Absent glacial erosion, mountain range height is limited by the rate of bedrock river incision and is thought to asymptote to a steady-state elevation as erosion and rock uplift rates converge. For glaciated mountains, there is evidence that range height is limited by glacial erosion rates, which vary cyclically with glaciations. The strongest evidence for glacial limitation is at midlatitudes, where range-scale hypsometric maxima (modal elevations) lie within the bounds of Late Pleistocene snow line variation. In the tropics, where mountain glaciation is sparse, range elevation is generally considered to be fluvially limited and glacial limitation is discounted. Here we present topographic evidence to the contrary. By applying both old and new methods of hypsometric analysis to high mountains in the tropics, we show that (a) the majority are subject to glacial erosion linked to a perched base level set by the snow line or equilibrium line altitude (ELA) and (b) many truncate through glacial erosion towards the cold-phase ELA. Evaluation of the hypsometric analyses at two field sites where glacial limitation is seemingly marginal reveals how glaciofluvial processes act in tandem to accelerate erosion near the cold-phase ELA during warm phases and to reduce their preservation potential. We conclude that glacial erosion truncates high tropical mountains on a cyclic basis: zones of glacial erosion expand during cold periods and contract during warm periods as fluvially driven escarpments encroach and destroy evidence of glacial action. The inherent disequilibrium of this glaciofluvial limitation complicates the concept of time-averaged erosional steady state, making it meaningful only on long timescales far exceeding the interval between major glaciations.
The height of non-glaciated mountain ranges is limited by erosion along rivers, with rock uplift steepening channels and accelerating channel incision until, in principle, both rates match. This mechanism ties mountain erosion to a low-elevation base level, and channel incision abruptly switches to deposition at this elevation. However, in many mountains there is a transition at high elevations from fluvial to glacial conditions, which introduces a second base level perched close to the glacial equilibrium line altitude (ELA). Above the ELA, rates of glacier incision and coupled supra-ice rock-slope erosion may match or exceed rock uplift rates; below the ELA, in the ablation zone, subglacial incision rates asymptote to zero while fluvial erosion is suppressed. Under these conditions, erosion towards the perched glacial base level is the essential height-limiting mechanism.
For well over a century (Penck, 1905), it has been thought that such glacial limitation of mountain height is widespread. In the late 20th century, the concept of glacial limitation was taken to an extreme and reframed as the “glacial buzz saw” hypothesis (Brozović et al., 1996). The hypothesis is rooted in the observation that mountain ranges high enough for Pleistocene glaciation do not usually rise much higher than the ELA (Fig. 1), and it claims that glacial erosion imposes a near-global topographic ceiling by cutting large swaths of terrain down to this elevation. It has been an important contribution to the broader realization that landscapes evolve under the interacting influence of tectonically driven crustal deformation and climatically modulated erosion (e.g., Molnar and England, 1990; Willett et al., 2006). Debate over the prevalence (or even existence) of a glacial buzz saw cuts to the core challenge of disentangling climatic and tectonic imprints on landscapes.
Hypsometric maximum of glaciated
The strongest proponents of a global glacial buzz saw have implied that erosion can outpace strong tectonic forcing once uplifted rock mass reaches the ELA – a climatically determined elevation. Detractors claim that the power of the glacial buzz saw has been overstated and instead suggest that in most cases glaciations have ornamented, rather than substantially cut, uplifted terrain, which remains at high elevations precisely because erosion is unable to remove it (van der Beek et al., 2009; Hall and Kleman, 2014).
Here, we present a new test of the hypothesis that the height of mountains with terrain above the ELA is set by a perched de facto glacial base level. We focus on a group of tropical mountains whose erosion is dominated by fluvial processes driven by and tightly linked to tectonic uplift, that are unaffected by recent volcanic construction and free of broad areas of internal drainage, and that only encounter the ELA at high elevations. Our rationale is that these mountains will be particularly sensitive recorders of glacial limitation precisely because non-glacial erosion processes have been imprinted on them so strongly. In other words, we assess the potential for glacial limitation in places where fluvial limitation is, a priori, more likely. We ultimately find evidence in nine ranges that glacial erosion limits mountain height. In the most marginally glaciated ranges, such as the Central Range of Taiwan and the Talamanca Range of Costa Rica, glacial and fluvial erosion work in tandem to limit range height to the ELA – a process we term “glaciofluvial limitation”. In contrast, the perched glacial base level is fully developed in those ranges where the most rock mass has been advected through the ELA. This surprising result adds new context to the concept of glacial limitation by showing that a high-elevation, glacial base level has been periodically introduced even in the warmest places.
Glacial limitation refers to the presence of an erosional base level at the ELA and the disconnection of glaciated terrain from fluvial base-level forcing (Egholm et al., 2009). The introduction of this perched base level results in the widening of glacial terrain near the ELA and does not necessarily affect the distribution of peak heights, which are often found on isolated spires far above the ELA (e.g., Brozović et al., 1997; Ward et al., 2012). The concept of base level is usually associated with fluvial erosion, in which the rate of river incision undergoes a zero crossing, typically at an elevation where channelized flow becomes unconfined and spreads laterally, such as at sea/lake level or at a piedmont-alluvial fan apex (Fig. 2a). Above fluvial base level, river channel incision drives hillslope erosion and competes with rock uplift to set range relief. Erosional behavior above and below the perched, glacial base level is analogous. Below the ELA, ice flow spreads laterally, ablates, and slows, driving subglacial erosion rates to zero. The (near-) ELA acts as an erosional base level, above which ice-driven erosion pushes headward into the landscape (Fig. 2b). Glacial erosion ultimately disconnects these landscapes from fluvial base level by blocking channel incision above the glacier terminus – an elevation where glacial erosion is also least effective. During glaciations, continued channel incision below the glacier terminus amplifies the disconnection between fluvial and glacial landscapes.
