Slow denudation within an active orogen: Ladakh Range, Northern India
A thesis submitted to the Graduate School of the University of Cincinnati
in partial fulfillment of the requirements for the degree of
Master of Science
in the Department of Geology of the College of Arts and Sciences
2011
by
Scott A. Reynhout B.S., Beloit College, 2008
Committee Members: Craig Dietsch, Ph.D. (chair) Lewis A. Owen, Ph.D. Marc W. Caffee, Ph.D.
i ABSTRACT
Cosmogenic 10Be measurements of both bedrock and basin denudation reveal strong topographic-climatic control on erosion within the Ladakh Range of northwestern India. Bedrock weathering rates, extremely slow below the contemporary equilibrium line altitude, increase by
an order of magnitude at the high-elevation range divide. Along the southern slope of the Range
south of the range divide, rates of ridge crest summit lowering vary from 0.02±0.03 m/Myr to
0.7±0.09 m/Myr, whereas rates of range divide summit lowering vary from 4.97±0.45 m/Myr to
13.13±1.17 m/Myr. Long-term denudation rates of nonglaciated basins are the lowest yet
reported. Accelerated glacial and periglacial erosion near the range divide drives a “frost
buzzsaw” in this region that works to selectively destroy high topography. The actively-eroding
glacial landscape stands in stark contrast to hillslopes at lower elevations, which have achieved
long-term, near equilibrium conditions and may represent extant Pliocene or older landscapes.
ii
iii ACKNOWLEDGEMENTS
SAR thanks the Geological Society of America for funding this research through the 2009 Arthur
D. Howard Award, and the University of Cincinnati Department of Geology for supporting this study as part of his M.S. degree research. T. Dorje of Discover Ladakh Adventure provided field support and Susan Ma of PRIME Lab provided technical support.
iv Table of Contents
Abstract ...... ii
Acknowledgements ...... iv
Introduction ...... 4
Regional setting ...... 5
Rationale and field methods ...... 8
Laboratory methods and post processing ...... 10
Results ...... 11
Discussion...... 13
Conclusion ...... 21
Figures and tables ...... 22
References ...... 26
1 List of Figures and Tables
Figure 1. Regional overview ...... 22
Figure 2. Denudation vs. elevation ...... 23
Figure 3. Denudation vs. basin metrics ...... 24
Table 1. Short denudation rate table ...... 25
2 List of Appendices
Appendix A. Full sample data table ...... 34
3 INTRODUCTION
The lower elevation limit of glaciation in mountains broadly defines a boundary above which rates and processes of erosion differ significantly from those controlled dominantly by fluvial and mass movement processes at lower altitudes. As efficient alpine erosive mechanisms (Hallet
1996), even small temperate glaciers can downcut rapidly (e.g. Loso et al., 2004; Riihimaki et
al., 2005), in some cases outpacing fluvial downcutting (Brocklehurst and Whipple, 2002). The
mean Quaternary equilibrium-line altitude (ELA) is often used as a benchmark for glaciation. In
mountain ranges around the world, ELAs correlate well to zones of focused erosion (Ballantyne,
2002; Berger and Spotila, 2008); to summit elevations in the Andes (Montgomery et al., 2001),
the eastern Himalaya (Brozovic et al., 1997) and Washington Cascades (Mitchell and
Montgomery, 2006); to cirque floor elevations in the Cascades (Mitchell and Montgomery,
2006) and Alps (Anders et al., 2010); and to peak hypsometric surface area worldwide (Egholm
et al., 2009), including the Andes (Montgomery et al., 2001) Sierra Nevada (US) (Brocklehurst and Whipple, 2004), Cascades (Mitchell and Montgomery, 2006), and the Lemhi Range and
Bitterroots (Foster et al., 2008). Collectively, the geomorphic significance of ELAs suggests the
existence of a “glacial buzzsaw” mechanism that acts to limit the vertical development of
topography at or above the glaciated zone, within ~ 1 km of the mean Quaternary ELA.
Here, a suite of terrestrial cosmogenic nuclide (TCN)-derived erosion rates are used to
quantitatively bracket the vertical transition from a transport limited, equilibrium landscape to an
actively-eroding glacial-periglacial landscape in the Ladakh Range of the Transhimalaya,
northern India, within the Himalayan-Tibetan orogen. This transition occurs in the vicinity of the
contemporary ELA, which suggests that glacial and periglacial processes account for the
4 majority of denudation in this, and possibility other, arid landscapes within the Himalayan-
Tibetan orogen. The order-of-magnitude difference between erosion in the inactive versus active landscape also implies that glaciation is reducing net topographic relief on the southern side of the Ladakh Range, albeit very slowly.
REGIONAL SETTING
The NW-striking Ladakh Range rises from the Indus and Shyok valleys at ~ 3000 m asl to >
6000 m asl, with a width perpendicular to strike of ≤ 50 km (Fig. 1). Bedrock within the range is
composed of the Ladakh batholith, the granodioritic roots of a pre-Himalayan island- and
continental-arc complex (Honegger et al., 1982), which is genetically related to the Kohistan arc
of northwest Pakistan and the Gangdese batholith of southern Tibet (Hodges, 2000). South of the
Range, the Indus Suture Zone (ISZ) separates the Ladakh batholith from the Indus molasse; to
the north, the Karakoram Fault (KF) in the east and the Shyok Suture Zone (SSZ) in the west
juxtapose the Karakoram terrane against the Ladakh batholith (Searle, 1986). Apatite fission
track (AFT) and (U-Th)/He apatite (AHe)and zircon (ZHe) ages from the Ladakh Range record a pulse of cooling at ~ 22 Ma (Kirstein et al., 2006) coincident with exhumation of the High
Himalaya. By late Miocene time, cooling of the Range had slowed considerably throughout the the batholith south of the range divide (Kirstein et al., 2009). Active faults in the Ladakh Range have not been recognized.
