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Dramatic increase in late Cenozoic alpine erosion rates recorded by cave sediment in the southern Rocky Mountains
Earth and Planetary Science Letters 297 (2010) 505–511
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Earth and Planetary Science Letters j o u r n a l h o m e p a g e : w w w. e l s ev i e r. c o m / l o c a t e / e p s l
Dramatic increase in late Cenozoic alpine erosion rates recorded by cave sediment in the southern Rocky Mountains Kurt A. Refsnider Institute of Arctic and Alpine Research and Department of Geological Sciences, University of Colorado, Boulder, 1560 30th St UCB 450, Boulder, Colorado 80309, United States
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Article history: Received 25 May 2010 Received in revised form 30 June 2010 Accepted 2 July 2010 Available online 24 July 2010 Editor: T.M. Harrison Keywords: erosion rate cosmogenic nuclide cave Rocky Mountains Pliocene and Pleistocene climate periglacial weathering
a b s t r a c t Apparent increases in sedimentation rates during the past 5 Ma have been inferred at sites around the globe to document increased terrestrial erosion rates, but direct erosion rate records spanning this period are sparse. Modern and paleo-erosion rates for a small alpine catchment (3108 m above sea level) in the Southern Rocky Mountains are measured using the cosmogenic radionuclides (CRNs) 10Be and 26Al in cave sediment, bedrock on the overlying landscape surface, and coarse bedload in a modern fluvial drainage. The unique setting of the Marble Mountain cave system allows the inherited erosion rates to be interpreted as basin-averaged erosion rates, resulting in the first CRN-based erosion rate record from the Rocky Mountains spanning 5 Myr. Pliocene erosion rates, derived from the oldest cave sample (4.9 ± 0.4 Ma), for the landscape above the cave are 4.9 ± 1.1 m Myr− 1. Mid Pleistocene erosion rates are nearly an order of magnitude higher (33.1 ± 2.7 to 41.3 ± 3.9 m Myr− 1), and modern erosion rates are similar; due to the effects of snow shielding, these erosion rate estimates are likely higher than actual rates by 10–15%. The most likely explanation for this dramatic increase in erosion rates, which likely occurred shortly before 1.2 Ma, is an increase in the effectiveness of periglacial weathering processes at high elevations related to a cooler and wetter climate during the Pleistocene, providing support for the hypothesis that changes in late Cenozoic climate are responsible for increased continental erosion. © 2010 Elsevier B.V. All rights reserved.
1. Introduction The nature of modern landscapes reflects the cumulative effects of tectonic, erosional, and sedimentary processes, but determining the contribution of each individual component from the geologic record is rarely straightforward. Changes in erosion rates during the late Cenozoic have been documented in or inferred from a broad range of settings and archives around the globe (Hay et al., 1988; Zhang et al., 2001; Molnar, 2004; Schaller et al., 2004; Balco and Stone, 2005; Schuster et al., 2005; Häuselmann et al., 2007). However, the interpretation of many records is complicated by changes in sediment budgets, basin accommodation space, sediment sources, and processes with long recurrence intervals (Sadler, 1981; Schumer and Jerolmack, 2009), so confidently identifying variations in erosion rates and attributing such variations to specific processes presents a serious challenge. Whether or not surface erosion rates increased during the late Cenozoic in the Southern Rocky Mountains remains an open question. Taking advantage of a unique geologic setting insulated from many of sedimentation-related complications mentioned above, I present modern and paleo-erosion rates spanning the past 5 Myr derived
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from a small, non-glacial alpine catchment (0.25 km2 averaging 3108 m above sea level; asl) in the otherwise glacially-sculpted Sangre de Cristo (SdC) Range of southern Colorado (Fig. 1A) using cosmogenic radionuclides (CRNs). Paleo-erosion rates are recorded by detrital sediment washed into a cave system situated below this catchment, and modern erosion rates are determined from surface bedrock and stream sediment within the catchment. The ability to interpret the cave sediment as reflecting basin-averaged paleoerosion rates yields the first CRN-based late Cenozoic terrestrial erosion rate record from the Rocky Mountains. 2. Marble Mountain cave system Marble Mountain (4043 m asl), located in the northern SdC Range, holds a cave network with some of the highest solution caves in North America, ranging in elevation from ~ 3500 to 3700 m above sea level (Fig. 1B). The SdC Range, situated along the eastern margin of the Rio Grande Rift, was exhumed beginning in the late Oligocene (Lindsey et al., 1986). The range is predominantly composed of Paleozoic sedimentary rocks. Carbonate strata (Pennsylvanian–Permian SdC Formation) only crop out on the east slope of Marble Mountain above tree line (3400 m asl locally), and above this limestone, the upper 200 m of the mountain are composed of arkosic sandstone and conglomerate (Fig. 2; Johnson et al., 1987). Deep glacial valleys have been carved along the north and west flanks of Marble Mountain
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Fig. 1. Location map and cross sections of the study area. (A) Marble Mountain is located in the Sangre de Cristo Range and above the Wet Mountain Valley. The location of Spanish Cave is denoted with the black circle. Cross sections shown in (B) are indicated by the solid and dashed lines. The approximate extents of glaciers in the South Colony and Sand Creek drainages during the Last Glacial Maximum are shown in dark grey (Refsnider et al., 2009). The polygon indicates the view shown in Fig. 2. The square in the inset shows the area in (A), and the line between the study area and Westcliffe denotes the location of Promontory Divide. (B) Cross sections through Marble Mountain showing Spanish Cave (grey lines), cave sediment (open circles) and surface (filled circles) sampling sites.
(Refsnider et al., 2009), but there is no evidence that the eastern side of the mountain has ever been glaciated, and the caves are not situated in a favorable location to have been affected by glacier meltwater streams. The largest cave in Marble Mountain is Spanish Cave (N2000 m of passage; Fig. 2), and there are at least nine other small caves b300 m in length. Passages are generally only a few meters across, with the largest rooms rarely exceeding 5 m across. The majority of the passages in Spanish Cave exhibit classic keyhole geometry with an upper phreatic tube that was subsequently incised by vadose flow as the water table dropped, resulting in sinuous canyons up to 25 m in depth. The upper 100 m of Spanish Cave, however, have a characteristically different morphology with fewer distinct phreatic tubes and considerably steeper passages (Figs. 1B and 2), though some of the highest passages in the cave system are phreatic tubes (Davis, 1960). Passages in the smaller caves are predominantly phreatic tubes with limited vadose entrenchment. Most of the caves are linked and once formed part of a larger system that has been
truncated by lowering of the landscape surface (Davis, 1960; Wilson and Withrow, 2000). Detrital sediment is found in several isolated locations within the cave system and contains predominantly limestone fragments, limey clay, and rounded feldspar and quartz clasts. The silicic minerals must be derived from surficial erosion and inwash of the overlying arkosic sandstone and conglomerate. The area of the modern watershed from which surficial material sourced for Spanish Cave is approximately 0.25 km, 75% of which is above the limestone–sandstone contact (Fig. 2). 3. Methods Quartz-bearing sediment samples were collected from three sites within the Marble Mountain cave system, as well as two sites on the surface above the caves (Figs. 1B and 2, Table 1). The Sand Crawl sample is from the floor of a horizontal phreatic tube at the boundary between the more vertical upper reaches of Spanish Cave and the more horizontal phreatic tubes of the lower part of the cave. A sample
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Fig. 2. View of the east face of Marble Mountain from the northeast showing cave passages and sampling locations. Solid and open circles denote surface and cave sampling sites, respectively. The dashed lines approximate the lower and upper boundaries of the limestone in which the caves are located. The thin solid black lines are 100-m contour intervals.
from T-Slot Crawl was collected from the edges of a phreatic tube incised subsequent to sediment deposition. These two samples were both buried more than 60 m below the landscape surface. The two passages of Burns Cave, at the lowest known level of cave formation, are filled with at least 1 m of fine-grained, horizontally-bedded sediment, and the third cave sample comes from the coarsest of these sediments, buried beneath approximately 15 m of bedrock. Samples from the landscape surface above the caves include a conglomeratic sample from the summit ridge and an amalgamation of ten sandstone and conglomerate clasts 4 to 8 cm in diameter from the dry streambed in Spanish Gulch above Burns Cave. No evidence of landslides or other mass movements is apparent on the eastern face of Marble Mountain. The mineralogy of the b2 mm size fraction of the cave sediment samples and the bulk surface bedrock and clast samples is determined using quantitative phase analysis of x-ray diffraction spectra with RockJock v. 11 (Srodon et al., 2001; Eberl, 2003). Uncertainties are b5%, though all detected phases must be present in at least trace concentrations.