In mountains thought to be glacially limited, it is common to find peaks that rise more than 1 km above the ELA (Brozović et al., 1997). Glacial limitation allows such high peaks if they are coupled to cirque/valley glacier incision that is effectively controlled by the ELA. The concept is similar to fluvial limitation, such that high peaks rise several kilometers above fluvial base level but are ultimately limited in height by the efficacy of erosion. In the glacially limited case, glacial erosion introduced near the ELA lowers mountains that would otherwise reach higher elevations in fluvially limited conditions.
Tropical mountains have traditionally been considered to be fluvially limited. Their high rates of rainfall drive high rates of fluvial incision and coupled hillslope erosion, and these processes are generally thought to dominate landscape evolution. Although such mountain ranges have been glaciated repeatedly throughout the Pleistocene (Porter, 1989; Hastenrath, 2009), the erosional consequences of snowfall and glacial ice have typically been deemed negligible and limited to very high elevations. Glacial erosion has therefore been discounted as a factor in limiting the height of tropical mountain ranges – with only a few exceptions (e.g., Ring, 2008; Margirier et al., 2016). Our goal in this paper is to assess the possibility of glacial limitation in tropical ranges and by extension to reevaluate the efficacy of glacial limitation in general.
We explore the possibility that glacial erosion has limited the rise of tropical mountains by looking closely at the morphology of tropical highlands. First, we review the evidence for glacial limitation globally and highlight how hypsometry has been the tool of choice in assessing glacial limitation. We then focus on the tropics and undertake topographic and geomorphic analyses of a broad selection of high tropical mountain ranges. We carry out range-scale hypsometric analysis of 10 such ranges to evaluate signs of correlation between mountain height, the ELA, and glacial erosion. Next, we develop a new methodology for analyzing catchment hypsometry in a progressive, nested fashion, which helps with the comparison of hypsometry in mountain ranges of very different sizes. Armed with insights from these analyses, we evaluate them in detail in the Talamanca Range of Costa Rica and the Central Range of Taiwan. These two ranges were chosen because of the seemingly marginal influence of glacial erosion in limiting their height. We find evidence to the contrary: the data suggest that glacial erosion has played a major role in the long-term evolution of both landscapes, but that evidence of its role has been compromised by the headward encroachment of surrounding fluvial terrain during interglacials. The poor preservation potential of glacial morphology in the tropics has led to an underestimation of its importance in limiting range height.
The first articulation of the idea that ice-driven erosion could limit mountain range height came over a century ago, when Dawson (1895) asserted that frost cracking had kept the ridgeline of Kamloops, British Columbia, at a relatively constant elevation. His claim was speculative and called attention to a striking similarity in the elevation of mountain summits and their apparent correlation with the bottom of the frost-cracking window. Penck (1905) put forth a more formal argument for glacial limitation of mountain height around the same time but instead focused on subglacial erosion. In this remarkable paper, Penck described the differences between glacial and fluvial valleys in the Western Alps and proposed that glacial erosion had flattened and widened preexisting fluvial valleys to accommodate ice flux. Penck reasoned that the topographic structure of the Western Alps is geologically young and (in the absence of any plate tectonics framework) hypothesized that glacial erosion had outpaced the rate of rock uplift and limited the height of the range. It was not until the late 20th century, however, that the potential for the widespread limitation of mountain height by glacial erosion was considered, when Porter (1977) and Broecker and Denton (1989) used a compilation of ELA estimates (paleo and modern) from the length of the American Cordillera to demonstrate a qualitative match between peak mountain height and the bounds of Pleistocene ELA fluctuation.
The modern form of this idea is the glacial buzz saw hypothesis, which was first
articulated by Brozović et al. (1996, 1997). They found that hypsometric
maxima (modal average elevations) of large (
In the years since, this hypothesis has gained widespread although incomplete
acceptance. A hypsometric maximum within the bounds of Late Pleistocene ELA
variation is now widely regarded as the signature of the glacial buzz saw
(e.g., Brozović et al., 1997; Brocklehurst and Whipple, 2004; Oskin and
Burbank, 2005; Mitchell and Montgomery, 2006; Egholm et al., 2009; Mitchell
and Humphries, 2015), although
The interpretation of the match between landscape hypsometric maximum and the ELA often links the hypsometric maximum to low-gradient terrain left near the ELA by cirque-driven headward erosion (e.g., Brozović et al., 1997; Egholm et al., 2009). It should be noted, however, that hypsometry is simply a measure of the fraction of the terrain covered by each contour band. A hypsometric maximum is not uniquely determined by topographic gradient: rather, the hypsometric maximum should be thought of as a contour whose combined length and thickness exceeds all other contours. A hypsometric maximum at the ELA is consistent with a lengthening of the ELA contour by headward cirque erosion as well the presence of low-gradient terrain at this elevation. In most cases, the hypsometric maxima found at the ELA probably arise both from contour width (low-gradient terrain) and contour length (erosional penetration) at the ELA.
The strongest empirical evidence for widespread glacial limitation comes
from the global analysis of Egholm et al. (2009), who analyzed large swaths
of topography (
Another line of evidence for glacial limitation has been an apparent acceleration in exhumation rate that coincides with the onset of Pleistocene glaciation. The thought is that if mountain ranges around the world have been subject to glacial limitation, a corresponding increase in exhumation rate should be apparent close to the onset of global glaciation in the thermo-chronometric record. This signal has been borne out on a regional level, such as in the Southern Andes (Thomson et al., 2010), Western Alps (Fox et al., 2015), and Cordillera Blanca (Margirier et al., 2016). Thus, the evidence for glacial limitation consists of both an apparent glacially driven increase in exhumation rate recorded in thermo-chronometry as well as a near-global topographic signature.
