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Development of a steep erosional gradient over a short distance in the hyperarid core of the Atacama Desert, northern Chile
Journal Pre-proof Development of a steep erosional gradient over a short distance in the hyperarid core of the Atacama Desert, northern Chile
Joel Mohren, Steven A. Binnie, Benedikt Ritter, Tibor J. Dunai PII:
S0921-8181(19)30553-3
DOI:
https://doi.org/10.1016/j.gloplacha.2019.103068
Reference:
GLOBAL 103068
To appear in:
Global and Planetary Change
Received date:
29 May 2019
Revised date:
17 October 2019
Accepted date:
25 October 2019
Please cite this article as: J. Mohren, S.A. Binnie, B. Ritter, et al., Development of a steep erosional gradient over a short distance in the hyperarid core of the Atacama Desert, northern Chile, Global and Planetary Change(2019), https://doi.org/10.1016/ j.gloplacha.2019.103068
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© 2019 Published by Elsevier.
Journal Pre-proof Development of a steep erosional gradient over a short distance in the hyperarid core of the Atacama Desert, northern Chile Joel Mohren*1, Steven A. Binnie1, Benedikt Ritter1, Tibor J. Dunai1
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Institute of Geology and Mineralogy, University of Cologne, Germany
Zülpicher Str. 49b, 50674 Cologne
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*Corresponding author. Email: [email protected]
Abstract
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Although generally considered to exist under hyperarid conditions over the long term, landscapes in
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many parts of the Coastal Cordillera of northern Chile have undergone fluvial erosion. Small-scale drainage systems in this mountain range are mostly isolated from river networks and associated
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processes originating in the Precordillera or the High Andes to the east, thus providing natural
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laboratories to investigate the interplay between erosion, atmospheric deposition, tectonics and local (micro-) climatic conditions. In this study, we present a set of cosmogenic
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Be and
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catchment-wide erosion rates along a short (2.5 km) E-W transect on the northern rim of the Río Loa Canyon in the Coastal Cordillera of northern Chile (latitude 21.4°S). Here, a flat sedimentary gravel surface, which was deposited before the Middle Miocene, becomes increasingly dissected and changes into a badland-like topography to the west. The 10Be erosion rates increase by approximately an order of magnitude from east to west, reflecting (1) localized tectonic movements, (2) geologically recent base level lowering, (3) time-integrated (micro-) climatic gradients and (4) the presence/absence of gypcrete. These findings are corroborated by analysis of geomorphologic parameters, which point towards the presence of two fundamentally different erosional regimes in this small study area. These regimes are sharply delineated along a topographically modest tectonic ridge. To the west, a detachment-limited erosion regime prevails, while in the east a transport-limited regime is dominant. The presence or absence of gypcrete, whose prevalence is governed by (micro-) climatic conditions, generally appears 1
Journal Pre-proof to reflect the respective erosional regimes. The erosion rates we infer point to a long-term process of differential drainage evolution in the study area, likely on timescales of millions of years.
Keywords: Geomorphology; Gypsum; Erosion; Crust; Cosmogenic nuclides; Atacama Desert
1 Introduction The formation of topography is directly linked to a variety of endogenic and exogenic processes, such
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as tectonics, erosion, aggradation and weathering. The efficiency of exogenic processes is influenced by parameters such as rock type, climate and vegetation cover and their likely interplay, but the differing
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timescales over which such Earth-shaping processes integrate and variations in their influence over time
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can blur our understanding of long-term topography formation. Thus, for the study of the interactions between Earth-shaping processes, environmental conditions and their impact on topography formation,
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it is often desirable to isolate the influence of individual variables, while maintaining consistency in other factors over the timescale that is being analyzed. For timescales on the order of > 104 yrs,
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encompassing climatic cycles, this requirement is rarely met. On the other hand, the question about how
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changing environmental conditions affect topography formation is – also in the context of modern