The central Ladakh Range (centered on Leh, 34°09'N, 77°34'E) exhibits pronounced
morphometric asymmetry. Basin size, valley width, and mean elevation increase north of the range divide. Jamieson et al. (2004) and Kirstein et al. (2006) attribute this asymmetry to the
5 northward propagation of the ISZ, which induced block tilting along the range‟s long axis.
Alternatively, Dortch et al. (in review) suggest that transpression along the KF produced higher
northerly elevations and subsidence of the southern edge of the range. Alluvial fans, small peaks
from buried spurs, aggradation of streams, and highly denuded mountain ridges and spurs all
suggest that south of its divide, the Ladakh Range has been tectonically inactive at least since the
Pleistocene.
The glacial history in the Ladakh Range is one of progressively smaller glacial advances after ~
430 ka (Owen et al., 2006). Burbank and Fort (1985) plotted the contemporary steady-state ELA
at 5280 m asl south of the range divide. Hobley et al. (2010) described the upper reaches of
glaciated catchments as possessing classical postglacial morphologies, with U-shaped valleys
and hummocky valley bottoms hosting significant paraglacial fans. Below the glaciated area,
tributary valleys are fluvial in character and lack evidence of glacial modification; Hobley et al.
(2010) consider these tributaries to be preglacial relict landscapes.
Regional climate is consistent with a cold desert, with patterns of precipitation and temperature
affected by the extreme altitudes, the Indian summer monsoon, mid-latitude westerly storm
tracks, and the rain shadow effect of the High Himalaya to the south and west. Contemporary
annual precipitation amounts to < 500 mm/yr across Ladakh (Bookhagen and Burbank, 2006), with a measured 50-year average of 92.6 mm/yr in Leh (Holmes, 1993). Precipitation increases
with altitude, although the magnitude of this increase has yet to be quantified (Derbyshire et al.,
1991; Hewitt, 1993; Owen et al., 2006). In Leh, January mean temperature maxima and minima
are -2.8°C and -14.0°C, respectively; July mean temperature maxima and minima are 24.7°C and
6 10.2°C, respectively (Osmaston, 1994). The environmental lapse rate averages to ~1°C per 170 m (Derbyshire et al., 1991). Vertical zonation of flora across the study area supports observations
of progressively colder and wetter conditions with altitude. Flora characteristic of subalpine
desert transition to steppic and then to alpine flora at progressively higher elevations; plant cover increases from 10 % in the subalpine desert to 65 % in the alpine grassland (Hartmann, 2009).
Using sediment flux from the Indus River, Garzanti et al. (2005) calculated an average erosion
rate of 50±50 m/Myr for Indus catchments draining the Ladakh arc. From AFT, AHe, and ZHe
ages and thermal modeling, Kirstein et al. (2009) proposed a multi-phase cooling history for the
Ladakh batholith. They derived a near-surface geotherm of 40C/km and calculated that the
southern Ladakh Range near Leh cooled rapidly between about 32 and 25 Ma, and since 20 Ma
cooled very slowly yielding a time-integrated exhumation rate since 20 Ma of 25m/Myr. The
relatively low erosion rates implied in these exhumation rates reflect both the tectonic quiescence
and arid climate of the region. Observed in-situ rock weathering in the Ladakh Range scales
positively with grain size, as evidenced by knobby weathering favoring preservation of fine-
grained mafic xenoliths. Haloclasty and grusification are common, particularly at low elevations,
and the sparse regolith consists mostly of coarse quartz sand in a fine-grained matrix. Hartmann
(2009) noted widespread calcification throughout the greater Ladakh region, with basic soil pH
values persisting even in granitic regions.
The amount of regolith mantle, vegetation cover, and humic content all increase with altitude in
unglaciated catchments, likely reflecting progressively more humid conditions with altitude. The
changing characteristics and quantity of regolith with elevation suggest that hillslope erosion in
7 unglaciated tributaries is supply-limited at lower elevations and transport-limited at higher elevations (Hobley et al., 2010). Granular disintegration dominates physical weathering in all but the highest elevations, where abundant talus and felsenmeer suggest that the predominant mode
of physical weathering is block removal. Below the highest elevations, frost heaved blocks, soil,
and small-scale patterned ground are evident, suggesting that freeze-thaw processes are
particularly active at higher elevations.
RATIONALE AND FIELD METHODS
TCN measurements of bedrock or catchment-wide denudation account for the sum of active
erosive and weathering processes at a point or given area (Gosse and Phillips, 2001) and have been extensively used to document relief production (Small et al., 1997; Hancock and Kirwan,
2007; Quigley et al., 2007), identify relict landscapes (Bierman and Caffee, 2002; Stroeven et al.,
2002), and quantify erosion in small basins (Bierman and Steig, 1996; Granger et al., 1996).
Denudation of summits and adjoining basins must be determined in concert to evaluate
topographic relief production. Physical erosion is dominated by mass wasting, and fluvial and
glacial processes whereas physical weathering involves various thermal stresses, including frost
weathering and thermal shock/fatigue, unloading and exfoliation, and salt-crystal weathering.