The idea that glacial erosion limits mountain height and that mountains would generally be higher in an ice-free world is not universally accepted. A key problem is that it is difficult to disentangle the topographic signatures of non-glacial versus glacial processes (e.g., Hall and Kleman, 2014). Ironically, one of the most problematic cases comes from the birthplace of the glacial buzz saw hypothesis, the northwest Himalaya, where Van der Beek et al. (2009) used a combination of thermo-chronometry and topographic analysis to propose that Pleistocene glaciation took advantage of a preexisting eroded surface – and that the correlation between region-scale hypsometric maxima and the Pleistocene ELA is simply a coincidence of low-gradient topography near the ELA. It has been proposed that this kind of topographic inheritance may comprise the majority of the global signature of glacial limitation (Hall and Kleman, 2014). Another matter of significant controversy is whether mountain ranges were actually subject to enhanced erosion rates during the Pleistocene. Schildgen et al. (2018) analyzed 30 locations where enhanced Pleistocene exhumation rates had been inferred from thermo-chronometric data (Herman et al., 2013) and found that the signal of enhanced exhumation in 27 of those sites could be explained either by bias introduced by inappropriate sampling or by a change in tectonic boundary conditions rather than by climate change. Inferences about global-scale exhumation patterns thus merit reevaluation. Our paper does not directly address these issues and instead evaluates the topographic imprint of glacial erosion, focusing on locations where it has not been recognized before.
The principal reason for skepticism of glacial limitation in the tropics is that regional hypsometry shows (a) very little area near the lower-limit of the Pleistocene ELA and (b) significant hypsometric maxima several kilometers below it (Fig. 1). Since most of these mountains are ridge-and-valley landscapes dominated by fluvial incision, it is reasonable to assume that if glacial erosion has taken place at all, it has only affected isolated portions of the mountain range. A similar rationale is that these ranges need to be well above 3000 m for cold-phase glacial erosion at all, and thus most glacial landscapes in the tropics sit above quickly eroding highlands whose height is thought to be controlled by the rate of fluvial incision. Nevertheless, the observation that mountain height rarely exceeds the ELA includes the tropics (Broecker and Denton, 1989; Egholm et al., 2009).
Seeking to broadly assess the potential for glacial limitation in tropical mountain ranges, we reviewed all those ranges close to the height of the cold-phase ELA (cpELA). Within this broad selection, we sought a subset uncontaminated by processes that would distort or complicate any signal of glacial limitation. In particular, we wanted to avoid any confusion between the passive uplift of low-relief terrain from that of in situ glacial erosion. The potential for such confusion is strongest in the Peruvian/Bolivian Andes, the Sierra Madre of Mexico and Guatemala, and the central highlands of Papua New Guinea, all of which are characterized by large, internally drained plateaus disconnected from external base level; as such, these regions were excluded from our analysis. Volcanism is another complicating factor, in that glaciated volcanoes have undergone a mix of construction and erosion that cannot easily be disentangled. Therefore, glaciated volcanoes such as those in East Africa (Mt. Kilimanjaro and Mt. Kenya) and Papua New Guinea (Mt. Giluwe) were excluded from our analysis. A final constraint was that each range must be circumferentially well-connected to external base level (sea level or lake level) by relatively short fluvial links.
The following 10 tropical mountain ranges remain and were selected for
analysis (Fig. 3):
Leuser Range, Aceh, Indonesia Central Range, Taiwan Talamanca Range, Costa Rica Crocker Range, Borneo Finisterre Range, Papua New Guinea Owen Stanley Range, Papua New Guinea Merauke Range, Papua Mérida Range, Venezuela Sierra Nevada de Santa Marta, Colombia Rwenzori, East Africa
For our tropics-wide comparison we used an ELA range of 3400–4000 m, which
represents the approximate vertical span of the Last Glacial Maximum (LGM) ELA estimated around the
tropics (e.g., Mark et al., 2005; Hastenrath, 2009) and is similar to the
vertical range of LGM ELA in particularly wide mountain ranges with strong
spatial variation in local climate (e.g., Merauke Range, Papua: Prentice et
al., 2005; Mérida Range, Colombia: Stansell et al., 2007). We adopt the
term cold-phase ELA (cpELA) to emphasize
that the ELA has repeatedly descended to roughly this elevation in tropical
mountain ranges during glacial periods in the Late Pleistocene (e.g., Farber
et al., 2005; Barrows et al., 2011), regardless of whether mountains were
high enough to intersect it. Importantly, the cpELA is not strictly
interchangeable with the global LGM (gLGM) ELA, since the timing of
the local LGM appears to have varied (by 10–20 kyr) in the
tropics and subtropics (e.g., Ono et al., 2004). Furthermore, ELA estimates
are based on geomorphic reconstructions that ignore tectonics: as such they
are potentially biased towards higher elevations by postglacial rock uplift.
Where glacial landforms are dated to the LGM such bias is likely negligible.
When age constraints are lacking and an LGM age for the glacial landscape is
in doubt, ELA estimates may be influenced by pre-LGM glacial landforms
subject to rock uplift for longer periods of time: in these cases, the risk
of bias towards higher elevations (
On a global basis, the ELA of interglacial periods (which for convenience we
term the “warm-phase” ELA or wpELA) has generally been 800–1000 m above
the cpELA (Porter, 1989; Broecker and Denton, 1989). We consider the wpELA to
be broadly interchangeable with the modern ELA, although warming and
commensurate glacier retreat during the late 20th century complicates the
definition of the modern ELA. Four of our 10 selected mountain ranges are
currently glaciated and thus intersect the wpELA: the Merauke Range, the
Rwenzori, the Sierra Nevada de Santa Marta, and the Mérida Range. In the
Merauke Range, the only remaining glacier is the Carstensz Glacier, where
Allison and Kruss (1977) and Prentice et al. (2005) both estimated a modern
ELA of 4650 m using aerial photographs (Prentice et al. (2005) found that
ice loss has accelerated at the Carstensz Glacier since the 1970s and
suggested that the most reliable modern ELA estimate was based on 1972
imagery). In the Rwenzori Mountains, Kaser and Osmaston (2002) compiled
aerial photographs and observations from several field expeditions to map
changes in glacier extent between 1955 and 1990. They used field mapping and
aerial photographs from the 1950s and 1960s to estimate a modern ELA of
4600–4700 m. They noted that ice loss has also accelerated in the Rwenzori
during the study period. Less work on wpELA estimation has been done in the
Sierra Nevada de Santa Marta. Wood (1970) mapped glacier extent there using
aerial photographs acquired in 1969. He compared his results to the mapping
efforts of the 1939 Cabot Expedition (Cabot et al., 1939) and demonstrated
significant ice retreat during this time. No wpELA estimate was provided.