Samples were prepared for isotopic measurement at the University of Colorado Cosmogenic Isotope Lab, and isotope ratios were measured at LLNL-CAMS. Local production rates are calculated using sea level, high-latitude 10Be and 26Al production rates of 4.41 ± 0.51 −1 and 29.77 ± 3.51 atoms g SiO−1 , respectively (2009 update to 2 yr Balco et al., 2008) and scaled for latitude and altitude using the model of Desilets et al. (2006). Site temperature and atmospheric pressure are calculated from NCEP reanalysis data (cf. Balco et al., 2008), and geomagnetic field intensity is taken as modern for the surface samples and the 800-ka mean intensity (Guyodo and Valet, 1999) for burial samples. CRN production, as well as associated uncertainties, due to nucleon spallation (96.39% and 95.44% of surface production rates for 10 Be and 26Al, respectively), slow muon interactions (1.92% and 2.42% for 10Be and 26Al, respectively), and fast muon interactions (1.69% and 2.15% for 10Be and 26Al, respectively) are calculated based on the results of Heisinger et al. (2002a,b). Particle attenuation lengths are taken as 160 ± 10, 1510 ± 100, and 4320 ± 500 g cm−2 for nucleons, slow muons, and fast muons, respectively (Gosse and Phillips, 2001;
Table 1 CRN-derived burial ages and erosion rates. Sample ID
Location
Sample type
Elevation/[depth*] (m)
[26Al]† (106 1 at g SiO− 2 )
[10Be]† (106 1 at g SiO− 2 )
[26Al]/ [10Be]
Burial age§ (Ma)
Erosion rate§ (m Ma− 1)
KAR07-MM4 KAR08-MM6 KAR08-MM2 KAR08-MM8# KAR08-MM10
Sand Crawl T-slot Crawl Burns Cave Marble Mtn summit ridge Spanish Gulch drainage
Sand Sand Sand Bedrock Cobbles
[95] [114] [206] 3930 3801 mean
0.470 ± 0.089 3.369 ± 0.207 2.450 ± 0.169 9.445 ± 0.555 5.355 ± 0.316
0.763 ± 0.013 0.842 ± 0.014 0.641 ± 0.016 1.753 ± 0.017 0.844 ± 0.019
0.62 ± 0.12 4.00 ± 0.25 3.82 ± 0.28 5.39 ± 0.32 6.34 ± 0.40
4.91 ± 0.41 (0.46) 1.10 ± 0.14 (0.19) 1.20 ± 0.16 (0.21) N.A. N.A.
4.85 ± 1.09 (1.30) 33.07 ± 2.73 (5.25) 41.28 ± 3.85 (6.79) 24.74 ± 1.01 (3.49) 45.41 ± 1.82 (6.45)
*Depth represents the vertical distance below the upper entrance of Spanish Cave. † Blank-corrected concentrations. ~ 0.25 mg of 9Be carrier was added to ~ 100 g quartz for burial samples and ~ 40 g quartz for surface samples. Quartz [Al] was measured by ICP-OES (5% measurement uncertainty). 26Al/27Al and 10Be/9Be ratios were measured at LLNL and normalized to standards KNSTD10650 and 07KNSTD3110, respectively. ±1 standard error for analytical uncertainties are shown. All blank corrections were b 2%. § Age and erosion rate uncertainties represent the ± 1 standard error for analytical uncertainties and rock density; values in parentheses include systematic uncertainties in production rates, decay constants, and particle attenuation lengths. Erosion rates for surface samples are calculated from 10Be data. See text for calculation details. No corrections are made for snow shielding; see discussion in text of how shielding affects calculated erosion rates. # Latitude = 37.951°N, longitude = 105.531°W. Sample thickness is 3 cm, no correction for topographic shielding is necessary, and density is assumed to be 2.5 ± 0.1 g cm− 3.
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Heisinger et al., 2002a,b). Production rates, scaling factors, and calculations are based on half-lives for 10Be and 26Al of 1.36 ± 0.07 and 0.708 ± 0.17 Myr, respectively. The cave depths at which burial samples were deposited are sufficient to prevent appreciable postburial CRN production. The CRN concentration in bare bedrock is a function of the CRN production rates and the rate at which the CRN-rich surface material is removed by erosion (Lal, 1991). Lower erosion rates provide a longer duration for CRNs to accumulate in the rock, resulting in higher CRN concentrations, and higher erosion rates correspond to lower CRN concentrations. The CRN concentrations in cave sediment can thus be inverted to determine an “inherited” erosion rate representative of the landscape surface on a ca. 10-ka time scale prior to burial. Burial ages and inherited erosion rates are calculated using an iterative approach described by Granger et al. (1997) to solve for these two unknown values with the 10Be and 26Al concentrations, and these equations have been expanded to account for production by slow and fast muons. Modern bedrock and basin-averaged erosion rate calculations follow Lal (1991) using his equation 21 expanded to also account for production by slow and fast muon interactions. Two uncertainties are calculated for each burial age and erosion rate value, one including the uncertainty in analytical measurements and estimates of rock density, and a second including additional systematic uncertainties in production rates, decay constants, and particle attenuation lengths, and the fraction of total production due to each production mechanism.
4. Results CRN data are presented in Table 1 and Fig. 3, and cave ages, inherited erosion rates, and modern erosion rates are shown in Table 1. The upper part of the cave system is older than ca. 5 Myr, and the lower half likely formed within the past ca. 1.2. The inherited erosion rate for the oldest sample is 4.9 ± 0.4 m Myr−1, whereas the two younger burial samples have inherited erosion rates of nearly an order of magnitude higher. Bulk mineralogy is similar for all samples (Fig. 4). Erosion rates derived from the two surface samples are broadly similar to the inherited erosion rates of the younger two burial
samples, suggesting that modern erosion rates are similar to erosion rates at ca. 1.2 Ma. It is not surprising that the summit ridge sample has a lower erosion rate than that of the basin-averaged rate derived from the Spanish Gulch sample, which presumably also includes clasts being eroded from steeper slopes below the summit ridge. The 26Al/ 10 Be ratio of the summit ridge sample does not overlap with the steady-erosion line in Fig. 3, which suggests that this sample experienced prolonged burial by snow or sediment and/or eroded by slab exfoliation. The conglomeratic bedrock on the narrow summit ridge crest is horizontally-bedded, and these beds dip moderately and often form planes along which the rock has fractured and weathered. If dislodged, the resulting blocks (10–40 cm thick) could easily tumble down the dip slope, exposing a “new” surface that had already accumulated a CRN inventory. This pattern of erosion can result in an anomalously low 26Al/10Be ratio for a surface sample (Gosse and Phillips, 2001) and is the most likely explanation for the unexpectedly low 26Al/10Be ratio of the summit ridge sample. The fact that the 26Al/ 10 Be ratio of the Spanish Gulch sample overlaps with the steadyerosion line suggests that the majority of the sediment washed into the cave system is not affected by prolonged burial or non-steady styles of erosion. Shielding due to snow is neglected in age and erosion rate calculations but warrants discussion. Cover by snow results in the attenuation of secondary cosmic-ray particles, reducing surface CRN production rates. While burial ages are essentially unaffected by the effects of snow shielding, failing to account for the reduced production rates will result in overestimates of inherited erosion rates. Snow data are available from a SNOTEL station in the South Colony drainage, immediately north of Marble Mountain at 3292 m asl (Fig. 1A), for the past 30 years (available from the Western Climate Data Center; http://www.wrcc.dri.edu). After adjusting CRN production rates for the monthly mean snow water equivalent at this SNOTEL station, the resulting erosion rates are 12–13% lower than those reported in Table 1. Snow cover on the east face of Marble Mountain undoubtedly persisted longer into summers during the LGM, but increases in precipitation were likely not substantial (Leonard, 2007; Refsnider et al., 2009; Brugger, 2010), and there is no evidence that the range crest was glaciated in this section of the SdC (Refsnider et al., 2009). Prior to the LGM, there are no constraints on alpine snow cover in the region. Because of these uncertainties, reported erosion rates are not adjusted for snow shielding but likely overestimate actual rates by at least 10%. The magnitude of this overestimate probably differs for all samples, but the corrections and associated uncertainties should not affect the interpretation of the results discussed below. 5. Discussion 5.1. Water table lowering and paleo-erosion rates
Fig. 3. Two-isotope diagram showing cave samples (open circles) and surface samples (filled circles); error bars show 1σ standard error based on analytical uncertainties. The solid black curve and associated erosion rates track the CRN inventory derived from continuous surface exposure at a constant erosion rate. Upon burial, samples will track downward parallel to the dotted lines, and the dashed burial isochrons indicate the 26 Al/10Be ratio after a period of continuous burial.