Chemical weathering consists of the chemical breakdown of rock by silicate hydrolysis,
hydration, dissolution, oxidation, and/or biological activity. Bedrock summit tors are eroded only
by weathering, deep-seated bedrock landslides, and hillslope retreat. Therefore, in the absence of
mass wasting, weathering is the only force that lowers summits and so denudation of tors is
“weathering limited” (Small et al., 1997; Kober et al., 2007). TCN measurements of summit
bedrock lowering thus measure the sum of all weathering processes actively lowering the summit
8 in question, and set the baseline weathering rate for a given area. The difference between summit weathering and basin-wide denudation is net relief production.
Twelve summit-basin pairs were sampled to determine 10Be TCN concentrations for erosion
determinations across the southern flank of the Ladakh Range (Fig. 1; CH, LH, PH, and TS
sample sets). Each pair comprised one bedrock sample, taken from a tor on a localized summit, and one sediment sample, taken from a basin draining the summit. The close proximity of summit and basin sampling sites ensures that summit samples provide an accurate minimum denudation rate for the whole basin, allowing for a quantitative assessment of relief production.
Bedrock was sampled from the tops of tors and from sharp narrow bedrock ridges on summits or interfluves. Poorly-sorted regolith mantles many of the tors, which exhibit laterally conformal
joint planes that extend into the subsurface. Samples were preferentially collected from bedrock prominences > 1 m high to help minimize any possible error from past burial. Coarse-grained granite comprised the majority of samples, although several summits were composed of fine- to
medium-grained granodiorite. Samples were located using a handheld GPS and photographed.
At each site, ~ 500 g of bedrock was collected to a depth of 1-5 cm using a hammer and chisel.
Topographic shielding was minimal (less than 20°) at each bedrock location.
Erosion of alpine landscapes reflects interplay among lithology, tectonics, and climate.
Measurement of in situ-produced TCNs has been reliably used to assess long-term erosion in
small catchments of uniform geology (Bierman and Steig, 1996; Granger et al., 1996). TCN
concentrations in fluvial sediments are a function of the average erosion rate, cosmic ray
attenuation length, and the TCN production rate averaged throughout the basin (Lal, 1991). This
9 method assumes uniform distribution of target minerals, as well as proportional denudation across the basin (Granger et al., 1996). The former can be reasonably inferred in regions of near- homogeneous lithology; we control for the latter by choosing small basins without evidence of
former glaciation or significant mass wasting, all of which affect the distribution of sediment and
may have an adverse effect on TCN-determined rates of basin denudation (Niemi et al., 2005;
Wittmann et al., 2007).
Sediment was sampled from small (< 5 km2) basins adjacent to the sampled summits.
Approximately 1 kg of quartz-rich sand was collected from no fewer than five points along a 20-
m stretch of the active channel and dried in the field.
LABORATORY METHODS AND POST PROCESSING
Quartz isolation, chemical separation of BeO, and cathode preparation were conducted in the
Department of Geology at the University of Cincinnati using the methods of Kohl and
Nishiizumi (1992), and detailed in Dortch et al. (2009). The 10Be/9Be ratios for WL- and NL-
sample series were measured at the Lawrence Livermore National Laboratory‟s Center for
Accelerator Mass Spectrometry (CAMS) and for the CR, LH, PH, and TS samples at the Purdue
Rare Isotope Measurement (PRIME) Laboratory.
Erosion rate calculations for the summit bedrock samples utilized the CRONUS 10Be-26Al
erosion rate calculator (v. 2.2; Balco, 2009) with the scaling factors of Lal (1991) and Stone
(2000). Basin erosion calculations follow the procedure outlined in Granger et al. (1996).
Production rates for the basins upstream of the sampling locations were determined after the
10 method outlined in Dortch et al. (in review). Although selective dissolution and quartz enrichment may take place in this environment, the amount is likely to be small, on the order of
~ 10% (Riebe et al., 2001a).
Catchments were delineated on a 3 arc second (~ 90 m) Shuttle Radar Tomography Mission
digital elevation model (CGIAR-CSI, 2008) using MATLAB, following the methods of Wobus
et al. (2006). Neutron, fast muon, and slow muon production rates within the basin were
calculated in MATLAB on a pixel-by-pixel basis using the scaling factors of Stone (2000), a revised sea-level high-latitude production rate of 4.49±0.39 10Be atoms gram (quartz)-1 year-1 and a 10Be half life of 1.36 Ma (Nishiizumi et al., 2007). Shielding corrections for each pixel were applied by binning surrounding pixels into azimuth increments of 30° and using the maximum
angle to the horizon to estimate shielding (Dortch et al., 2011).
RESULTS
Collectively, across the southern flank of the range and below the contemporary ELA, rates of
basin denudation reveal slow erosion rates of < 2 m/Myr. Rates of ridge crest summit lowering
are even slower, ranging from 0.02±0.03 m/Myr to 0.7±0.09 m/Myr (Fig. 2). Two ridge crest
summit samples (PH45P, TS40P) have reached secular equilibrium between 10Be production and
decay, indicating a theoretical condition of zero erosion. Rates of range divide summit lowering
are an order of magnitude higher and vary from 4.97±0.45 m/Myr to 13.13±1.17 m/Myr.