Modern glaciers are presently less than 2 km
We conducted a detailed geomorphic study of two mountain ranges: the
Talamanca Range of Costa Rica and the Central Range of Taiwan. The most
recent research has converged on the hypothesis that high elevations were
attained in both ranges after a rapid Plio–Pleistocene acceleration in rock
uplift rate (Morell et al., 2012; Zeumann and Hampel, 2017; Hsu et al.,
2016). This explanation leaves room for only incidental glacial erosion,
despite the long history of glacial geomorphic study in both places (e.g.,
Kano, 1932, 1935; Panzer, 1935; Weyl, 1955). A similar case can be made for
other tropical mountain ranges, such as the Finisterre Range of Papua New
Guinea, where proposed patterns of landscape evolution have not considered
the influence of glacial erosion (Abbott et al., 1997; Hovius et al., 1998),
despite early – but not widely recognized – reports of glacial remnants
there (e.g., Loeffler, 1971). We ultimately narrowed our foci to the
Talamanca Range and the Central Range – both because of the extensive prior
studies of their geodynamics and glacial geomorphology and because of the
apparent disconnect between these two lines of research. In the Talamanca
Range, despite extensive geomorphic mapping carried out over several decades
(Weyl, 1955; Hastenrath, 1973; Bergoeing, 1978; Barquero and Ellenberg, 1986;
Shimizu, 1992; Orvis and Horn, 2000; Lachniet and Seltzer, 2002), the glacial
chronology has been poorly constrained (Orvis and Horn, 2000); to address
this issue, we targeted Cerro Chirripó for
The Talamanca Range is a high section of the Central American Volcanic Arc
that stretches for
Several studies have attempted to link the erosional history of the Talamanca Range to the onset of Cocos Ridge subduction (Morell et al., 2012; Zeumann and Hampel, 2017). A significant disequilibrium observed in Talamanca river networks is thought to record a switch to an (ongoing) higher rate of rock uplift during the last 3 Myr. Zones of anomalously low-relief topography found at moderately high elevations (between 2000 and 3000 m), which are thought to represent an eroded surface that has been advected to its present-day elevation, have been cited as further evidence of a Plio-Pleistocene switch in rock uplift rate (Morell et al., 2012). Importantly, this prior work has excluded glacial erosion as a factor in the long-term evolution of the mountain range. Our work does not address the interpretation that low-relief landscapes found between 2000 and 3000 m elevation are part of an uplifted, eroded landscape, and neither do our results contradict the claim that a major shift in the rate of rock uplift has occurred recently in the Talamanca. Rather, we make a case that sufficient rock mass existed above the ELA for glaciation to occur during LGM and possibly during cold stages prior to the LGM.
The highest landscape of the Talamanca Range is the Chirripó massif, a
low-relief terrain spanning an area of
Orvis and Horn (2000) provided the first rigorous ELA estimate for maximum
ice extent at Chirripó. They suggested an ELA of
The Central Range of Taiwan is the product of the oblique collision of the Luzon Arc and Eurasia, and it is comprised of metamorphosed marine sediments and pre-Cenozoic basement (Suppe, 1981). Due to the oblique nature of collision, it is thought that deformation has propagated from north to south for 5–7 Ma (Byrne and Liu, 2002). The northern 150 km of the Central Range have long been considered a type example of a fluvially driven, steady-state mountain belt (Willett and Brandon, 2002; Stolar et al., 2007). Recent work has countered this idea by putting forth the case that a rapid increase in exhumation rate occurred along the entire strike of the range starting between 1 and 2 Ma (Hsu et al., 2016). Low-relief surfaces found at high elevations (2800–3000 m) in Taiwan are thought to be remnants of an eroded surface formed sometime prior to 1–2 Ma (Ouimet et al., 2015).
Despite contemporary skepticism that glacial erosion has affected Taiwan,
evidence supporting this contention has a long and well-established history,
with the first observations dating back to the early 20th century (Kano,
1932, 1935; Panzer, 1935). Several groups have reported glacial remnants near
the highest peaks of Taiwan's Central Range (Chu et al., 2000; Cui et al.,
2002; Hebenstreit and Böse, 2003; Böse, 2004; Hebenstreit, 2006;
Hebenstreit et al. 2011; Ono et al., 2005; Carcaillet et al., 2007; Siame et
al., 2007) which lie within
ELA estimation in Taiwan has been a challenge for a number of reasons. One
reason is that fluvial erosion of glaciated landscapes has been severe in
places, which makes the maximum extent of glaciation in Taiwan difficult to
ascertain. Previous work has indicated the presence of a glacial diamict at
2250 m in a valley flanking Nanhudashan, far below unambiguous glacial
valleys (Hebenstreit et al., 2006). Using the relict configuration of
glaciated valleys, Hebenstreit (2006) estimated an ELA of 3355 m at
Nanhudashan, specifically employing the terminal-to-summit altitudinal method
(TSAM). Hebenstreit (2006) also used TSAM to estimate an undated ELA of
3400 m at Yushan, a third glaciated massif in southwest Taiwan. Other work
has used the maximum vertical extent of lateral moraines in both Xueshan (Cui
et al., 2002) and at Yushan (Böse, 2004) to propose an ELA of
The principal data sets used in this study are digital topography,
high-resolution satellite imagery, and rock samples collected from Cerro
Chirripó for
We generated a contiguous digital elevation model (DEM) of each tropical mountain range by mosaicking tiles of 1 arcsec SRTM digital topography. Voids were patched with a 3 arcsec void-filled SRTM DEM. Each DEM was projected using a local Lambert azimuthal equal-area projection. Pit filling and drainage delineation were then performed on each DEM using the TopoToolbox (version 2.1) package in MATLAB (Schwanghart and Kuhn, 2010; Schwanghart and Scherler, 2014).