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climate change and its impact on future conditions – one of the most important issues in the field of geomorphology, underlining the relevance of the long-term perspective. In this study we analyze a unique landscape located within the hyperarid core of the Atacama Desert in northern Chile (cf. Ritter et al., 2018b; Ritter et al., 2019). The onset and secular variation of aridity in the Atacama Desert, as well as spatial variability, are still a matter of scientific debate. Estimates for the onset of hyperaridity range between the Pleistocene and the Oligocene (e.g. Alpers and Brimhall, 1988; Amundson et al., 2012; Betancourt et al., 2000; Carrizo et al., 2008; de Porras et al., 2017; Dunai et al., 2005; Evenstar et al., 2009; Evenstar et al., 2017; Gayo et al., 2012; Hartley and Chong, 2002; Jordan et al., 2014; Nishiizumi et al., 2005; Rech et al., 2006; Reich et al., 2009; Ritter et al., 2018b; Ritter et al., 2019; Sáez et al., 2012; Sáez et al., 1999; Sillitoe and McKee, 1996; Veit et al., 2016; Wang et al., 2015). The geological record points to predominantly arid conditions in the region since 150 Ma (Hartley et al., 2005). Long-term stable climatic conditions in the Coastal Cordillera of northern Chile 2
Journal Pre-proof are achieved by (1) the steady latitudinal location of the South American continent since the MidCretaceous (Somoza and Zaffarana, 2008); (2) the rainshadow effect of the High Andes blocking westward moisture transport originating from the Amazonian region (e.g. Houston and Hartley, 2003; Weischet, 1966); (3) the cold Peru-Chile Current cooling the air masses and thus reducing the possible amount of absolute humidity to be transported with onshore winds along the western coast of South America (e.g. Houston, 2006a; Stoertz and Ericksen, 1974), and (4) the perennial atmospheric subsidence caused by the Hadley circulation (e.g. Houston, 2006b; Houston and Hartley, 2003) and the
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Rutllant circulation (Rutllant, 2003). Taken together, these factors lead to an excess of mean annual
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potential evaporation by two to three orders of magnitude when compared to mean annual precipitation
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in the Atacama Desert (Houston, 2006a). Some areas within the hyperarid core of the desert, which is located between ~19°S and ~22°S and includes the Coastal Cordillera and the Central Depression (cf.
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Ritter et al., 2018b), currently receive less than 2 mm yr-1 of precipitation from rainfall (e.g. Houston,
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2006a). In these areas, signs of significant fluvial erosion are rare, and most of the landscape appears to be fossilized under a blanket of gypsum dust and gypcrete (Ericksen, 1981; Mortensen, 1929; Mortimer,
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1980). Exposure ages in these ‘fossilized’ areas are predominantly of Miocene age (cf. Carrizo et al.,
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2008; Dunai et al., 2005; Evenstar et al., 2017; Nishiizumi et al., 2005; Ritter et al., 2018b). Further to the west, within the Coastal Cordillera in the vicinity of the coast, apparently well-developed
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drainage networks indicate a fluvial response to base level change and/or to increasing precipitation. However, significant fluvial erosion in this area (~19°S to ~22°S) is restricted to (1) the Río Loa, a major drainage system crossing the entire Coastal Cordillera from the High Andes to the Pacific Ocean and (2) local catchments, i.e. catchments hosted entirely within the Coastal Cordillera. In the latter, only a few incise down to the base level (sea-level or coastal abrasion platform) as many of these local drainages terminate as hanging valleys high on the coastal cliff, illustrating that incision into the face of the cliff is slower than cliff retreat (Mortimer, 1980). The catchments that do incise to base level have very low denudation rates, ~2 m Myr-1 for catchments near our study area, in spite of their steep gradients (~400 to ~700 m mean local relief on 5 km radius; Starke et al., 2017). Our study area incorporates a zone of fluvially shaped topography near the coast, in the immediate vicinity of the northern rim of the Rio Loa canyon. This particular setting allows the investigation of 3
Journal Pre-proof spatially resolved erosion rates and surface processes along a west-east transect that displays a progressive reduction in the amount of fluvial incision despite experiencing a common base-level (i.e. the Río Loa, Fig. 1). The catchments are small (< 0.3 km2; Figs. 1, 2) and so record the affects of longterm local climatic conditions, rather than an allochthone climate signal (cf. Ritter et al., 2018b; Ritter et al., 2019). Furthermore, the catchments in the study area incise into a rather homogeneous and uniform substrate, consisting of two sedimentary units which have been successively deposited in the time period between Eocene?-Upper Oligocene and Miocene (Sepúlveda and Vásquez, 2019; Vásquez et al., 2018).