Detrital sedimentary deposits in caves have previously been used to infer changes in fluvial erosion rates in systems where the depth of shallow phreatic cave passage formation is controlled by water table lowering due to stream incision (Granger et al., 1997, 2001; Stock et al., 2004; Häuselmann et al., 2007). The elevation of the vadose– phreatic transition in Marble Mountain is not directly controlled by the elevation of a single local stream. While the boundaries of this watershed have undoubtedly evolved over the past 5 Myr, its areal extent must have been restricted to between the caves and the summit ridge. The Marble Mountain cave system is therefore unique in that few geologic archives allow inherited CRN erosion rates to be interpreted as basin-averaged erosion rates, essentially analogous to the Spanish Gulch surface sediment sample. The morphology of the Spanish Cave system, which can be divided into three sections vertically based on passage development, provides support for the cave passage age estimates and the paleo-erosion rates
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Fig. 4. Mineralogy of the b2 mm size fraction of cave sediment samples and bulk surface bedrock and clast samples. All analyses have degrees of fit ≤0.12. The mineralogical segmentation of each bar matches the same order as the minerals listed in the legend. These data suggest that there have been no dramatic changes in the lithology of material being eroded and washed into the cave system.
derived from the detrital sediment. The uppermost section above Sand Crawl (0 to 95 m depth below the upper Spanish Cave entrance), is predominantly vadose passages with only a few phreatic tubes. The section below Sand Crawl (95 to 135 m depth), characterized by incised phreatic tubes connecting small rooms, is where the majority of passage development took place in both Spanish Cave and White Marble Halls. T-Slot Crawl (114 m depth) is in the middle of this section. The main entrance passages to these caves are the lowest extensive phreatic tubes in this section. In the lower section of Spanish Cave, only one additional phreatic tube level has been found, which is 90 m lower and at the same elevation as Burns Cave (206 m depth). Spanish Spring is approximately 40 m below this lowest level. The extensive vadose enlargement and development of the upper section of the cave system suggests considerable antiquity, in general agreement with the ca. 5 Ma age from Sand Crawl. The 1.1 ± 0.1 Ma age from only 19 m lower in the cave system at T-Slot Crawl suggests that the water table remained at a similar elevation for several million years, allowing considerable time for the extensive network of phreatic passages in the middle section to develop. Subsequently, the water table dropped rapidly to the level of Burns Cave as indicated by the overlapping ages from T-Slot Crawl and Burns Cave. This lowering of the vadose–phreatic transition was rapid enough that virtually no phreatic passage development occurred. The pattern of water table lowering recorded by both the cave morphology and sediment burial ages is related to springs controlling the local base level (Klimchouk et al., 2000). The elevation of Spanish Spring (Fig. 1B) and springs at similar elevations near the valley floor on the west side of Marble Mountain presumably control the level of the modern water table. Phreatic tubes within the cave system thus record higher base levels within the watershed, and the rate of water table lowering must somehow be related to landscape lowering rates. The prolonged period of relative water table stability between Sand Crawl and T-Slot Crawl is in accordance with the low surface erosion rates during between ca. 5 and 1.2 Ma. Rapid lowering of the water table from the level of T-Slot Crawl to Burns Cave corroborates the subsequent increase in surface erosion rates, and the level of the spring controlling the base level is now approximately 40 m lower than Burns Cave. These morphological relationships and ages suggest that the network of phreatic passages in the middle section of the cave system developed between ca. 5 and 1.2 Ma, and the dramatic increase in erosion rates most likely occurred shortly before 1.2 Ma. Modern and Pleistocene basin-averaged and summit ridge erosion rates from Marble Mountain are similar to the CRN-derived erosion rates (all reflecting ca. 10-ka time scales) measured from other ranges in the Rocky Mountains comprised predominantly of crystalline
bedrock. Bedrock ridges in the Teton Range of Wyoming are eroding at ~22 m Myr−1 (Tranel et al., 2009), and basin-averaged erosion rates in unglaciated catchments in the Front, Medicine Bow, and Laramie ranges of northern Colorado and southern Wyoming are 18– 30 m Myr−1 (Dethier et al., 2002). Basin-averaged erosion rates from non-glacial mountain catchments in Idaho range from 20 to 100 m Myr−1 (Kirchner et al., 2001). Basin-averaged erosion rates from steep, unglaciated catchments in the core of the Wasatch Range of Utah range from 170 to 790 m Myr−1, and these higher and more variable rates are likely due to glacial erosion in main canyons steepening the profile of the non-glacial tributary drainages from which samples were collected (Stock et al., 2009). Interestingly, the low Pliocene erosion rate recorded by the oldest Spanish Cave sample is more similar to bare-bedrock erosion rates (generally b10 m Myr−1) on Rocky Mountain summit flats (Small et al., 1997) and on ridges, hilltops, and summits across the eastern United States (Bierman et al., 1995; Granger et al., 2001; Hancock and Kirwan, 2007) than to modern basinaveraged erosion rates in the Rocky Mountains (Dethier et al., 2002; Stock et al., 2009). Estimates of mean fluvial incision rates in the mountains of central Colorado over the past 0.64 Myr are in the range of 100 to 150 m Myr−1 (Dethier, 2001), considerably higher than the Marble Mountain erosion rates. The result of this contrast is an increase in landscape relief, which both supports the conclusion of Small and Anderson (1998) that modern relief in the Rocky Mountains is currently increasing and suggests that such an increase has been underway for at least 0.6 Myr if alpine erosion rates have been relatively constant over this interval. 5.2. The erosion rate increase The late Cenozoic increase in erosion rates of nearly an order of magnitude for Marble Mountain can be explained by a combination of four potential factors: (1) a change in the bedrock lithology being eroded, (2) increases in local glacial and fluvial incision rates, (3) an increase in surface uplift rates due to local tectonism, or (4) a change in weathering processes. The first two factors are likely of little importance in this setting. Strata dip moderately to the west, so the eastern slope of Marble Mountain exposes a sequence of sandstone and conglomerate units. While it is possible that more easily erodible beds may have previously been exposed, the b2 mm size fractions of each CRN sample have similar mineralogical compositions (Fig. 4). This implies that there have been no dramatic changes in the composition of bedrock being eroded on the upper slopes of Marble Mountain. While
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there is evidence for mid to late Pleistocene increases in fluvial incision rates of major rivers in the Southern Rocky Mountains immediately to the east (Dethier, 2001 and references therein), Marble Mountain is situated above Promontory Divide in the Wet Mountain Valley (Fig. 1A), and waves of incision advancing south from the Arkansas River drainage and north from the Huerfano River drainage have not yet begun to dissect the Plio–Pleistocene valley fill near this divide. Glacial scouring in valleys to the north and west of Marble Mountain should not have had a direct effect on the unglaciated east-facing slopes above the Spanish Cave system. The west edge of the SdC Range follows the SdC Fault (Fig. 1A), which locally forms the eastern boundary of the Rio Grande Rift. On the basis of faulted alluvial fan surfaces from at least three glaciations, McCalpin (1983, 1986) estimated mean vertical slip rates of 0.1 to 0.2 mm yr−1 (50 to 100 m Myr−1) for the SdC Fault. Seismic stratigraphy and well log correlations show units within the Santa Fe Group (representing the past 20 Myr) consistently thickening toward the eastern margin of the San Luis Valley (Brister and Gries, 1994). Alamosa Formation sediments (3.0 to 0.4 Ma) also thicken to the east (Machette et al., 2007), implying a relatively continuous increase in accommodation space in the basin during the late Cenozoic. There is also evidence for a shift in the focus of tectonic activity farther south along the fault based on geomorphic features best explained by higher slip rates along the northern end of the SdC Range prior to the late Pleistocene (Ruleman and Machette, 2007). Thus, long-term mean slip rates along the SdC Fault immediately west of Marble Mountain have likely not increased substantially and may even have decreased since the mid Pleistocene. A shift in weathering regimes related to climatic change since the mid Pliocene provides the most reasonable mechanism to explain the increase in erosion rates. Pliocene climate in the Southern Rocky Mountains was characterized by more humid conditions and lower seasonality than in the Pleistocene (Thompson, 1991). Rogers et al. (1992) reconstructed San Luis Valley climatic conditions between 2.6 and 0.7 Ma and identified a gradual transition to cooler temperatures (ca. 3 °C above modern) and increasing aridity beginning at ~ 2.5 Ma and continuing until 1.6 Ma. Conditions 3 to 5 °C warmer