Hypsometry refers to the frequency distribution of elevation. With
normalization, it becomes the probability distribution of elevation (also
known as the altitude–area distribution: e.g., Strahler, 1952; Mitchell and
Montgomery, 2006). It has traditionally been represented as a “hypsometric
integral”, which is equivalent to a cumulative distribution function
(CDF) (Strahler, 1952; Montgomery et al., 2001). When represented as a probability
density function or PDF (e.g., Egholm et al., 2009), it is typically computed
as a histogram. Here we instead use kernel density estimation to compute the elevation PDF. This standard
technique in statistics takes account of sampling (counting) uncertainty and
generates a smooth function that is easier to interpret. Specifically,
hypsometry is calculated as follows:
Hypsometry has been a popular but rather blunt instrument for assessing the
influence of glacial erosion on landscapes. The advent of SRTM digital
topography, first at 3 arcsec and then at 1 arcsec resolution, has
facilitated the global deployment of hypsometry; it has also entailed the
need to set a meaningful scale at which to segment the data. One approach
(Egholm et al., 2009) is to compute the elevation distribution for each
Progressive hypsometry in detail.
We designed an algorithm that concisely describes how hypsometry varies with
the scale of analysis. The development of this algorithm was motivated by the
observation that the hypsometric maximum of large regions can shift by
several kilometers depending on the boundaries of analysis. The algorithm,
which we term “progressive hypsometry” (PH), involves the measurement of
hypsometric maxima in nested catchments whose outlets span from the lowest to
the highest elevations in a mountain range. Progressive hypsometry consists
of three major components: (i) segmentation of the landscape into large
catchments; (ii) calculation of hypsometry along flow paths;
(iii) segmentation into nested subcatchments characterized by a shared modal
elevation. We first segment the targeted mountain range into large ( Map channel network:
define a channel network in each super-catchment using an arbitrary flow
accumulation area threshold traverse downstream from each channel head extend each chain Map PH along the network (Fig. 4):
traverse each chain map along each chain a nested series of subcatchments, one at every channel
pixel for each nested subcatchment, estimate its elevation PDF, its modal
elevation record as a set of Identify all PH “benches”, characteristic nested-catchment modal
elevations (Fig. 5):
perform change-point detection along each chain define the outlet elevation designate the groups of between-jump modal elevations define each bench modal elevation record as a set of concatenate all
We performed progressive hypsometry on the 10 selected mountain belts,
using a low-elevation reference level of 150–250 m in each mountain range.
This reference elevation focuses the analysis just above large depositional
plains, which improves the efficiency of the algorithm. The Rwenzori are an
exception, since they are more than 1000 km from the nearest coast and rise
sharply above lowlands with several large lakes at 1200 m. We used a
low-elevation reference level of 1200 m for the Rwenzori.
Mapping single progressive hypsometry chains on the mountain range
scale.
We compiled observations made during field campaigns at Cerro Chirripó (in 2014, 2016) and Nanhudashan (in 2015), satellite/aerial imagery, and maps produced by previous workers to map glacial landforms and estimate maximum ice extent in both the Talamanca Range and the Central Range. Glacial landscapes are only present at Cerro Chirripó in the Talamanca Range; in Taiwan, glacial remnants are best preserved at Nanhudashan: we focused on these two massifs. We also selected glaciated catchments to compare glacial geomorphic mapping with detailed hypsometric analysis. Glacial advances at both focus sites terminated near 3000 m, so for ease of comparison all glaciated catchments were extracted and delineated (in a GIS) using a common 3000 m outlet elevation.
Glacial valleys at Chirripó generally have a much lower gradient than the
fluvial valleys flanking them, and the very presence of well-preserved
glacial landforms, such as sharp-crested moraines, together with the apparent
absence of Holocene landsliding, indicates that postglacial erosion has been
slow in these glaciated zones. In contrast, signs of relatively fast erosion,
such as frequent landsliding, are apparent in the surrounding fluvial valleys
in satellite imagery and air photographs and were confirmed in the field.
There have been no direct measurements of erosion rate in these fluvial
catchments, but it is thought that a fluvial erosion rate of
To guide our qualitative assessment, we developed a set of rules. First, we
used a binary slope map (with a threshold of 35
Scarp encroachment ratio (SER) calculation.
To quantitatively describe the pattern of fluvial scarp encroachment into glaciated terrain, we define two new metrics. The first, which we term the ELA-relative modal elevation (ERME), measures the difference between hypsometric maxima of glaciated catchments and the estimated LGM ELA. At both Cerro Chirripó and Nanhudashan, this metric is calculated using the local estimate of the LGM ELA. Both massifs are small enough that large climatic gradients are not likely to drive substantial differences in the position of the ELA, as has been documented in much larger tropical mountain ranges (Prentice et al., 2005; Stansell et al., 2007). We hypothesized that the duration and intensity of glacial erosion in all glaciated catchments was similar during the LGM and that glacial erosion would thus leave a characteristic modal elevation at the ELA in all glaciated catchments. We further hypothesized that fluvial scarp encroachment of glaciated terrain would bias hypsometric maxima at the ELA to higher elevations.