Summit bedrock samples cluster into two distinct populations, with samples taken from the range
divide exhibiting higher weathering rates than samples taken from ridge crests. Plotted by
11 elevation, summit lowering—equivalent in this case to rock weathering—does not exceed ~ 1
m/Myr below an elevation of 5100 m asl (Fig. 2). Above 5600 m, weathering rates increase by
an order of magnitude.
Rates of basin denudation were compared to specific basin metrics to assess systematic
relationships between basin morphometry and denudation (Fig. 3; e.g. Ahnert, 1970; Stock et al.,
2009; Delunel et al., 2010). No correlation is observed between basin denudation and area, relief, or mean slope; basin denudation exhibits a weak positive correlation with mean basin elevation.
These results may be due to the basins being too similar in size to see a notable change in basin
denudation and basin statistics.
Although the study presented here focuses on small-scale basin denudation, large trunk valleys
sampled for basin denudation by Dortch et al. (in review) also span the study area (Fig. 1).
Compared to the basins sampled in this current study, the larger (101-102 km2) catchments
examined by Dortch et al. (in press) feature perennial streams that incise into glacial valley fill
(Hobley et al., 2010). These streams drain high reaches of the range which host active glaciers or have in the past. Erosion rates across these large glaciated catchments are between 10x and 20x faster than rates from the small catchments in this current study, a result that corresponds well to
global samples bearing on glacial vs. nonglacial denudation (Hallet et al., 1996; Norton et al.,
2010).
All basins from this study denude faster than their paired summits, implying a small amount of
relief production, 0.60 to 1.79 m/Myr. This suggests that hillslopes below 5100 m in the Ladakh
12 Range have remained remarkably static throughout the climatic fluctuations of the Quaternary, and may represent extant landscapes from the Pliocene or earlier. Rates of erosion from sources higher than ~ 5100 m asl are much different, with summit lowering rates characteristic of the
western U.S. (Small et al., 1997) and basin erosion rates comparable to granitic alpine basins in
the Sierra Nevada (Riebe et al., 2001b).
DISCUSSION
These new data help reveal the spatial scales at which a mountain landscape is modified by different weathering and erosive processes, and how a threshold elevation separates a high,
actively eroding landscape from a lower one in which where there is little change. Denudation
rates of unglaciated bedrock and catchments across the Ladakh Range south of its range divide
reveal a landscape near equilibrium. Rates of bedrock weathering adjacent to glacially-altered catchments, as well as basin denudation of these catchments, are up to an order of magnitude
higher than their nonglacial counterparts. It is likely that glaciation and the onset of periglacial conditions as the principal agents responsible for the tenfold increase in basin denudation and bedrock weathering in high elevation reaches of the Ladakh Range.
The range of summit lowering rates from the range divide is typical of summit lowering rates recorded from granitic alpine regions across the western United States (Small et al., 1997), the
Tibetan Plateau (Lal et al., 2003), the Cairngorm Mountains of Scotland (Phillips et al., 2006), and from temperate bedrock in the eastern United States (Hancock and Kirwan, 2007). Ridge
crest rates determined in this study, 0.02±0.03 m/Myr to 0.7±0.09 m/Myr, are more similar to the
very low rates of bedrock erosion recorded in Antarctica (Nishiizumi et al., 1991; Fogwill et al.,
13 2004), the Namib Desert (Bierman and Caffee, 2001; Van der Wateren and Dunai, 2001), the
Northern Territories of Australia (Bierman and Caffee, 2002), the Atacama Desert (Nishiizumi et al., 2005; Kober et al., 2007), and the Negev Desert (Matmon et al., 2009). The rates calculated
for this study correspond well to the prediction of Burbank and Fort (1985) of very slow erosion at low elevations within the Ladakh Range.
Our calculated maximum long-term basin denudation rates vary from 0.83±0.12 m/Myr to
2.01±0.27 m/Myr. These low rates are unprecedented in the corpus of TCN erosion literature and
are comparable only to conventional sediment yields from streams in Idaho (Kirchner et al.,
2001) and southwestern Australia (Tomkins et al., 2007). These rates are much lower than the
long-term denudation rates of 20-50 m/Myr for the southern Ladakh Range estimated by
Garzanti et al. (2005) and Kirstein et al. (2009), suggesting non-uniform erosion across the range. TCN-derived erosion rates elsewhere in the Himalaya range from 1200±100 m/Myr in the
southern Tibetan Plateau to 2700 m/Myr in the High Himalaya (Vance et al., 2003).
Calculated rates of denudation from TCN data average erosion over the time necessary to erode
the attenuation depth of basin sediment (attenuation depth / erosion rate), suggesting that slow erosion has been sustained since at least 0.3 Ma (Lal et al., 1991). Low erosion rates (> 0.5-1.5 m/Myr) imply long response times to episodic erosion events, wherein rates of basin denudation do not equilibrate to temporal perturbations in erosion rates but rather remain true to the long- term average (Bierman and Steig, 1996). Basin response time to a perturbation in erosion rate without elevation fluctuations is equivalent to the time necessary to erode 2 to 3 spallation
production e-folding lengths (Parker and Perg, in press). Assuming a conservative e-folding
14 length of 0.60 m x 2, basin denudation rates must have been constant since at least 0.6 Ma, while bedrock weathering rates have been constant since no later than 1.7 Ma. We are thus confident that our rates accurately reflect long-term erosion averaged across the sampled basins.