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Irrespective of their age, both of these alluvial layers presumably behave similarly to fluvial erosion,
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meaning that substrate/lithological parameters can be considered constant throughout the area studied.
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As such, our study location provides a unique natural laboratory for the investigation of the temporal
the hyperarid core of the Atacama Desert.
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and spatial variability of precipitation and its impacts on the formation of topography on the fringe of
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Our results from cosmogenic nuclide measurements and digital terrain model analysis confirm the presence of a time-integrated gradient in erosional style and rates. We will show that the likely causes
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for this gradient relate to: moisture sources (Pacific); the interception of that moisture by subtle,
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2 Regional setting
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intervening, fault-driven relief; and the resulting presence/preservation of gypcrete.
2.1 Topographic characteristics
The study area is a spatially short (~2500 m) E-W transect located on the northern rim of the Río Loa canyon, at an elevation of about 700 m a.s.l. within the Coastal Cordillera (21.42° S, 69.98° W; Fig. 1). Approximately 8 km to the west, the Río Loa drains into the Pacific Ocean; the closest cities are Iquique to the north (~130 km), Quillagua to the southeast (~50 km) and Tocopilla to the south (~80 km).
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Figure 1. Sentinel 1B true color image (L2A product) showing the location of the study area in the
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Coastal Cordillera at the northern rim of the Río Loa canyon, northern Chile. Fault systems in the vicinity of the study area are taken from the geological map of this region (Vásquez et al., 2018) or
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based on own observations and data analysis. A strong E-W gradient in fluvial erosion is apparent, with
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pronounced incision in the west and limited incision to the east. The study area straddles this gradient where it is most pronounced. Three climate stations (#11, #12, #13), from which meteorological data
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has been analyzed for this study (Hoffmeister, 2018), are indicated by white stars.
We decided to conduct our work in this area because the topography and surface patterns change remarkably along the short spatial transect. In the eastern portion, the landscape can generally be characterized as smooth, with gypsum-rich chuca and costra (sensu Ericksen, 1981) covering almost the entire surface, which is gently dipping to the southwest. The small-scale drainage systems (< 0.3 km2) are predominantly N-S oriented and terminate as hanging valleys in the rim of the much larger and more deeply incised Río Loa canyon. Local relief is generally low and slopes show high convexities (Fig. 2). A significant proportion of the eastern sector does not show any signs of fluvial erosion as surfaces are flat and encrusted, with costra being exposed. Gypsum crust has formed polygonal structures and channel bottoms are partly covered by chuca. In some places, especially along 5
Journal Pre-proof slopes and a few meters below the rim of the northern flank of the Loa canyon, a > 20 cm thick salt layer
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mainly consisting of halite and sometimes forming nodules, crops out (Fig. 10).
Figure 2. Handheld camera (A-B) and DJI Spark drone (C-F) pictures illustrating the geomorphological E-W contrasts in the study area. The channels of the western catchments (groups W1 and W2; see Fig. 3) have incised relatively deeply into the gravels (A, C). In contrast to that, the relief in the eastern catchments (groups E1 and E2) is much lower (B, D with BA15-016 in the foreground). A local ridge 6
Journal Pre-proof sharply separates encrusted (east) from non-encrusted surfaces (west; E). However, small patches of gypsum crust can be found in the western catchment BA15-004, covering a small slope and a plateau above subcatchment BA15-006 (dashed polygons in C and F). In some areas, salt crops out along slopes (white arrows in C and F; see also Fig. 10).
By contrast, fluvial erosion has formed a well-developed drainage system in the western portion of the transect. In comparison to the eastern drainages, incision appears to be much greater. Signs of sub-recent
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discharge, such as minor accumulations of clay in some channels and centimeter-scaled incision into
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this layer are observable, as are scars and deposits from slope failures. Although the local erosional base
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level is determined by the incision of the Río Loa to the south, the drainage patterns are mainly W-E/EW oriented. This is due to the presence of a N-S oriented local ridge into which the catchments have
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incised. The ridge is partly covered by a thick gypsum crust, up to ~40 cm on the ridge peaks and
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spatially coinciding with the sharply defined and easy-to-trace westward limit of gypsum surface cover (Fig. 2). To the west of the ridge, gypsum encrustation is absent, apart from some highly localized
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patches, and bare gravel is exposed at the surface. Many surface cracks, similar to those described by
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eastern sectors.