and more arid than modern persisted between 1.6 and ~ 0.93 Ma and were followed by the onset of high-amplitude glacial–interglacial oscillations in both temperature and moisture. Evidence of increased runoff from upland sources beginning at ~ 1 Ma likely indicates the expansion of glaciers and seasonal snowpack in the SdC Range and San Juan Mountains, and a 3- to 6-fold increase in sedimentation rates in the San Luis Valley between 2 and 1 Ma may be related to the onset of extensive glaciation in these mountains (Rogers et al., 1992). The apparent rapid water table lowering and vadose incision within the Spanish Cave system indicated by the overlapping ages from T-Slot Crawl and Burns Cave at ca. 1.2 Ma is best explained by the intensification of glaciation and valley floor lowering on the west (where springs are located near the modern valley floor) and north sides of Marble Mountain (e.g., Häuselmann et al., 2007). Cooler temperatures during the mid to late Pleistocene would result in bedrock at high elevations in the SdC Range experiencing more time within the frost-cracking window (−3 to −8 °C; Anderson, 1998), and increased moisture from snowmelt could accelerate periglacial weathering rates and the downslope transport of eroded material by frost creep, gelifluction, and fluvial action. Erosion rates in at least some small alpine catchments in the Alps increase with elevation, which Delunel et al. (2010) interpret as being dominantly due to frost-cracking processes. The mean modern winter (October to April) temperature for the Marble Mountain catchment is −2.7 °C (cf. Refsnider et al., 2009), near the upper limit of the frost-cracking window. Temperatures of at least 3 °C warmer than present would substantially reduce the duration during which bedrock within this catchment would be within the frost-cracking window. Additionally, landscapes may also be unable to reach a state of equilibrium with
climate in the presence of millennial-scale climate variability, resulting in continuous adjustments in erosion rates (Zhang et al., 2001; Molnar, 2004; Schaller et al., 2004). 6. Conclusions The Marble Mountain record of paleo-erosion rates provides the first direct, CRN-based evidence for dramatic erosion rate increases in the Rocky Mountains between 4.9 and 1.2 Ma, and based on passage development and morphology, this increase likely occurred shortly before 1.2 Ma. This record also suggests that modern alpine erosion rates are similar to rates at ca. 1.2 Ma. Because this record comes from a locality relatively insulated from the direct influence of glacial and major fluvial systems, the transition from relatively warm Pliocene conditions into the cooler Pleistocene climate with an associated increase in the effectiveness of periglacial weathering processes at high elevations is the most likely explanation for the observed increase in erosion rates. Acknowledgements This project was made possible with funding from the National Speleological Society. Dylan Rood at LLNL-CAMS provided Be and Al isotope measurements. Skip Withrow, Jim Wilson, and Margaret Barnes were instrumental in sample collection, Skip Withrow shared indispensible cave survey data, Donald Davis and Fred Luiszer provided stimulating speleogenesis discussions, Gifford Miller, Miriam Dühnforth, and Keith Brugger kindly provided comments on the manuscript, and reviews by Darryl Granger and an anonymous referee included insightful and helpful critiques. References Anderson, R.S., 1998. Near-surface thermal profiles in alpine bedrock: implications for the frost weathering of rock. Arc. Alp. Res. 30, 362–372. Balco, G., Stone, J.O.H., 2005. Measuring middle Pleistocene erosion rates with cosmicray-produced nuclides in buried alluvial sediment, Fisher Valley, southeastern Utah. Earth Planet. Sci. Lett. 30, 1051–1067. Balco, G., Stone, J.O., Lifton, N.A., Dunai, T.J., 2008. A complete and easily accessible means of calculating surface exposure ages or erosion rates from 10Be and 26Al measurements. Quat. Geochronol. 3, 174–195. Bierman, P.R., Gillespie, A.R., Caffee, M.W., Elmore, D., 1995. Estimating erosion rates and exposure ages with 36Cl produced by neutron activation. Geochim. Cosmochim. Acta 59, 3779–3798. Brister, B.S., Gries, R.R., 1994. Tertiary stratigraphy and tectonic development of the Alamosa Basin (northern San Luis Valley), Rio Grande rift, south-central Colorado. In: Keller, G.R., Cather, S.M. (Eds.), Basins of the Rio Grande Rift: Structure, Startigraphy, and Tectonic Setting: Geological Society of America Special Paper, vol. 291, pp. 39–58. Brugger, K.A., 2010. Climate in the southern Sawatch Range and Elk Mountains, Colorado, USA, during the Last Glacial Maximum: inferences using a simple degreeday model. Arc. Ant. Alp. Res. 42, 164–178. Davis, D.G., 1960. The Caves of Marble Mountain, Colorado, Described and Located. S. Col. Grotto Spec. Bull., vol. 2. Delunel, R., van der Beek, P.A., Carcaillet, J., Bourlès, D.L., Valla, P.G., 2010. Frost-cracking control on catchment denudation rates: Insights from in situ produced 10Be concentrations in stream sediments (Ecrins–Pelvoux massif, French Western Alps). Earth Planet. Sci. Lett. 293, 72–83. Desilets, D., Zreda, M., Prabu, T., 2006. Extended scaling factors for in situ cosmogenic nuclides: new measurements at low latitude. Earth Planet. Sci. Lett. 246, 265–276. Dethier, D.P., 2001. Pleistocene incision rates in the western United States calibrated using Lava Creek B tepha. Geology 29, 783–786. Dethier, D.P., Ouimet, W., Bierman, P.R., Finkel, R.C., 2002. Long-term erosion rates derived from 10Be in sediment from small catchments, northern Front Range and southern Wyoming. Abstr. Programs - Geol. Soc. Am. 34, 409. Eberl, D.D., 2003. User guide to RockJock: a program for determining quantitative mineralogy from X-ray diffraction data. U.S. Geological Survey, Open File Report 03-78. 40 pp. Gosse, J.C., Phillips, F.M., 2001. Terrestrial in situ cosmogenic nuclides: theory and application. Quat. Sci. Rev. 20, 1475–1560. Granger, D.E., Kirchner, J.W., Finkel, R.C., 1997. Quaternary down-cutting rate of the New River, Virginia, measured from differential decay of cosmogenic 26Al and 10Be in cave-deposited alluvium. Geology 25, 107–110. Granger, D.E., Fabel, D., Palmer, A.N., 2001. Pliocene–Pleistocene incision of the Green River, Kentucky, determined from radioactive decay of cosmogenic 26Al and 10Be in Mammoth Cave sediments. Geol. Soc. Amer. Bull. 113, 825–836.
K.A. Refsnider / Earth and Planetary Science Letters 297 (2010) 505–511 Guyodo, Y., Valet, J.-P., 1999. Global changes in intensity of the Earth's magnetic field during the past 800 kyr. Nature 399, 249–252. Hancock, G., Kirwan, M., 2007. Summit erosion rates deduced from 10Be: implications for relief production in the central Appalachians. Geology 35, 89–92. Häuselmann, P., Granger, D.E., Jeannin, P.-Y., Lauritzen, S.-E., 2007. Abrupt glacial valley incision at 0.8 Ma dated from cave deposits in Switzerland. Geology 35, 143–146. Hay, W.W., Sloan, J.L., Wold, C.N., 1988. Mass/age distribution and composition of sediments on the ocean floor and the global rate of sediment subduction. J. Geophys. Res. 93, 14933–14940. Heisinger, B., Lal, D., Jull, A.J.T., Kubik, P., Ivy-Ochs, S., Knie, K., Nolte, E., 2002a. Production of selected cosmogenic radionuclides by muons: 2. Capture of negative muons. Earth Planet. Sci. Lett. 200, 357–369. Heisinger, B., Lal, D., Jull, A.J.T., Kubik, P., Ivy-Ochs, S., Neumaier, S., Knie, K., Lazarev, V., Nolte, E., 2002b. Production of selected cosmogenic radionuclides by muons 1. Fast muons. Earth Planet. Sci. Lett. 200, 345–355. Johnson, B.R., Lindsey, D.A., Bruce, R.M., Soulliere, S.J., 1987. Reconnaissance geologic map of the Sangre de Cristo Wilderness Study Area, south-central Colorado. U.S. Geological Survey map MF-1635-B. Kirchner, J.W., Finkel, R.C., Riebe, C.S., Granger, D.E., Clayton, J.L., King, J.G., Megahan, W.F., 2001. Mountain erosion over 10 yr, 10 k.y., and 10 m.y. time scales. Geology 29, 591–594. Klimchouk, A., Ford, D.C., Palmer, A.N., Dreybrodt, W. (Eds.), 2000. Speleogenesis: Evolution of Karst Aquifers. National Speleological Society, Huntsville. 527 pp. Lal, D., 1991. Cosmic ray labeling of erosion surface: in-situ nuclide production rates and erosion models. Earth Planet. Sci. Lett. 161, 231–241. Leonard, E., 2007. Modeled patterns of Late Pleistocene glacier inception and growth in the Southern and Central Rocky Mountains, USA: sensitivity to climate change and paleoclimatic implications. Quat. Sci. Rev. 26, 2152–2166. Lindsey, D.A., Andriessen, P.A.M., Wardlaw, B.R., 1986. Heating, cooling, and uplift during Tertiary time, northern Sangre de Cristo Range, Colorado. Geol. Soc. Amer. Bull. 97, 541–561. Machette, M.N., Marchetti, D.W., Thompson, R.A., 2007. Ancient Lake Alamosa and the Pliocene to Middle Pleistocene evolution of the Rio Grande. Rocky Mountain Section Friends of the Pleistocene Field Trip—Quaternary Geology of the San Luis Basin of Colorado and New Mexico. Geological Survey Open-File Report 2007– 1193, pp. 157–167. McCalpin, J.P., 1983. Quaternary geology and of the northern Sangre de Cristo Mountains, south-central Colorado. Col. School. Mines Q. 77, 1–97. McCalpin, J.P., 1986. Quaternary tectonics of the Sangre de Cristo and Villa Grove fault zones. In: Rogers, W.P., Kirkham, R.M. (Eds.), Contributions to Colorado Seismicity and Tectonics—A 1986 Update: Colorado Geological Survey Special Publication, vol. 28, pp. 59–64. Molnar, P., 2004. Late Cenozoic increase in accumulation rates of terrestrial sediment: how might climate change have affected erosion rates? Annu. Rev. Earth Plan. Sci. 32, 67–89.