The calculated rates of very slow basin denudation are in accord with post-late Miocene
denudation rates of the southern, low elevation margin of the Ladakh Range inferred from
Kirstein et al.‟s (2009) modeled cooling history of low-temperature thermochronometers. Rapid
denudation between about 32 and 27 Ma gave way to much slower cooling at ~ 20 Ma, and
sustained, near-zero denudation at ~ 10 Ma when the rocks must have essentially reached the
surface. Coupled with limited sediment drape on hillslopes, near-zero denudation implies
extremely slow regolith generation over a 107 year time scale. Slow bedrock weathering rates
explain the preservation of very old (≥ 430 ka) moraines in the region (Owen et al., 2006), and suggest that slow denudation has persisted throughout the Quaternary, potentially to the Pliocene or earlier.
Basin metrics show no correlation with denudation (Fig. 3), in contrast to other workers who have found strong correlations between denudation and either relief (e.g. Ahnert, 1970;
Summerfield and Hulton, 1994) or mean slope (Montgomery and Brandon, 2002; Roering et al.,
2007). Most of these studies focus on perennial fluvial denudation of large basins, whereas
basins sampled here are small, with ephemeral flow. Hillslopes show evidence of scree
development, small-scale mass wasting, and frost heave, suggesting that downhill creep may be
responsible for sediment movement within the small basins we sampled (Roering et al., 2001).
15 The results suggest that nonglacial hillslope channels host inefficient streams, lacking the ability
to incise their channels or affect hillslopes through channel-hillslope couplings (Burbank et al,
1996). Sampled basins in this study occupy nonglaciated spaces characteristic of Domains II and
III of Hobley et al. (2010). Particularly in their dominantly aggradational Domain III, base level
fluctuations should not have an effect on these streams. In Domain II, where localized incision
takes place, our results suggest that any perturbations of base level have limited effects on long-
term denudation of these hillslopes.
Frost weathering control of basin denudation was inferred by Delunel et al. (2010), utilizing
mean basin elevation as a proxy for frost weathering intensity. The positive trend between basin
denudation and mean basin elevation in our data (Fig. 3B), in the absence of any other
relationships to basin metrics, could represent periglacial influence on small catchments.
Alternatively, increasing vegetation and frost-heaved regolith mantle encountered with elevation
could mark the indirect impact of orographic precipitation on erosion rates (Nearing et al., 2005;
Roering et al., 2010).
Rates of denudation increase significantly in samples from high altitude. We note that the ELAs
of 382 contemporary glaciers across the Ladakh Range are remarkably consistent, with an
average elevation of 5455±130 m (Dortch et al., in review). Above ~ 5100 m asl, talus-mantled
hillslopes are common, while below, hillslope mantles consist of coarse, poorly-sorted sediment.
Talus and block fields, as products of macro gelivation, are characteristic features of periglacial
landscapes; the absence of talus fields below 5100 m suggests that frost weathering declines in
relative importance to granular disintegration. Bedrock above 5600 m asl weathers at rates
16 characteristic of global alpine regions. This distinct change in weathering character, hillslope composition, and absolute weathering rates implies that significant frost-cracking is responsible
for the order-of-magnitude increase of weathering rates at high altitude.
Hales and Roering (2007) provided a model in which the growth of segregation ice within
substrate is most intense at an elevation within the range of elevation at which mean annual
temperature (MAT) is just above 0°C. Using a MAT of 4.9°C at Leh and a regional
environmental lapse rate of 0.6°C/100 m (Derbyshire et al., 1991; Schäfer et al., 2008), the zone
of maximum segregation ice growth in the Ladakh Range south of the range divide falls slightly
above ~ 4300 m asl, below where significant development of talus-mantled hillslopes occurs. We
suggest a nonlinear effect of climate on weathering, wherein frost cracking mechanisms are
limited by the arid conditions at low elevations. Sufficient precipitation for widespread
development of segregation ice occurs only at high elevations where there is greater
precipitation.
While the rates of weathering bedrock in the Ladakh Range by frost-cracking are not exceptional
in a global context (André, 1997; Small et al., 1997; Heimsath and McGlynn, 2008), they are
sufficient in this setting to operate a „frost buzzsaw‟ (Hales and Roering, 2009), in which
summits above the weathering threshold lower ten times faster than those below the threshold,
reducing net relief. Whipple et al. (1999) highlighted frost-cracking as a process that could limit
relief production from isostatic peak uplift.
17 The denudation rates from Dortch et al. (in review) incorporate catchments above the contemporary ELA, and show a similar order-of-magnitude increase over their lower elevation
counterparts. TCN-derived rates of basin-averaged erosion measure erosion averaged over the
entire catchment area; however, the data shows that the contribution from hillslopes below 5100
m asl comprises no more than ~ 10% of the measured erosion. Faster apparent denudation rates
must thus reflect: 1) increased contribution of sediment from glaciated areas; 2) reworking of
glacial debris; or 3) lateral channel erosion. Some reworking of glacial degrees is likely, as
glacial deposits on the valley floor are currently being incised by perennial streams (Hobley et al., 2010). However, subglacial erosion and paraglacial processes common to the glaciated
reaches of catchments provide more potential sediment to the main channel (Hallet et al., 1996;
Wittmann et al., 2007; Hobley et al., 2010). These observations suggest that glaciation acts to reduce range-scale basin relief, an effect predicted by Tomkin and Braun (2002).
