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Structure and Function of an Alpine EcosystemNiwot Ridge, Colorado$

William D. Bowman and Timothy R. Seastedt

Print publication date: 2001

Print ISBN-13: 9780195117288

Published to Oxford Scholarship Online: November 2020

DOI: 10.1093/oso/9780195117288.001.0001

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Geomorphic Systems of Green Lakes Valley

Geomorphic Systems of Green Lakes Valley

Chapter:
(p.45) 4 Geomorphic Systems of Green Lakes Valley
Source:
Structure and Function of an Alpine Ecosystem
Author(s):

Nel Caine

Publisher:
Oxford University Press
DOI:10.1093/oso/9780195117288.003.0005

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.

Keywords:   Avalanches, Blockfield zone, Cliffs, Debris flows, Erosion, Fine sediment system, Geochemical denudation, Kiowa Peak, Landforms

Introduction

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.

Geomorphic Systems of Green Lakes Valley

Figure 4.1. Location Maps of the Indian Peaks Areas and Green Lakes Valley. Sites within Green Lakes Valley are identified: Ak, Arikaree; N, Navajo; G5, Green Lake 5; G4, Green Lake 4; G3, Green Lake 3; Gl, Green Lake 1; I, Lake Albion Inlet; S, Lake Albion Spillway; M, Martinelli Snowpatch; Al, Albion. The Green Lakes are numbered 1-5.

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

Elevation(m)

Area(ha)

Relief(m)

Relative Relief(m/ha)

Mean Slope (°)

Channel Length(km)

Stream Order

Drainage Density(m/ha)

Stream Gradient (°)

Lake Area(ha)

Arikaree

3790

9.2

220

32.6

36.0

0.35

1

38

6.4

0.2

Navajo

3730

41.8

360

8.6

29.1

1.14

2

27

4.4

0.2

Green Lake 5

3620

135.2

470

3.5

22.0

4.65

3

34

7.6

4.3

Green Lake 4

3550

220.9

540

2.4

20.0

9.33

3

42

3.6

8.9

Green Lake 3

3450

306.0

640

2.1

20.1

10.43

3

34

6.9

16.4

Albion Inlet

3345

354.7

745

2.1

21.6

11.65

3

33

7.6

23.2

Spillway

3345

546.2

745

1.4

17.7

16.18

3

30

0.0

44.8

Albion

3250

709.7

840

1.2

17.5

25.67

3

36

2.5

55.0

Green Lake 1

3415

47.3

235

5.0

23.6

0.55

2

12

13.2

4.1

Martinelli

3410

8.0

155

19.4

18.0

0.35

1

44

8.1

0.0

Geomorphic Systems of Green Lakes Valley

Figure 4.2. Catchment characteristics in the Green Lakes system, (a) Basin form (L/W) andbasin area; (b) valley step spacing and downvalley distance; (c) basin relief and basin area;(d) relative relief and basin area; (e) mean slope and basin area; (f) stream gradient and down-valley distance; (g) channel length and basin area; (h) drainage density and basin area; (i) to-tal lake area and basin area; and (j) relative lake area and basin area.

Geomorphic Systems of Green Lakes Valley

Figure 4.3. Relative relief and basin area for the Green Lakes system. The power relationfitted by least squares has the form: RR = 82.49 A−0.646 with R2 = 0.981.

Geomorphic Systems of Green Lakes Valley

Figure 4.4. Aerial oblique view of Niwot Ridge from the northeast. The valley of South St, Vrain Creek is in the foreground with the Green Lakes Valley beyond the broad interfluve of Niwot Ridge. Note the correspondence in elevation between Niwot Ridge and the saddle be-tween Kiowa Peak and Mount Albion and the common elevation of the summits along the Continental Divide and along the Gore Range (background) (photo by Nel Caine).

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).

Geomorphic Systems of Green Lakes Valley

Figure 4.5. Relict patterned ground on the crest of Niwot Ridge at the D1 station. Buildings(right) and jeep road give scale of polygons (photo by Nel Caine).

Geomorphic Systems of Green Lakes Valley

Figure 4.6. Relict sorted patterns on Niwot Ridge 200 m west of the Saddle research site (photo by Nel Caine).

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.

Geomorphic Systems of Green Lakes Valley

Figure 4.7. Talus beneath the north side of Kiowa Peak (photo by Nel Caine).

Geomorphic Systems of Green Lakes Valley

Figure 4.8. Talus conebelow couloir on thesouth face of Niwot Ridge. Green Lake 5 is inthe foreground (photo by Nel Caine).