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Earth and Planetary Science Letters j o u r n a l h o m e p a g e : w w w. e l s ev i e r. c o m / l o c a t e / e p s l
Dramatic increase in late Cenozoic alpine erosion rates recorded by cave sediment in the southern Rocky Mountains Kurt A. Refsnider Institute of Arctic and Alpine Research and Department of Geological Sciences, University of Colorado, Boulder, 1560 30th St UCB 450, Boulder, Colorado 80309, United States
a r t i c l e
i n f o
Article history: Received 25 May 2010 Received in revised form 30 June 2010 Accepted 2 July 2010 Available online 24 July 2010 Editor: T.M. Harrison Keywords: erosion rate cosmogenic nuclide cave Rocky Mountains Pliocene and Pleistocene climate periglacial weathering
a b s t r a c t Apparent increases in sedimentation rates during the past 5 Ma have been inferred at sites around the globe to document increased terrestrial erosion rates, but direct erosion rate records spanning this period are sparse. Modern and paleo-erosion rates for a small alpine catchment (3108 m above sea level) in the Southern Rocky Mountains are measured using the cosmogenic radionuclides (CRNs) 10Be and 26Al in cave sediment, bedrock on the overlying landscape surface, and coarse bedload in a modern fluvial drainage. The unique setting of the Marble Mountain cave system allows the inherited erosion rates to be interpreted as basin-averaged erosion rates, resulting in the first CRN-based erosion rate record from the Rocky Mountains spanning 5 Myr. Pliocene erosion rates, derived from the oldest cave sample (4.9 ± 0.4 Ma), for the landscape above the cave are 4.9 ± 1.1 m Myr− 1. Mid Pleistocene erosion rates are nearly an order of magnitude higher (33.1 ± 2.7 to 41.3 ± 3.9 m Myr− 1), and modern erosion rates are similar; due to the effects of snow shielding, these erosion rate estimates are likely higher than actual rates by 10–15%. The most likely explanation for this dramatic increase in erosion rates, which likely occurred shortly before 1.2 Ma, is an increase in the effectiveness of periglacial weathering processes at high elevations related to a cooler and wetter climate during the Pleistocene, providing support for the hypothesis that changes in late Cenozoic climate are responsible for increased continental erosion. © 2010 Elsevier B.V. All rights reserved.
1. Introduction The nature of modern landscapes reflects the cumulative effects of tectonic, erosional, and sedimentary processes, but determining the contribution of each individual component from the geologic record is rarely straightforward. Changes in erosion rates during the late Cenozoic have been documented in or inferred from a broad range of settings and archives around the globe (Hay et al., 1988; Zhang et al., 2001; Molnar, 2004; Schaller et al., 2004; Balco and Stone, 2005; Schuster et al., 2005; Häuselmann et al., 2007). However, the interpretation of many records is complicated by changes in sediment budgets, basin accommodation space, sediment sources, and processes with long recurrence intervals (Sadler, 1981; Schumer and Jerolmack, 2009), so confidently identifying variations in erosion rates and attributing such variations to specific processes presents a serious challenge. Whether or not surface erosion rates increased during the late Cenozoic in the Southern Rocky Mountains remains an open question. Taking advantage of a unique geologic setting insulated from many of sedimentation-related complications mentioned above, I present modern and paleo-erosion rates spanning the past 5 Myr derived
E-mail address: [email protected]. 0012-821X/$ – see front matter © 2010 Elsevier B.V. All rights reserved. doi:10.1016/j.epsl.2010.07.002
from a small, non-glacial alpine catchment (0.25 km2 averaging 3108 m above sea level; asl) in the otherwise glacially-sculpted Sangre de Cristo (SdC) Range of southern Colorado (Fig. 1A) using cosmogenic radionuclides (CRNs). Paleo-erosion rates are recorded by detrital sediment washed into a cave system situated below this catchment, and modern erosion rates are determined from surface bedrock and stream sediment within the catchment. The ability to interpret the cave sediment as reflecting basin-averaged paleoerosion rates yields the first CRN-based late Cenozoic terrestrial erosion rate record from the Rocky Mountains. 2. Marble Mountain cave system Marble Mountain (4043 m asl), located in the northern SdC Range, holds a cave network with some of the highest solution caves in North America, ranging in elevation from ~ 3500 to 3700 m above sea level (Fig. 1B). The SdC Range, situated along the eastern margin of the Rio Grande Rift, was exhumed beginning in the late Oligocene (Lindsey et al., 1986). The range is predominantly composed of Paleozoic sedimentary rocks. Carbonate strata (Pennsylvanian–Permian SdC Formation) only crop out on the east slope of Marble Mountain above tree line (3400 m asl locally), and above this limestone, the upper 200 m of the mountain are composed of arkosic sandstone and conglomerate (Fig. 2; Johnson et al., 1987). Deep glacial valleys have been carved along the north and west flanks of Marble Mountain
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Fig. 1. Location map and cross sections of the study area. (A) Marble Mountain is located in the Sangre de Cristo Range and above the Wet Mountain Valley. The location of Spanish Cave is denoted with the black circle. Cross sections shown in (B) are indicated by the solid and dashed lines. The approximate extents of glaciers in the South Colony and Sand Creek drainages during the Last Glacial Maximum are shown in dark grey (Refsnider et al., 2009). The polygon indicates the view shown in Fig. 2. The square in the inset shows the area in (A), and the line between the study area and Westcliffe denotes the location of Promontory Divide. (B) Cross sections through Marble Mountain showing Spanish Cave (grey lines), cave sediment (open circles) and surface (filled circles) sampling sites.