The results presented here and in Dortch et al. (in review) suggest that climate modulates erosion
indirectly, through glaciation and periglacial processes. While there is no relationship between
altitude (a proxy for precipitation) and weathering/denudation below 5100 m, high glaciated
basins and intensively frost-weathered bedrock denude at rates that are up to an order of
magnitude faster compared to their lower elevation counterparts. Glaciers further affect
denudation by providing abundant unconsolidated sediment and a year-round source of flow to
trunk streams, whereas nonglaciated basins are affected only by ephemeral streams lacking sufficient power to move much sediment from hillslopes.
18 Although paired basin-peak denudation rates appear to indicate a small amount of relief production, extensive aggradation at lower elevations suggests that alluviation may outpace adjacent basin denudation, resulting in net destruction of relief. Isolated bedrock knobs along the
Indus River attest to the extensive burial of the landscape. Inferring relief production from basin-
averaged denudation can be fraught with complications, as apparent relief production may not accurately reflect actual relief production. Hillslopes may denude faster than summits, indicating
relief production, but burial of the lower slopes may actually reduce relief. Intra-basin sediment
storage and remobilization (Dühnforth et al., 2008; Hobley et al., 2010) or aggradation of the
lower reaches can lead to over- or underestimation of produced relief.
Preferential, long-term erosion of glaciated areas in the Ladakh Range supports the assertion that glaciers supply the majority of the alluvial sediment found within each trunk valley (Hobley et al., 2010). This sediment is ultimately delivered to the Indus Valley, implying that southwards fan propagation from the Ladakh Range, as well as overall denudation of the upper Indus catchment, is significantly influenced by glaciation. The observed disparity between erosion of glacial and nonglacial areas provides a mechanism by which glaciers can erode themselves out of
existence, a phenomenon first suggested by MacGregor et al. (2000). Progressively less
extensive glaciations have been documented worldwide, including within the Ladakh Range
(Kaufman et al., 2004; Owen et al., 2006; Kaplan et al., 2009, Dortch et al., 2011). Glacial retreat
in the Ladakh Range may also be linked to progressively decreasing precipitation supply to the
Transhimalaya (Owen et al., 2006) and/or subsidence of the southern half of the range (Dortch et al., in review).
19 The summer of 2010 was marked by extreme monsoon precipitation events across much of northern Pakistan and northwestern India. Flash floods on the nights of August 4-6 caused
extensive destruction across the Ladakh Range. Debris flows near Phyang (PH- samples in this
study) were responsible for widespread damage and 14 deaths; up to 8 m of incision into alluvium was reported (H. Munack, pers. comm.). Slow hillslope denudation rates reported in
this study are at apparent odds with the events of August 2010, and imply either: 1) the
magnitude of observed flooding is exceptional over a 106-107 yr timescale; or 2) intense rainfall
events are, in fact, responsible for minimal sediment stripping from hillslopes in arid
environments. The latter conclusion is supported by recent work demonstrating frequent alluvial
fan resurfacing in hyper-arid environments (Haug et al., 2010). Moreover, extreme flood events likely occurred during periods of enhanced monsoons throughout the Holocene and between 35-
45 ka (Fang, 1991; Gasse et al., 1991; Shi et al. 2001; Bookhagen et al., 2005; Dortch et al.,
2009) and since enhanced monsoons follow precession as modeled by Prell and Kutzbach
(1987), numerous increased precipitation events likely occurred since the onset of the south
Asian monsoon c. 8 Ma (Zhisheng et al., 2001). These rates demand caution from researchers
utilizing sediment-budget approaches to evaluate hazards (e.g. Kirchner et al., 2001), as long-
term erosion rates may underestimate actual flood frequencies.
20 CONCLUSIONS
Nonglaciated catchments across the southern Ladakh Range reveal summit bedrock weathering rates ranging from 0.02±0.03 m/Myr to 13.13±1.17 m/Myr, and basin denudation rates ranging
from 0.83±0.12 m/Myr to 2.01±0.27 m/Myr. These rates help define an active landscape
dominated by glacial processes and weathering by frost-cracking, and inactive hillslopes at near
equilibrium. The boundary between these two domains appears to parallel the contemporary
ELA, suggesting the action of glacial- and frost-buzzsaw mechanisms that reduce high
topography. This effect appears to be sustained over 106-107 year timescales, thus providing
quantitative evidence of a long-term climatic limit on alpine topography.
21 FIGURES AND TABLES
Figure 1. Study area in the southern Ladakh Range near Leh. Gray lines mark interfluves and blue lines mark channels based on a 3 arc second SRTM DEM. Sampling sites for summit erosion are shown as red dots and green triangles; catchments used for basin- averaged erosion are demarcated by red polygons. Blue squares denote sampling sites for basin-averaged erosion from Dortch et al. (in review). The teal shaded area shows the approximate extent of Leh stage glaciation (130-200 ka) mapped by Owen et al. (2006); black polygons are locations of present-day glaciers.
22
Figure 2. Basin-averaged denudation rate vs. maximum elevation for summit erosion rates (red dots and green triangles), basin-averaged erosion from this study (open red circles), and basin-averaged erosion from Dortch et al. (in review) (open blue squares). Maximum elevation marks the absolute elevation of a summit sample or the highest elevation encountered within a measured basin. The contemporary mean ELA for the southern Ladakh Range of 5455±130 m from Dortch et al. (in review) is marked by the dashed blue line. The grey box shows basin denudation and bedrock weathering rates vs. maximum elevation with an expanded y-axis.
23
Figure 3. Basin-averaged denudation rates versus (A) catchment area, (B) mean elevation, (C) catchment relief, and (D) mean hillslope angle.