The second metric, which we call the scarp encroachment ratio (SER), is an
approximation of the headward distance traveled by the fluvial escarpment
into each glaciated catchment (Fig. 6). This distance is expressed as a ratio
of scarp-affected terrain to all terrain in glaciated catchments. In each
glaciated catchment, we found the area below and above the escarpment,
Six samples were collected for
In Fig. 8a2–j2 we compare the range-scale hypsometry of each of the selected mountain ranges to the tropics-wide cpELA band at 3400–4000 m. Most of the selected ranges show that the fractional area occupied by each elevation band decreases steadily with increasing altitude, although in some cases, such as the Finisterre Range (Fig. 8e2), topographic plateaus are evident at high elevations and are recorded as secondary hypsometric maxima. In some cases, the highest elevations fall within the bounds of the tropical cpELA, and in other cases peaks extend far above it, but none of these mountain ranges has a significant hypsometric maximum at or above the cpELA. In other words, each range was high enough for cold-phase glacial erosion (with the exception of the Leuser Range), and yet the extent and prevalence of glaciated landscapes in all of them appears small relative to the size of fluvial landscapes.
Progressive hypsometry and mountain range hypsometry for selected
mountain ranges (Fig. 3).
Nine of the 10 selected mountain ranges have catchments with a PH modal
elevation at the cpELA, which we consider a record of those catchments
having established a perched glacial base level (Fig. 8b1–j1). Variability
in glacial influence between each range is best assessed on the basis of
paired outlet and modal elevations. For example, catchments in some ranges
have an outlet elevation as low as 150 m and a modal elevation between the
cpELA and wpELA. In these mountain ranges, the progressive hypsometry of the
largest catchments is thus dominated by glacial erosion. In Fig. 8f1–j1,
these large catchments are represented by a left spread of points within the
ELA band. In other ranges, catchments with a modal elevation within the
bounds of the cpELA are only found above
Cerro Chirripó glacial geomorphology.
These results stand in contrast to range-scale hypsometry, which indicates minor glacial influence across all of the selected mountain ranges. Progressive hypsometry instead highlights variability in the prevalence of glaciated terrain and points to glacial erosion having had an influence in most of these mountain ranges and a particularly significant role in some. Our two focus sites in Costa Rica and Taiwan are examples of the former – where glacial erosion has apparently had only a marginal influence.
Classic examples of glacial landforms are found at Cerro Chirripó (Fig. 9). Several valleys host kilometer-scale lateral moraines, and most valleys are blanketed by glacial till and recessional moraines. Striated and scoured bedrock is prevalent in the paleo-accumulation zone in the major valleys emanating from the Chirripó peak. Most glaciated catchments at Chirripó have a modal elevation within several meters of the estimated local LGM ELA of 3500 m (Fig. 11c). Two high catchments have a modal elevation that is somewhat higher than the local ELA. An important feature is that these catchments appear to be heavily modified by scarp encroachment, which leads us to infer that fluvial scarp encroachment has erased the lower portion of their glacially eroded topography and has biased their hypsometric maxima to higher elevations. Field observations corroborate this inference (e.g., Fig. 6c–e).
Nanhudashan glacial geomorphology.
Our
Glacial landforms are not preserved as clearly at Nanhudashan (Fig. 10) as they are at Chirripó, although scoured bedrock and recessional moraines have been mapped and dated to the last glacial (Siame et al., 2007; Hebenstreit et al., 2011). Perhaps the most striking feature of the glacial landscape at Nanhudashan is its asymmetry. The southeastern glacial valley of Nanhudashan is about 2 km long and has the best preserved glacial remnants. A small plateau that fed north-flowing glacial ice has been largely removed by scarp encroachment from the southwest. Recessional moraines in the northwestern glacial valley disappear abruptly below an erosion front just below 3400 m. In the southwest, almost no glacial remnants remain. Glaciated catchments at Nanhudashan have a hypsometric maximum at or above the estimated LGM ELA of 3400 m. In the best preserved glacial valley (southeast), the hypsometric maximum is within 35 m of the ELA (Fig. 11d). The hypsometric maxima of the other two glaciated catchments are more than 100 m above the estimated LGM ELA (Fig. 11d). Scarp encroachment has been far more severe in these valleys.
ERME and SER reveal similar patterns of glacial erosion and scarp encroachment at Chirripó and Nanhudashan; only scarp encroachment is more advanced at Nanhudashan (Fig. 12). Of the three glacial catchments at Nanhudashan, two show greater scarp encroachment than any at Chirripó, a pattern reflected in their ERME and SER values. The escarpments propagating into glacially eroded terrain at Nanhudashan, some of which were mapped in previous studies (e.g., Hebenstreit et al., 2011; Hebenstreit, 2006; Willett et al., 2014), are particularly spectacular. When combined, our observations at Chirripó and Nanhudashan capture a continuum of scarp encroachment into glacially eroded landscapes and the alteration of glacial-type hypsometry.
Glacial valley hypsometry at Chirripó and Nanhudashan.
ELA-relative modal elevation (ERME) and scarp encroachment ratio
(SER) for each catchment at Chirripó (blue) and Nanhudashan (green)
(names found in Fig. 11). Kendall's tau (0.64) reported for the entire data
set. One-tailed significance test yields a
Our hypsometric analyses reveal signs of glacial limitation recorded in the topography of multiple tropical mountain ranges. While the application of traditional hypsometry provides hints of glacial influence, the new tool of progressive hypsometry amplifies its topographic signal and exposes evidence that glacial erosion has had a profound effect on the height of these mountain ranges. This inference was tested at focus sites in Taiwan and Costa Rica, expressly chosen because their exposure to glacial erosion has been apparently weak to negligible. Even in these seemingly marginal examples, we find substantial evidence that valleys with a modal elevation near the ELA originate in glacial erosion. Our observations support the claim that a glacial base level has been periodically introduced in mountain ranges throughout the tropics.