(Refsnider et al., 2009), but there is no evidence that the eastern side of the mountain has ever been glaciated, and the caves are not situated in a favorable location to have been affected by glacier meltwater streams. The largest cave in Marble Mountain is Spanish Cave (N2000 m of passage; Fig. 2), and there are at least nine other small caves b300 m in length. Passages are generally only a few meters across, with the largest rooms rarely exceeding 5 m across. The majority of the passages in Spanish Cave exhibit classic keyhole geometry with an upper phreatic tube that was subsequently incised by vadose flow as the water table dropped, resulting in sinuous canyons up to 25 m in depth. The upper 100 m of Spanish Cave, however, have a characteristically different morphology with fewer distinct phreatic tubes and considerably steeper passages (Figs. 1B and 2), though some of the highest passages in the cave system are phreatic tubes (Davis, 1960). Passages in the smaller caves are predominantly phreatic tubes with limited vadose entrenchment. Most of the caves are linked and once formed part of a larger system that has been
truncated by lowering of the landscape surface (Davis, 1960; Wilson and Withrow, 2000). Detrital sediment is found in several isolated locations within the cave system and contains predominantly limestone fragments, limey clay, and rounded feldspar and quartz clasts. The silicic minerals must be derived from surficial erosion and inwash of the overlying arkosic sandstone and conglomerate. The area of the modern watershed from which surficial material sourced for Spanish Cave is approximately 0.25 km, 75% of which is above the limestone–sandstone contact (Fig. 2). 3. Methods Quartz-bearing sediment samples were collected from three sites within the Marble Mountain cave system, as well as two sites on the surface above the caves (Figs. 1B and 2, Table 1). The Sand Crawl sample is from the floor of a horizontal phreatic tube at the boundary between the more vertical upper reaches of Spanish Cave and the more horizontal phreatic tubes of the lower part of the cave. A sample
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Fig. 2. View of the east face of Marble Mountain from the northeast showing cave passages and sampling locations. Solid and open circles denote surface and cave sampling sites, respectively. The dashed lines approximate the lower and upper boundaries of the limestone in which the caves are located. The thin solid black lines are 100-m contour intervals.
from T-Slot Crawl was collected from the edges of a phreatic tube incised subsequent to sediment deposition. These two samples were both buried more than 60 m below the landscape surface. The two passages of Burns Cave, at the lowest known level of cave formation, are filled with at least 1 m of fine-grained, horizontally-bedded sediment, and the third cave sample comes from the coarsest of these sediments, buried beneath approximately 15 m of bedrock. Samples from the landscape surface above the caves include a conglomeratic sample from the summit ridge and an amalgamation of ten sandstone and conglomerate clasts 4 to 8 cm in diameter from the dry streambed in Spanish Gulch above Burns Cave. No evidence of landslides or other mass movements is apparent on the eastern face of Marble Mountain. The mineralogy of the b2 mm size fraction of the cave sediment samples and the bulk surface bedrock and clast samples is determined using quantitative phase analysis of x-ray diffraction spectra with RockJock v. 11 (Srodon et al., 2001; Eberl, 2003). Uncertainties are b5%, though all detected phases must be present in at least trace concentrations.
Samples were prepared for isotopic measurement at the University of Colorado Cosmogenic Isotope Lab, and isotope ratios were measured at LLNL-CAMS. Local production rates are calculated using sea level, high-latitude 10Be and 26Al production rates of 4.41 ± 0.51 −1 and 29.77 ± 3.51 atoms g SiO−1 , respectively (2009 update to 2 yr Balco et al., 2008) and scaled for latitude and altitude using the model of Desilets et al. (2006). Site temperature and atmospheric pressure are calculated from NCEP reanalysis data (cf. Balco et al., 2008), and geomagnetic field intensity is taken as modern for the surface samples and the 800-ka mean intensity (Guyodo and Valet, 1999) for burial samples. CRN production, as well as associated uncertainties, due to nucleon spallation (96.39% and 95.44% of surface production rates for 10 Be and 26Al, respectively), slow muon interactions (1.92% and 2.42% for 10Be and 26Al, respectively), and fast muon interactions (1.69% and 2.15% for 10Be and 26Al, respectively) are calculated based on the results of Heisinger et al. (2002a,b). Particle attenuation lengths are taken as 160 ± 10, 1510 ± 100, and 4320 ± 500 g cm−2 for nucleons, slow muons, and fast muons, respectively (Gosse and Phillips, 2001;
Table 1 CRN-derived burial ages and erosion rates. Sample ID
Location
Sample type
Elevation/[depth*] (m)
[26Al]† (106 1 at g SiO− 2 )
[10Be]† (106 1 at g SiO− 2 )
[26Al]/ [10Be]
Burial age§ (Ma)
Erosion rate§ (m Ma− 1)
KAR07-MM4 KAR08-MM6 KAR08-MM2 KAR08-MM8# KAR08-MM10
Sand Crawl T-slot Crawl Burns Cave Marble Mtn summit ridge Spanish Gulch drainage
Sand Sand Sand Bedrock Cobbles
[95] [114] [206] 3930 3801 mean
0.470 ± 0.089 3.369 ± 0.207 2.450 ± 0.169 9.445 ± 0.555 5.355 ± 0.316
0.763 ± 0.013 0.842 ± 0.014 0.641 ± 0.016 1.753 ± 0.017 0.844 ± 0.019
0.62 ± 0.12 4.00 ± 0.25 3.82 ± 0.28 5.39 ± 0.32 6.34 ± 0.40
4.91 ± 0.41 (0.46) 1.10 ± 0.14 (0.19) 1.20 ± 0.16 (0.21) N.A. N.A.
4.85 ± 1.09 (1.30) 33.07 ± 2.73 (5.25) 41.28 ± 3.85 (6.79) 24.74 ± 1.01 (3.49) 45.41 ± 1.82 (6.45)
*Depth represents the vertical distance below the upper entrance of Spanish Cave. † Blank-corrected concentrations. ~ 0.25 mg of 9Be carrier was added to ~ 100 g quartz for burial samples and ~ 40 g quartz for surface samples. Quartz [Al] was measured by ICP-OES (5% measurement uncertainty). 26Al/27Al and 10Be/9Be ratios were measured at LLNL and normalized to standards KNSTD10650 and 07KNSTD3110, respectively. ±1 standard error for analytical uncertainties are shown. All blank corrections were b 2%. § Age and erosion rate uncertainties represent the ± 1 standard error for analytical uncertainties and rock density; values in parentheses include systematic uncertainties in production rates, decay constants, and particle attenuation lengths. Erosion rates for surface samples are calculated from 10Be data. See text for calculation details. No corrections are made for snow shielding; see discussion in text of how shielding affects calculated erosion rates. # Latitude = 37.951°N, longitude = 105.531°W. Sample thickness is 3 cm, no correction for topographic shielding is necessary, and density is assumed to be 2.5 ± 0.1 g cm− 3.
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Heisinger et al., 2002a,b). Production rates, scaling factors, and calculations are based on half-lives for 10Be and 26Al of 1.36 ± 0.07 and 0.708 ± 0.17 Myr, respectively. The cave depths at which burial samples were deposited are sufficient to prevent appreciable postburial CRN production. The CRN concentration in bare bedrock is a function of the CRN production rates and the rate at which the CRN-rich surface material is removed by erosion (Lal, 1991). Lower erosion rates provide a longer duration for CRNs to accumulate in the rock, resulting in higher CRN concentrations, and higher erosion rates correspond to lower CRN concentrations. The CRN concentrations in cave sediment can thus be inverted to determine an “inherited” erosion rate representative of the landscape surface on a ca. 10-ka time scale prior to burial. Burial ages and inherited erosion rates are calculated using an iterative approach described by Granger et al. (1997) to solve for these two unknown values with the 10Be and 26Al concentrations, and these equations have been expanded to account for production by slow and fast muons. Modern bedrock and basin-averaged erosion rate calculations follow Lal (1991) using his equation 21 expanded to also account for production by slow and fast muon interactions. Two uncertainties are calculated for each burial age and erosion rate value, one including the uncertainty in analytical measurements and estimates of rock density, and a second including additional systematic uncertainties in production rates, decay constants, and particle attenuation lengths, and the fraction of total production due to each production mechanism.