24 Table 1. Locations of 10Be TCN samples, concentrations, 10Be concentrations, and minimum erosion rates. Sample 10Be number Location Elevation concentration Erosion 5 Latitude Longitude (m) (10 atoms/g SiO2) Erosion rate (oN) (oW) (m/Ma) Summit samples WL-1 34.1054 77.8280 5650 12.228 ± 0.279 6.06 ± 0.56 WL-2 34.1053 77.8280 5649 14.681 ± 0.171 4.97 ± 0.45 WL-3 34.1052 77.8281 5652 10.035 ± 0.138 7.47 ± 0.67 NL-1 34.3608 77.3682 5634 5.810 ± 0.135 13.13 ± 1.17 NL-3 34.3608 77.3682 5635 8.493 ± 0.196 8.87 ± 0.80 CR40P 33.9672 77.7758 4098 118.452 ± 5.099 0.02 ± 0.03 CR45P 34.0107 77.7899 4566 57.775 ± 1.905 0.52 ± 0.08 CR50P 34.0314 77.7830 4876 67.145 ± 3.408 0.53 ± 0.08 LH40P 34.1782 77.6090 4053 50.148 ± 1.652 0.45 ± 0.07 LH45P 34.1953 77.6292 4492 87.703 ± 3.922 0.23 ± 0.05 LH50P 34.2400 77.6295 5078 117.697 ± 4.177 0.21 ± 0.05 PH40P 34.1954 77.4523 4293 59.328 ± 1.975 0.42 ± 0.07 PH45P 34.2067 77.4617 4474 153.671 ± 6.373 0 PH50P 34.2344 77.4783 4998 58.848 ± 2.309 0.70 ± 0.09 TS40P 34.0122 77.7378 3942 126.928 ± 5.754 0 TS45P 34.0236 77.2697 4499 88.229 ± 3.264 0.22 ± 0.05 TS50P 34.0314 77.7830 4876 88.602 ± 2.351 0.32 ± 0.06 Basin-wide samples CR40B 33.9647 77.7839 3698 24.495 ± 0.753 1.34 ± 0.18 CR45B 33.9914 77.8011 3835 25.076 ± 1.114 1.50 ± 0.21 CR50B 34.0322 77.7981 4240 25.478 ± 1.385 1.76 ± 0.25 LH40B 34.1707 77.6041 3690 16.898 ± 0.758 1.87 ± 0.26 LH45B 34.1870 77.6232 4020 36.080 ± 1.527 1.08 ± 0.15 LH50B 34.2386 77.6195 4590 25.973 ± 0.964 1.95 ± 0.27 PH40B 34.1836 77.4586 3573 22.178 ± 0.496 1.42 ± 0.19 PH45B 34.1876 77.4680 3618 25.004 ± 1.068 1.33 ± 0.18 PH50B 34.2264 77.4872 4372 35.784 ± 1.522 1.30 ± 0.18 TS40B 34.0106 77.7256 3413 33.790 ± 1.578 0.83 ± 0.12 TS45B 34.0328 77.7622 3879 18.786 ± 0.579 2.01 ± 0.27 TS50B 34.0361 77.7578 3812 27.722 ± 0.767 1.39 ± 0.19
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33 Table 2. Locations for 10Be TCN samples, sample sizes, topographic shielding factors, concentrations, and analytical results and ages. d 10 9 e,f Sample number Lithology Location Elevation Thickness Production rate Shielding Quartz Be carrier Be/ Be 10Be concentrationf,g Agef,h,i Erosion ratej External a b and location latitude longitude Spallation Muons factor error error error uncertainty o o ‐13 5 name N W mcmatoms/g/yr g gx 10 10 atoms/g SiO2 ka ka g/cm2/yr m/Myr m/Myr Summits WL‐1 Granite 34.1054 77.8280 5650 4 118.65 0.879 1.0000 20.1433 0.3550 73.33 1.67 122.28 2.79 110 10 0.0016 6.1 0.56 WL‐2 Granite 34.1053 77.8280 5649 4 118.60 0.879 1.0000 20.6297 0.3489 91.75 1.07 146.81 1.71 130 12 0.0013 4.9 0.45 WL‐3 Granite 34.1052 77.8281 5652 4 118.75 0.880 1.0000 21.2405 0.3508 64.22 0.88 100.35 1.38 88 7.9 0.002 7.5 0.67 NL‐1 Granite 34.3608 77.3682 5634 4 118.76 0.876 1.0000 20.3953 0.3515 35.63 0.83 58.10 1.35 51 4.6 0.0036 13 1.17 NL‐3 Granite 34.3608 77.3682 5635 4 118.81 0.877 1.0000 22.1261 0.3542 56.08 1.29 84.93 1.96 74 6.8 0.0024 8.9 0.8 CR40P Granite 33.9672 77.7758 4098 3 57.68 0.609 1.0000 21.5529 0.3571 789.10 33.97 1184.52 50.99 5700 3100 0.00001 0.02 0.03 CR45P Granite 34.0107 77.7899 4566 3 72.80 0.685 1.0000 13.7281 0.3550 246.60 8.13 577.75 19.05 910 110 0.00014 0.52 0.08 CR50P Granite 34.0314 77.7830 4876 2 84.28 0.745 1.0000 19.0127 0.3457 407.60 20.69 671.45 34.08 900 110 0.00014 0.53 0.08 LH40P Granite 34.1782 77.6090 4053 2 56.63 0.605 1.0000 6.8151 0.3552 106.20 3.50 501.48 16.52 1000 130 0.00012 0.45 0.07 LH45P Granite 34.1953 77.6292 4492 1 70.54 0.683 1.0000 13.1548 0.3509 362.90 16.23 877.03 39.22 1600 