Glacial limitation is expressed differently in tropical mountains than in
midlatitude mountains because the high cpELA imposes a fundamentally
different relationship between tropical glaciers and their flanking fluvial
systems. At midlatitudes, the cpELA descends to
Progressive hypsometry for mountains uplifted through ELA and for
glacial limitation.
The long-term evolution of the high-elevation, glacial base level has proceeded differently between tropical mountain ranges, and for this reason it is important to consider the potential for glacial limitation in the context of all of the mountains analyzed. Three factors determine the development and preservation potential of a glacial base level in the tropics: (i) the volume and pattern of rock uplift through the cpELA; (ii) the efficacy of glacial erosion; and (iii) fluvially driven destruction of glaciated terrain. Valleys are initially disconnected from fluvial base level once enough rock mass has passed through the cpELA for glacier formation. If glacial erosion is strong enough during the initial phases of glaciation, it drives the development of a modal elevation at the cpELA. During interglacials, fluvially driven scarp encroachment removes portions of perched, glaciated valleys. If glaciated landscapes above the cpELA are spared scarp encroachment, continued rock uplift and subsequent cold-phase glacial erosion drives further development of the glacial base level at the cpELA. If not, evidence of the perched glacial base level is erased. Progressive hypsometry appears to capture a continuum of this behavior.
Landscapes can be thought of as a mosaic of catchments. Progressive hypsometry objectively distinguishes these catchments on the basis of their modal elevation, and in doing so offers a fine-scale perspective on the distribution of terrain (Fig. 8a1–j1). The pairing of modal elevation and catchment outlets also provides an indication of how catchment geometry varies with elevation. For example, catchments with a large vertical gap (Fig. 13c) between their modal elevation and outlet must occupy disproportionately broad areas at high elevations. This pattern reflects either perched, low-gradient terrain or erosional penetration and contour elongation at the modal elevation.
Headward glacial erosion both penetrates the landscape near the ELA and promotes the formation of low-gradient terrain. When mountains first pass through the cpELA, this effect is minor. If glacial erosion repeatedly acts on the same landscape through multiple glaciations, the relative strength of the modal elevation at the ELA increases. In Fig. 8a1–j1, we highlight those catchments with a (PH) modal elevation within the bounds of the cpELA and wpELA. In the Central Range of Taiwan, Talamanca Range, Finisterre Range, Crocker Range, and Owen Stanley Range, these glaciated catchments sit at the highest elevations and are nested inside fluvial catchments (Fig. 8b1–f1). In the Merauke Range, Mérida Range, Sierra Nevada de Santa Marta, and the Rwenzori (Fig. 8g1–j1), rock mass has been advected high above the cpELA (reaching as high as 5700 m in Santa Marta). In these four mountain ranges, however, the largest vertical gaps between catchment modal elevation and outlet are found at the cpELA.
Our data show that once a catchment reaches the cpELA, the bulk of its terrain does not rise much higher. The strongest evidence of this limitation phenomenon is found in Fig. 8f1–j1, where the (PH) modal elevation of the largest catchments sits between the cpELA and wpELA. Furthermore, we do not observe any significant catchment (PH) modal elevations above the wpELA, even though peak elevations often extend high above it. These observations are consistent with the introduction of a glacial base level at the cpELA, with supra-ELA terrain ultimately tied to an expanding zone of cirque/valley glacier incision. Progressive hypsometry thus supports the idea that glacial limitation is a viable mechanism in some tropical mountain ranges.
Summary of evidence for glacial limitation. Left spread (LS) refers to high-elevation hypsometric maxima with low-elevation outlets in progressive hypsometry (Figs. 8a1–j1, 13c). LS-ELA match: The strength of the match between the left spread and the ELA is assessed in Fig. 8a1–j1. Sub-ELA gap: the sub-ELA gap refers to the absence of left spread below the ELA. Glacial features: z – cirques; y – terminal moraines; x – lateral moraines; w – retreat moraines; u – overdeepenings. Glacial features in each place have been documented in the literature and were confirmed by the authors in the available imagery (e.g., Google Earth).
The Leuser Range is the only range among those analyzed that bears no signal of glacial erosion, and so we consider its progressive hypsometry as a null reference model by which to compare the progressive hypsometry of other glaciated ranges. In Fig. 13, we vertically exaggerate the progressive hypsometry of the Leuser Range to schematically illustrate the evolution of a mountain range that is spared glacial limitation as it rises through the cpELA. In the two null examples (Fig. 13a, b), rock uplift alone drives the growth of landscapes above the ELA, or the long-term effect of glacial erosion is compromised by fluvial erosion during interglacial periods. In Fig. 13a, the range is high enough for glaciation, but glacial erosion is either absent or negligible or largely erased by fluvial erosion. Similarly, in Fig. 13b, glacial erosion has been either been negligible or severely compromised, and no significant record of glacial erosion has been left, even though more rock mass has passed through the ELA. In contrast, in Fig. 13c, the process of glacial limitation focuses catchment growth at the ELA and glacial landscapes are preserved enough to leave a strong signal in progressive hypsometry. In this case, catchments enlarge near the ELA as more rock mass is pushed through it, strengthening the modal elevation observed there. Only isolated spires of rock reach higher elevations.
Considered in isolation, most of the mountain ranges analyzed would not be classified as glacially limited. Yet across the tropics, the highest mountain ranges bear the strong appearance of glacial limitation (Table 1). We next explore the potential for glacial limitation is those mountain ranges where a glacial influence is present but weaker, namely, the Central Range of Taiwan, Finisterre Range, Talamanca Range, Crocker Range, and Owen Stanley Range. The size of these glacial landscapes is limited by either a dearth of rock mass that has been advected through the cpELA or by the fluvial destruction of glaciated terrain. These mountains are therefore not glacially limited in the same sense that the other selected mountain ranges are. Nevertheless, their highest elevations all coincide with cpELA.