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

Elevation(m)

Area(ha)

Da(m)

Dbmax(m)

Volume (*106 m3)

Length(m)

Width(m)

D Vc(m)

Green Lake 5

3620

40

4.0

8.0

1.21

900

250

1.50

Green Lake 4

3550

53

4.1

13.0

2.15

850

240

0.93

Green Lake 3

3450

75

8.6

16.0

6.47

600

350

1.62

Green Lake 2

3400

68

7.7

16.0

5.23

700

240

1.44

Green Lake 1

3425

41

3.7

7.0

1.50

500

440

1.59

Lake Albion

3345

216

6.0

15.0

13.00

1000

850

1.20

Sources: Data derived from surveys of McNeely (1983) and Harbor (personal communication).

The Development of the Landscape

Pre-Quaternary History

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.

Pleistocene Glaciation

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).

Geomorphic Systems of Green Lakes Valley

Figure 4.9. Pleistocene glaciation in the North Boulder Creek Basin. Sources: Benedict (1973); Madole (1976); Richmond (1986); Davis (1988).

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).

Geomorphic Systems of Green Lakes Valley

Figure 4.10. Arikaree Glacier: aerial oblique from the northeast. The moraines at the footof the southern half of the glacier are of “Triple Lakes” age, with late-Holocene glacial de-posits on their proximal face. The relict rock glacier (left) distal to these moraines is of late-Pinedale age (photo by Nel Caine).

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

  Rockfall

10.0

44.6

  Talus accumulation

1.5

2.87

  Talus shift

206, 500

14.43

  Debris flow

9.4

15.32

  Rock glacier flow

373, 000

3.59

  Subtotal

579, 521

80.81

Fine sediment system

  Soil creep and solifluction

95, 440

0.90

  Surficial wasting

240

0.03

  Lake sedimentation

13.5

1.00

  Fluvial sediment export

0.38

0.05

  Eolian transport

20

?

  Subtotal

95, 714

2.25

Geochemical system

  Solute transport

1.9

13.36

Source: Caine (1986) with selected updates and additions.

Geomorphic Systems of Green Lakes Valley

Figure 4.11. Alpinegeomorphic systems.

Source: Caine (1974).

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.

Geomorphic Systems of Green Lakes Valley

Figure 4.12. Mass budget of the debris flow of June 26, 1988. The flow, on the south val-ley wall between Green Lake 4 and Green Lake 5, occurred in response to a 52-mm rainstorm(recurrence interval of about 100 years in the Green Lakes area). Inset is longitudinal profileof the flow (true scale).

Geomorphic Systems of Green Lakes Valley

Figure 4.13. Debrisflow deposits on thesouth valley wall between Green Lake 4and Green Lake 5. This flow occurred on June26, 1988 as a responseto a rainstorm that produced 52 mm of precip-itation in 8 hours.Previous debris flow activity on this slope occurred on July 15, 1965 in response to 21 mm ofrain in 4 hours (photoby Nel Caine).

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).

(p.62) The Fine Sediment System

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

Elevation

Zone

Source

Highest

Glacier

Active glaciers, perennial snow, ice

1, 2

Blockfield zone

Bedrock, blockfield, talus

3, 4

Sorted patterned ground, “unbound” solifluction

3, 4

Solifluction zone

Hummocks

3

Vegetated solifluction forms, micro-terracettes

5

Lowest

Forest zone

Treeline, krummholz

Sources: 1. Rapp and Rudberg (1960); 2. Caine (1978); 3. Rudberg (1972); 4. Hollerman (1967); and 5. Hastenrath (1974).

(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

T.D.S.

Cations

SiO2

Cl

Water (mm)

Sediment

Precipitation

5.63

0.58

0.01

0.23

1218

1.00a

Arikaree

15.76

2.51

1.58

0.32

3845

Navajo

8.89

1.71

1.27

0.25

1545

0.27

Green Lake 5

7.22

1.51

0.82

0.19

984

Green Lake 4

8.46

1.76

1.07

0.20

868

0.44

Albion Inlet

6.93

1.66

0.62

0.12

592

Spillway

7.84

1.87

0.60

0.11

601

Albion

10.85

2.45

1.24

0.25

640

0.59

Green Lake 1

4.13

0.97

0.38

0.03

91

Martinelli

9.97

1.91

1.59

0.22

1206

0.54

Materials Budgets

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.

Morphodimatic Landforms

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

Interfluves

Limit (m)

Valley Locations

Limit (m)

Glacier zone

Arikaree glacier

3800

Blockfield zone

Autochthonous blockfields

3700

Tongue-shaped rock glacier

?a

Allochthonous (?) blockfields

3700

Lobate rock glacier

3650

Rockfall talus

3200

Solifluction zone

Sorted Patterned ground

3625

Sorted patterned ground

3550

Non-sorted circles

3500

Non-sorted circles

3450

Stone-banked terraces

3475

Sub-nival pavements

3425

Hummocks

3450

Hummocks

3350

Turf-banked lobes and terraces

3400

Sources: Outcalt and MacPhail (1965), Benedict (1970), White (1971), Madole (1972), Fahey (1975), and Warburton (1991).

(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.

Conclusion

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.

(p.69) References

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