology 2: 19–24.
Small, E. E., and Anderson. R. S. 1995. Geomorphically driven late Cenozoic uplift in the Sierra Nevada, California. Science 270: 277–280.
Souchez, R. A., and Lemmens. M. M. 1987. Solutes. In Glacio-fluvial sediment transfer, edited by Gumall A. M. and Clark. M. J. Chichester, England: Wiley.
Stoecker, R. 1976. Pocket gopher distribution in relation to snow in the alpine tundra. In Ecological impacts of snowpack augmentation in the San Juan Mountains, Colorado, edited by Steinhoff H. W. and Ives. J. D. Fort Collins: Colorado State University, CSU-DWS 7052-4.
Sturchio, N. C., Pierce, K. L. Murrell, M. T. and Sorey. M. L. 1994. Uranium-series ages oftravertines and timing of the last glaciation in the northern Yellowstone area, Wyoming-Montana. Quaternary Research 41: 265–277.
Summer, R. M. 1982. Field and laboratory studies on soil erodibility, southern Rocky Mountains, Colorado. Earth Surfaces Processes and Landforms 7: 253–266.
Swanson, F. J., Kratz, T. K. Caine, N. and Woodmansee. R. G. 1988. Landform effects onecosystem patterns and processes. Bioscience 38: 92–98.
Thorn, C. E. 1975. Influence of late-lying snow on rock-weathering rinds. Arctic and Alpine Research 7: 373–378.
Thorn, C. E. 1976. Quantitative evaluation of nivation in the Colorado Front Range. Geological Society of America Bulletin 87: 1169–1178.
Thorn, C. E. 1978. A preliminary assessment of the geomorphic role of pocket gophers in the alpine zone of the Colorado Front Range. Geografiska Annaler 60A: 181–187.
Thorn, C. E. 1982. Gopher disturbance: Its variability by Braun-Blanquet vegetation units in the Niwot Ridge alpine tundra zone, Colorado Front Range, U. S. A. Arctic and Alpine Research 14: 45–51.
Thorn, C. E., and Darmody. R. G. 1985. Grain-size distribution of the insoluble component of contemporary eolian de-posits in the alpine zone, Front Range, Colorado, U. S. A. Arctic and Alpine Research 17: 433–442.
Thorn, C. E., and Loewenherz. D. S. 1987. Alpine mass wasting in the Indian Peaks area, Front Range, Colorado. In Geomorphic systems of North America, edited by Graf. W. L. Boulder, CO: Geological Society of America, Centennial Special Vol. 2.
Wahlstrom, E. E. 1947. Cenozoic physiographic history of the Front Range, Colorado. Geological Society of America Bulletin 58: 551–572.
Wahrhaftig, C. 1987. Foreword to Rock glaciers, edited by Giardino, J. R. Shroder, J. F. and Vitek. J. D. Winchester, MA: Allen and Unwin.
Waldrop, H. A. 1964. The Arapaho Glacier: A sixty-year record. University of Colorado Studies. Series in Geology 3.
Warburton, J. 1991. Absence of frost sorting at an experimental site, Green Lakes Valley, Colorado Front Range, U. S. A. Permafrost and Periglacial Processes 2: 113–122.
Wernicke, B., Clayton, R. Ducea, M. Jones, C. H. Park, S. Ruppert, S. Saleeby, J. Snow, J. K. Squires, L. Fliedner, M. Jiracek, G. Keller, R. Klemperer, S. Luetgert, J. Malin, P. Miller, K. Mooney, W. Oliver, H. and Phinney. R. 1996. Origin of high mountains in thecontinents: The southern Sierra Nevada. Science 271: 190–193.
White, S. E. 1968. Rockfall, alluvial and avalanche talus in the Colorado Front Range. Geological Society of America Special Papers 115: 237.
White, S. E. 1971. Rock glacier studies in the Colorado Front Range, 1961 to 1968. Arctic and Alpine Research 3: 43–64.
White, S. E. 1972. Alpine subnival boulder pavements in Colorado Front Range. Geological Society of America Bulletin 83: 195–200.
White, S. E. 1976. Rock glaciers and block fields: Review and new data. Quaternary Research 6: 77–97.
White, S. E. 1981. Alpine mass movement forms (noncatastrophic): Classification, description and significance. Arctic and Alpine Research 13: 127–137.
White, S. E. 1982. Physical and geological nature of the Indian Peaks, Colorado Front Range. In Ecological studies in the Colorado alpine: A festschrift for John W. Marr, edited by Halfpenny. J. C. Occasional Paper 37, Institute of Arctic and Alpine Research, University of Colorado, Boulder.
Williams, M. W., Brown, A. and Melack. J. M. 1993. Geochemical and hydrologic controlson the composition of surface water in a high-elevation basin, Sierra Nevada. Limnology and Oceanography 38: 775–797.
Williams, M. W., Baron, J. S. Caine, N. Sommerfeld, R. and Sanford. R. 1996. Nitrogen sat-uration in the Rocky Mountains. Environmental Science and Technology 30: 640–646.
Zielinski, G. A., and Davis. P. T. 1987. Late Pleistocene age for the type Temple Lakemoraine, Wind River Range, Wyoming, U. S. A. Geographie Physique et Quaternaire 41: 397–401.