We have presented evidence for the periodic introduction and subsequent
removal of a glacial base level in the Talamanca Range and the Central Range of
Taiwan. Both ranges were glacially eroded during the LGM; their glaciated
catchments both have a modal elevation near the cpELA; and in both
landscapes, postglacial fluvially driven erosion has erased significant
fractions of their glaciated terrain. These observations are consistent with
three possibilities.
The question of which scenario best describes the Talamanca Range and the
Central Range hinges on whether glaciations have been a rare (or even
isolated) occurrence in these places. Unfortunately, our evidence highlights
barriers to answering this question. Specifically, glacial landscapes are
prone to erasure in both places (Fig. 12). In Taiwan, erasure of glacial
landscapes appears to progress rapidly, and these
landscapes are unlikely to survive the long,
Given the pattern of glacial limitation observed throughout the tropics, it is possible that the absence of significant terrain at the cpELA in Costa Rica and Taiwan is due to particularly effective scarp encroachment rather than ineffective glacial erosion. Comparable mountain ranges in the tropics have all been subject to significant glacial erosion upon passage through the cpELA, and it would be a remarkable coincidence if the Talamanca Range and the Central Range as well as the Finisterre Range, Crocker Range, and Owen Stanley Range were all independently limited at the same elevation close to the cpELA. Furthermore, fluvial erosion has not limited rock mass to elevations below the cpELA in either Costa Rica or Taiwan. Moderately high-elevation (2000–3000 m), low-relief landscapes in both ranges that have been disconnected from external base level are further evidence that fluvial limitation in general is unlikely (Morell et al., 2012; Ouimet et al., 2015). In other words, if fluvial erosion can limit the height of these mountain ranges, it would likely do so at an elevation above the cpELA and would ultimately be subject to glacial limitation in the future. We therefore propose glaciofluvial limitation as the most parsimonious explanation for the coincidence of high peaks and the cpELA in the Talamanca Range and the Central Range of Taiwan (Fig. 14).
Mountain ranges are thought to evolve towards a steady state ultimately determined by the efficacy of fluvial erosion. Steady state is assessed on the basis of material flux, in which the accretion of rock and sediment are balanced by denudation, and topography such that mountain range height and width are invariant over appropriate spatial and temporal averaging scales (Willet and Brandon, 2002). Thermal steady state (time-invariant temperature field) is a prerequisite for flux steady state (Willett and Brandon, 2002). For decades, the Central Range of Taiwan has been considered a steady-state, fluvially limited mountain range (Suppe, 1981; Willett and Brandon, 2002; Stolar et al., 2007) on the basis of its relatively uniform height and width (Stolar et al., 2007) and a correlation between long- and short-term denudation rates (Dadson et al., 2003).
It is difficult to reconcile the steady-state model with evidence for periodic glacial erosion at high elevations in Taiwan. This phenomenon is equivalent to introducing an upper boundary condition on mountain growth, which violates the conditions required for fluvially driven steady state; essentially, glacial erosion imposes an upper limit on river channel steepness. Furthermore, it would be a striking coincidence for the Central Range to have reached a fluvially driven steady state at the cpELA, an elevation at which a glacial base level is apparent in eight other similar mountain ranges. We suggest that in the long term, an oscillatory pattern of glacial and fluvial erosion maintains the Central Range at its regionally uniform height.
Evidence for glacial limitation is widespread in high tropical mountains and has largely been overlooked. One reason is that the greater relative extent of fluvial versus glacial landscapes obscures the effects of glacial erosion when carrying out traditional hypsometric analysis. A second reason is that fluvially driven scarp encroachment can erase signs of glaciation. We have addressed both issues and have presented evidence that glacial erosion has had a profound effect on limiting range height in rapidly uplifting and eroding tropical mountain ranges. Even in the most marginally glaciated mountain ranges in the tropics, glaciofluvial limitation may drive long-lived, cyclic competition between glacial erosion and fluvial scarp encroachment, thereby preventing the establishment of steady state in the classical sense.
Schematic of glaciofluvial limitation.
The MATLAB package, data, and results used in the progressive hypsometry
analysis are located in three repositories and are available for download.
Scripts for running the progressive hypsometry analysis can be downloaded
from the PHtools repository:
We have provided documentation detailing how to implement the software as
well as an explanation of the organization of data and results in their
respective repositories at
In addition, as part of the Supplement we have also provided the raw data
used for
The supplement related to this article is available online at:
MTC helped design the project and formulate the analyses, performed fieldwork, laboratory analyses, and topographic analyses, and wrote the initial draft of the paper. CPS conceived and helped design the project and guided the analysis and data interpretation. MRK helped guide the research focus, assisted in leading field expeditions to Costa Rica, and assisted in data interpretation. MRK and JMS both guided laboratory work. All authors contributed to crafting the initial draft. MTC and CPS led the revision efforts.
The authors declare that they have no conflict of interest.
This research was supported by the Climate Center of Lamont-Doherty Earth Observatory and NASA GISS. We are grateful to the Área de Conservación La Amistad Pacífico of the Sistema Nacional de Áreas de Conservación of Costa Rica for permission to conduct research at Cerro Chirripó and to Chirripó National Park for providing accommodations during fieldwork. Lionel Siame generously led Maxwell T. Cunningham and Colin P. Stark in the field at Nanhudashan. We are grateful to Roseanne Schwartz, Jean Hanley, and Jeremy Frisch, who all assisted with the preparation of samples for exposure age analysis. We are also grateful to David Egholm and Peter van der Beek for their constructive remarks during review, which greatly enhanced the quality of the paper. This work is LDEO contribution no. 8781. Edited by: David Lundbek Egholm Reviewed by: Peter van der Beek