4. Results CRN data are presented in Table 1 and Fig. 3, and cave ages, inherited erosion rates, and modern erosion rates are shown in Table 1. The upper part of the cave system is older than ca. 5 Myr, and the lower half likely formed within the past ca. 1.2. The inherited erosion rate for the oldest sample is 4.9 ± 0.4 m Myr−1, whereas the two younger burial samples have inherited erosion rates of nearly an order of magnitude higher. Bulk mineralogy is similar for all samples (Fig. 4). Erosion rates derived from the two surface samples are broadly similar to the inherited erosion rates of the younger two burial
samples, suggesting that modern erosion rates are similar to erosion rates at ca. 1.2 Ma. It is not surprising that the summit ridge sample has a lower erosion rate than that of the basin-averaged rate derived from the Spanish Gulch sample, which presumably also includes clasts being eroded from steeper slopes below the summit ridge. The 26Al/ 10 Be ratio of the summit ridge sample does not overlap with the steady-erosion line in Fig. 3, which suggests that this sample experienced prolonged burial by snow or sediment and/or eroded by slab exfoliation. The conglomeratic bedrock on the narrow summit ridge crest is horizontally-bedded, and these beds dip moderately and often form planes along which the rock has fractured and weathered. If dislodged, the resulting blocks (10–40 cm thick) could easily tumble down the dip slope, exposing a “new” surface that had already accumulated a CRN inventory. This pattern of erosion can result in an anomalously low 26Al/10Be ratio for a surface sample (Gosse and Phillips, 2001) and is the most likely explanation for the unexpectedly low 26Al/10Be ratio of the summit ridge sample. The fact that the 26Al/ 10 Be ratio of the Spanish Gulch sample overlaps with the steadyerosion line suggests that the majority of the sediment washed into the cave system is not affected by prolonged burial or non-steady styles of erosion. Shielding due to snow is neglected in age and erosion rate calculations but warrants discussion. Cover by snow results in the attenuation of secondary cosmic-ray particles, reducing surface CRN production rates. While burial ages are essentially unaffected by the effects of snow shielding, failing to account for the reduced production rates will result in overestimates of inherited erosion rates. Snow data are available from a SNOTEL station in the South Colony drainage, immediately north of Marble Mountain at 3292 m asl (Fig. 1A), for the past 30 years (available from the Western Climate Data Center; http://www.wrcc.dri.edu). After adjusting CRN production rates for the monthly mean snow water equivalent at this SNOTEL station, the resulting erosion rates are 12–13% lower than those reported in Table 1. Snow cover on the east face of Marble Mountain undoubtedly persisted longer into summers during the LGM, but increases in precipitation were likely not substantial (Leonard, 2007; Refsnider et al., 2009; Brugger, 2010), and there is no evidence that the range crest was glaciated in this section of the SdC (Refsnider et al., 2009). Prior to the LGM, there are no constraints on alpine snow cover in the region. Because of these uncertainties, reported erosion rates are not adjusted for snow shielding but likely overestimate actual rates by at least 10%. The magnitude of this overestimate probably differs for all samples, but the corrections and associated uncertainties should not affect the interpretation of the results discussed below. 5. Discussion 5.1. Water table lowering and paleo-erosion rates
Fig. 3. Two-isotope diagram showing cave samples (open circles) and surface samples (filled circles); error bars show 1σ standard error based on analytical uncertainties. The solid black curve and associated erosion rates track the CRN inventory derived from continuous surface exposure at a constant erosion rate. Upon burial, samples will track downward parallel to the dotted lines, and the dashed burial isochrons indicate the 26 Al/10Be ratio after a period of continuous burial.
Detrital sedimentary deposits in caves have previously been used to infer changes in fluvial erosion rates in systems where the depth of shallow phreatic cave passage formation is controlled by water table lowering due to stream incision (Granger et al., 1997, 2001; Stock et al., 2004; Häuselmann et al., 2007). The elevation of the vadose– phreatic transition in Marble Mountain is not directly controlled by the elevation of a single local stream. While the boundaries of this watershed have undoubtedly evolved over the past 5 Myr, its areal extent must have been restricted to between the caves and the summit ridge. The Marble Mountain cave system is therefore unique in that few geologic archives allow inherited CRN erosion rates to be interpreted as basin-averaged erosion rates, essentially analogous to the Spanish Gulch surface sediment sample. The morphology of the Spanish Cave system, which can be divided into three sections vertically based on passage development, provides support for the cave passage age estimates and the paleo-erosion rates
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Fig. 4. Mineralogy of the b2 mm size fraction of cave sediment samples and bulk surface bedrock and clast samples. All analyses have degrees of fit ≤0.12. The mineralogical segmentation of each bar matches the same order as the minerals listed in the legend. These data suggest that there have been no dramatic changes in the lithology of material being eroded and washed into the cave system.
derived from the detrital sediment. The uppermost section above Sand Crawl (0 to 95 m depth below the upper Spanish Cave entrance), is predominantly vadose passages with only a few phreatic tubes. The section below Sand Crawl (95 to 135 m depth), characterized by incised phreatic tubes connecting small rooms, is where the majority of passage development took place in both Spanish Cave and White Marble Halls. T-Slot Crawl (114 m depth) is in the middle of this section. The main entrance passages to these caves are the lowest extensive phreatic tubes in this section. In the lower section of Spanish Cave, only one additional phreatic tube level has been found, which is 90 m lower and at the same elevation as Burns Cave (206 m depth). Spanish Spring is approximately 40 m below this lowest level. The extensive vadose enlargement and development of the upper section of the cave system suggests considerable antiquity, in general agreement with the ca. 5 Ma age from Sand Crawl. The 1.1 ± 0.1 Ma age from only 19 m lower in the cave system at T-Slot Crawl suggests that the water table remained at a similar elevation for several million years, allowing considerable time for the extensive network of phreatic passages in the middle section to develop. Subsequently, the water table dropped rapidly to the level of Burns Cave as indicated by the overlapping ages from T-Slot Crawl and Burns Cave. This lowering of the vadose–phreatic transition was rapid enough that virtually no phreatic passage development occurred. The pattern of water table lowering recorded by both the cave morphology and sediment burial ages is related to springs controlling the local base level (Klimchouk et al., 2000). The elevation of Spanish Spring (Fig. 1B) and springs at similar elevations near the valley floor on the west side of Marble Mountain presumably control the level of the modern water table. Phreatic tubes within the cave system thus record higher base levels within the watershed, and the rate of water table lowering must somehow be related to landscape lowering rates. The prolonged period of relative water table stability between Sand Crawl and T-Slot Crawl is in accordance with the low surface erosion rates during between ca. 5 and 1.2 Ma. Rapid lowering of the water table from the level of T-Slot Crawl to Burns Cave corroborates the subsequent increase in surface erosion rates, and the level of the spring controlling the base level is now approximately 40 m lower than Burns Cave. These morphological relationships and ages suggest that the network of phreatic passages in the middle section of the cave system developed between ca. 5 and 1.2 Ma, and the dramatic increase in erosion rates most likely occurred shortly before 1.2 Ma. Modern and Pleistocene basin-averaged and summit ridge erosion rates from Marble Mountain are similar to the CRN-derived erosion rates (all reflecting ca. 10-ka time scales) measured from other ranges in the Rocky Mountains comprised predominantly of crystalline
bedrock. Bedrock ridges in the Teton Range of Wyoming are eroding at ~22 m Myr−1 (Tranel et al., 2009), and basin-averaged erosion rates in unglaciated catchments in the Front, Medicine Bow, and Laramie ranges of northern Colorado and southern Wyoming are 18– 30 m Myr−1 (Dethier et al., 2002). Basin-averaged erosion rates from non-glacial mountain catchments in Idaho range from 20 to 100 m Myr−1 (Kirchner et al., 2001). Basin-averaged erosion rates from steep, unglaciated catchments in the core of the Wasatch Range of Utah range from 170 to 790 m Myr−1, and these higher and more variable rates are likely due to glacial erosion in main canyons steepening the profile of the non-glacial tributary drainages from which samples were collected (Stock et al., 2009). Interestingly, the low Pliocene erosion rate recorded by the oldest Spanish Cave sample is more similar to bare-bedrock erosion rates (generally b10 m Myr−1) on Rocky Mountain summit flats (Small et al., 1997) and on ridges, hilltops, and summits across the eastern United States (Bierman et al., 1995; Granger et al., 2001; Hancock and Kirwan, 2007) than to modern basinaveraged erosion rates in the Rocky Mountains (Dethier et al., 2002; Stock et al., 2009). Estimates of mean fluvial incision rates in the mountains of central Colorado over the past 0.64 Myr are in the range of 100 to 150 m Myr−1 (Dethier, 2001), considerably higher than the Marble Mountain erosion rates. The result of this contrast is an increase in landscape relief, which both supports the conclusion of Small and Anderson (1998) that modern relief in the Rocky Mountains is currently increasing and suggests that such an increase has been underway for at least 0.6 Myr if alpine erosion rates have been relatively constant over this interval. 5.2. The erosion rate increase The late Cenozoic increase in erosion rates of nearly an order of magnitude for Marble Mountain can be explained by a combination of four potential factors: (1) a change in the bedrock lithology being eroded, (2) increases in local glacial and fluvial incision rates, (3) an increase in surface uplift rates due to local tectonism, or (4) a change in weathering processes. The first two factors are likely of little importance in this setting. Strata dip moderately to the west, so the eastern slope of Marble Mountain exposes a sequence of sandstone and conglomerate units. While it is possible that more easily erodible beds may have previously been exposed, the b2 mm size fractions of each CRN sample have similar mineralogical compositions (Fig. 4). This implies that there have been no dramatic changes in the composition of bedrock being eroded on the upper slopes of Marble Mountain. While