250 0.00006 0.23 0.05 LH50P Granite 34.2400 77.6295 5078 3 92.87 0.776 1.0000 13.5029 0.3509 499.90 17.74 1176.97 41.77 1700 260 0.00006 0.21 0.05 PH40P Granite 34.1954 77.4523 4293 1 64.02 0.650 1.0000 13.6031 0.3576 249.10 8.29 593.28 19.75 1100 130 0.00011 0.42 0.07 PH45P Granite 34.2067 77.4617 4474 3 70.00 0.670 1.0000 20.3172 0.3541 973.20 40.36 1536.71 63.73 00000 PH50P Granite 34.2344 77.4783 4998 3 89.62 0.761 1.0000 15.1033 0.3557 275.80 10.82 588.48 23.09 720 82 0.00019 0.7 0.09 TS40P Granite 34.0122 77.7378 3942 1 53.27 0.593 1.0000 15.4062 0.3521 613.00 27.79 1269.28 57.54 00000 TS45P Granite 34.0236 77.2697 4499 3 70.72 0.675 1.0000 21.4930 0.3532 592.60 21.92 882.29 32.64 1700 250 0.00006 0.22 0.05 TS50P Granite 34.0314 77.7830 4876 3 84.28 0.739 1.0000 10.7982 0.3552 297.30 7.89 886.02 23.51 1300 170 0.00009 0.32 0.06 Basins (this study) d 10 9 e,f 10 f,g f,h,i Sample number Lithology Location Elevation Thickness Production rate Shielding Quartz Be carrier Be/ Be Be concentration Age Erosion rate External c and location latitude longitude Net Error factor error error uncertainty o o ‐13 5 name N W mcmatoms/g/yr g gx 10 10 atoms/g SiO2 ka ka g/cm2/yr m/Myr m/Myr CR40B Sediment 33.9647 77.7839 3698 ‐ 54.74 7.25 * 20.8835 0.3560 158.60 4.88 244.95 7.53 ‐‐‐ 1.3 0.18 CR45B Sediment 33.9914 77.8011 3835 ‐ 62.52 8.28 * 22.1386 0.3454 177.40 7.88 250.76 11.14 ‐‐‐ 1.5 0.21 CR50B Sediment 34.0322 77.7981 4240 ‐ 74.57 9.88 * 13.3057 0.3540 105.70 5.74 254.78 13.85 ‐‐‐ 1.8 0.25 LH40B Sediment 34.1707 77.6041 3690 ‐ 52.74 6.99 * 9.8387 0.3446 53.25 2.39 168.98 7.58 ‐‐‐ 1.9 0.26 LH45B Sediment 34.1870 77.6232 4020 ‐ 64.97 8.61 * 20.4073 0.3500 232.20 9.83 360.80 15.27 ‐‐‐ 1.1 0.15 LH50B Sediment 34.2386 77.6195 4590 ‐ 84.45 11.19 * 20.1980 0.3486 166.10 6.17 259.73 9.64 ‐‐‐ 1.9 0.27 PH40B Sediment 34.1836 77.4586 3573 ‐ 52.54 6.96 * 12.7289 0.3499 89.05 1.99 221.78 4.96 ‐‐‐ 1.4 0.19 PH45B Sediment 34.1876 77.4680 3618 ‐ 55.38 7.34 * 13.2696 0.3559 102.90 4.40 250.04 10.68 ‐‐‐ 1.3 0.18 PH50B Sediment 34.2264 77.4872 4372 ‐ 77.60 10.28 * 15.7561 0.3550 175.30 7.46 357.84 15.22 ‐‐‐ 1.3 0.18 TS40B Sediment 34.0106 77.7256 3413 ‐ 46.81 6.20 * 20.5068 0.3467 220.60 10.30 337.90 15.78 ‐‐‐ 0.83 0.12 TS45B Sediment 34.0328 77.7622 3879 ‐ 63.00 8.34 * 21.0646 0.3554 122.90 3.79 187.86 5.79 ‐‐‐ 2.0 0.27 TS50B Sediment 34.0361 77.7578 3812 ‐ 64.33 8.52 * 17.2378 0.3547 148.70 4.12 277.22 7.67 ‐‐‐ 1.4 0.19 aConstant (time-invariant) local production rate based on Lal (1991) and Stone (2000). A sea level, high-latitude value of 4.5 ± 0.3 at 10Be g-1 quartz was used. bConstant (time-invariant) local production rate based on Heisinger et al. (2002a,b). cShielding factors calculated according to the process outlined in Dortch et al. (2011) dA density of 2.7 g cm-3 was used for all surface samples. eIsotope ratios were normalized to 10Be standards prepared by Nishiizumi et al. (2007) with a value of 2.85 x 10-12 and using a 10Be half life of 1.36 x 106 years. fUncertainties are reported at the 1σ confidence level. gPropagated uncertainities include error in the blank, carrier mass (1%), and counting statistics. hPropagated error in the model ages include a 6% uncertainty in the production rate of 10Be and a 4% uncertainty in the 10Be decay constant. iBeryllium-10 model ages were calculated with the CRONUS-Earth online calculator, version 2.2 (Balco et al., 2008; http://hess.ess.washington.edu/). jZero erosion rates indicate samples were saturated with respect to secular equilibrium.
34