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there is evidence for mid to late Pleistocene increases in fluvial incision rates of major rivers in the Southern Rocky Mountains immediately to the east (Dethier, 2001 and references therein), Marble Mountain is situated above Promontory Divide in the Wet Mountain Valley (Fig. 1A), and waves of incision advancing south from the Arkansas River drainage and north from the Huerfano River drainage have not yet begun to dissect the Plio–Pleistocene valley fill near this divide. Glacial scouring in valleys to the north and west of Marble Mountain should not have had a direct effect on the unglaciated east-facing slopes above the Spanish Cave system. The west edge of the SdC Range follows the SdC Fault (Fig. 1A), which locally forms the eastern boundary of the Rio Grande Rift. On the basis of faulted alluvial fan surfaces from at least three glaciations, McCalpin (1983, 1986) estimated mean vertical slip rates of 0.1 to 0.2 mm yr−1 (50 to 100 m Myr−1) for the SdC Fault. Seismic stratigraphy and well log correlations show units within the Santa Fe Group (representing the past 20 Myr) consistently thickening toward the eastern margin of the San Luis Valley (Brister and Gries, 1994). Alamosa Formation sediments (3.0 to 0.4 Ma) also thicken to the east (Machette et al., 2007), implying a relatively continuous increase in accommodation space in the basin during the late Cenozoic. There is also evidence for a shift in the focus of tectonic activity farther south along the fault based on geomorphic features best explained by higher slip rates along the northern end of the SdC Range prior to the late Pleistocene (Ruleman and Machette, 2007). Thus, long-term mean slip rates along the SdC Fault immediately west of Marble Mountain have likely not increased substantially and may even have decreased since the mid Pleistocene. A shift in weathering regimes related to climatic change since the mid Pliocene provides the most reasonable mechanism to explain the increase in erosion rates. Pliocene climate in the Southern Rocky Mountains was characterized by more humid conditions and lower seasonality than in the Pleistocene (Thompson, 1991). Rogers et al. (1992) reconstructed San Luis Valley climatic conditions between 2.6 and 0.7 Ma and identified a gradual transition to cooler temperatures (ca. 3 °C above modern) and increasing aridity beginning at ~ 2.5 Ma and continuing until 1.6 Ma. Conditions 3 to 5 °C warmer and more arid than modern persisted between 1.6 and ~ 0.93 Ma and were followed by the onset of high-amplitude glacial–interglacial oscillations in both temperature and moisture. Evidence of increased runoff from upland sources beginning at ~ 1 Ma likely indicates the expansion of glaciers and seasonal snowpack in the SdC Range and San Juan Mountains, and a 3- to 6-fold increase in sedimentation rates in the San Luis Valley between 2 and 1 Ma may be related to the onset of extensive glaciation in these mountains (Rogers et al., 1992). The apparent rapid water table lowering and vadose incision within the Spanish Cave system indicated by the overlapping ages from T-Slot Crawl and Burns Cave at ca. 1.2 Ma is best explained by the intensification of glaciation and valley floor lowering on the west (where springs are located near the modern valley floor) and north sides of Marble Mountain (e.g., Häuselmann et al., 2007). Cooler temperatures during the mid to late Pleistocene would result in bedrock at high elevations in the SdC Range experiencing more time within the frost-cracking window (−3 to −8 °C; Anderson, 1998), and increased moisture from snowmelt could accelerate periglacial weathering rates and the downslope transport of eroded material by frost creep, gelifluction, and fluvial action. Erosion rates in at least some small alpine catchments in the Alps increase with elevation, which Delunel et al. (2010) interpret as being dominantly due to frost-cracking processes. The mean modern winter (October to April) temperature for the Marble Mountain catchment is −2.7 °C (cf. Refsnider et al., 2009), near the upper limit of the frost-cracking window. Temperatures of at least 3 °C warmer than present would substantially reduce the duration during which bedrock within this catchment would be within the frost-cracking window. Additionally, landscapes may also be unable to reach a state of equilibrium with
climate in the presence of millennial-scale climate variability, resulting in continuous adjustments in erosion rates (Zhang et al., 2001; Molnar, 2004; Schaller et al., 2004). 6. Conclusions The Marble Mountain record of paleo-erosion rates provides the first direct, CRN-based evidence for dramatic erosion rate increases in the Rocky Mountains between 4.9 and 1.2 Ma, and based on passage development and morphology, this increase likely occurred shortly before 1.2 Ma. This record also suggests that modern alpine erosion rates are similar to rates at ca. 1.2 Ma. Because this record comes from a locality relatively insulated from the direct influence of glacial and major fluvial systems, the transition from relatively warm Pliocene conditions into the cooler Pleistocene climate with an associated increase in the effectiveness of periglacial weathering processes at high elevations is the most likely explanation for the observed increase in erosion rates. Acknowledgements This project was made possible with funding from the National Speleological Society. Dylan Rood at LLNL-CAMS provided Be and Al isotope measurements. Skip Withrow, Jim Wilson, and Margaret Barnes were instrumental in sample collection, Skip Withrow shared indispensible cave survey data, Donald Davis and Fred Luiszer provided stimulating speleogenesis discussions, Gifford Miller, Miriam Dühnforth, and Keith Brugger kindly provided comments on the manuscript, and reviews by Darryl Granger and an anonymous referee included insightful and helpful critiques. References Anderson, R.S., 1998. Near-surface thermal profiles in alpine bedrock: implications for the frost weathering of rock. Arc. Alp. Res. 30, 362–372. Balco, G., Stone, J.O.H., 2005. Measuring middle Pleistocene erosion rates with cosmicray-produced nuclides in buried alluvial sediment, Fisher Valley, southeastern Utah. Earth Planet. Sci. Lett. 30, 1051–1067. Balco, G., Stone, J.O., Lifton, N.A., Dunai, T.J., 2008. A complete and easily accessible means of calculating surface exposure ages or erosion rates from 10Be and 26Al measurements. Quat. Geochronol. 3, 174–195. Bierman, P.R., Gillespie, A.R., Caffee, M.W., Elmore, D., 1995. Estimating erosion rates and exposure ages with 36Cl produced by neutron activation. Geochim. Cosmochim. Acta 59, 3779–3798. Brister, B.S., Gries, R.R., 1994. Tertiary stratigraphy and tectonic development of the Alamosa Basin (northern San Luis Valley), Rio Grande rift, south-central Colorado. In: Keller, G.R., Cather, S.M. (Eds.), Basins of the Rio Grande Rift: Structure, Startigraphy, and Tectonic Setting: Geological Society of America Special Paper, vol. 291, pp. 39–58. Brugger, K.A., 2010. Climate in the southern Sawatch Range and Elk Mountains, Colorado, USA, during the Last Glacial Maximum: inferences using a simple degreeday model. Arc. Ant. Alp. Res. 42, 164–178. Davis, D.G., 1960. The Caves of Marble Mountain, Colorado, Described and Located. S. Col. Grotto Spec. Bull., vol. 2. Delunel, R., van der Beek, P.A., Carcaillet, J., Bourlès, D.L., Valla, P.G., 2010. Frost-cracking control on catchment denudation rates: Insights from in situ produced 10Be concentrations in stream sediments (Ecrins–Pelvoux massif, French Western Alps). Earth Planet. Sci. Lett. 293, 72–83. Desilets, D., Zreda, M., Prabu, T., 2006. Extended scaling factors for in situ cosmogenic nuclides: new measurements at low latitude. Earth Planet. Sci. Lett. 246, 265–276. Dethier, D.P., 2001. Pleistocene incision rates in the western United States calibrated using Lava Creek B tepha. Geology 29, 783–786. Dethier, D.P., Ouimet, W., Bierman, P.R., Finkel, R.C., 2002. Long-term erosion rates derived from 10Be in sediment from small catchments, northern Front Range and southern Wyoming. Abstr. Programs - Geol. Soc. Am. 34, 409. Eberl, D.D., 2003. User guide to RockJock: a program for determining quantitative mineralogy from X-ray diffraction data. U.S. Geological Survey, Open File Report 03-78. 40 pp. Gosse, J.C., Phillips, F.M., 2001. Terrestrial in situ cosmogenic nuclides: theory and application. Quat. Sci. Rev. 20, 1475–1560. Granger, D.E., Kirchner, J.W., Finkel, R.C., 1997. Quaternary down-cutting rate of the New River, Virginia, measured from differential decay of cosmogenic 26Al and 10Be in cave-deposited alluvium. Geology 25, 107–110. Granger, D.E., Fabel, D., Palmer, A.N., 2001. Pliocene–Pleistocene incision of the Green River, Kentucky, determined from radioactive decay of cosmogenic 26Al and 10Be in Mammoth Cave sediments. Geol. Soc. Amer. Bull. 113, 825–836.
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