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Allmendinger and González (2010, and references therein), can be observed in both the western and the
2.2 Geological and geomorphological background Large parts of the Coastal Cordillera have been strongly affected by extensional tectonics since the Oligocene (Hartley et al., 2000). The tectonic regime is likely caused by the strong plate coupling between the Nazca Plate and the South American Plate during the subduction process and triggers crustal shortening (e.g. Allmendinger and González, 2010). Numerous N-S striking normal faults belonging to the Atacama Fault System (AFS; e.g. González et al., 2003), visible all over the Coastal Cordillera, are evidence for stress release. The faults promoted the formation of horst and graben structures, which can be identified along the Río Loa Canyon (González et al., 2003; Sepulveda and Vásquez, 2018, 2019; Vásquez et al., 2018). The bedrock below the alluvial surface of the study area is exposed along the ~600 m deep Loa Canyon and mainly comprises of Carboniferous metamorphic rocks (Caleta Loa 7
Journal Pre-proof complex, Csbcl) and the Upper Triassic to Lower Jurassic Sierra de Lagunas strata, mainly composed of andesites (TrJsl; Vásquez et al., 2018). Uplift of these rocks accommodated by the Atacama Fault has possibly been initiated during the Eocene and continues until the present day (Cosentino and Jordan, 2017; Juez-Larré et al., 2010). During the Early Miocene, tectonic uplift isolated large areas of the Coastal Cordillera from the influence of Precordilleran sediment aggradation (Evenstar et al., 2017; Mortimer et al., 1974). In a narrow corridor between 19° and 21.6°S latitude, trench-parallel shortening also causes E-W oriented stress release along reverse faults with pure dip slip movement (Allmendinger
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and González, 2010; Allmendinger et al., 2005). The E-W striking reverse fault systems are believed to
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have been active at least since Late Miocene and some could have remained active until the present day
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(Allmendinger and González, 2010; Allmendinger et al., 2005; Carrizo et al., 2008; González et al., 2015; Ritter et al., 2018b). The study area is located between two of these E-W striking reverse faults
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systems, namely the Chipana fault in the north and the Adamito fault to the south (Skármeta and
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Marinovic, 1981; Vásquez et al., 2018; Fig. 1). These faults are believed to define the southern and northern limits of the Chipana-Adamito block, which has possibly undergone uplift since 16-4 Ma
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(Vásquez et al., 2018) or, in the case of the Adamito fault, since 14-10 Ma (Ritter et al., 2018b). The
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uplift likely enhanced the isolation of the area from sediment sources to the north, south and east (cf. Ritter et al., 2018b; Vásquez et al., 2018). This interpretation is supported by the presence of two
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chronostratigraphic units that cover the underlying bedrock in the study area by a thickness of up to 300 m, as recently mapped by Sepúlveda and Vásquez (2019) and Vásquez et al. (2018). According to these authors, rounded to well-rounded conglomerates with a variable degree of consolidation and of possibly Eocene to Upper Oligocene age (Cañón del Loa beds, CLB) are partially covered by generally angular and moderately consolidated gravels with a maximum age of 24.69 ± 0.09 Ma (Alto Hospico gravels, AHG). On the side of the Loa canyon opposite to our study area, i.e. on its southern slopes, a tephra layer located within this unit 30 to 50 m below the surface has been dated to 21.4 ± 0.5 Ma, while a second tephra layer located close to the base of the AHG yielded an age of 23.50 ± 0.16 Ma (Vásquez et al., 2018). The minimum age obtained from this unit on the Chipana-Adamito block is 16.14 ± 0.02 Ma (all ages derived from 40Ar/39Ar dating in biotite; Vásquez et al., 2018), indicating an Early Miocene age for the entire AHG unit in the vicinity of our study area. It is likely that the majority 8
Journal Pre-proof of these gravels were deposited in an alluvial paleoenvironment under arid paleoclimatic conditions (Sepúlveda and Vásquez, 2019; Vásquez et al., 2018). The gravels of both strata are of varying lithology, including volcanic, metamorphic, intrusive and sedimentary rocks (cf. Sepulveda and Vásquez, 2018). Clast sizes as observed in the field cover the entire gravel and boulder spectra; Sepulveda and Vásquez (2018) report mean clast sizes of 6-25 cm for the CLB and 1-2 cm for the AHG, respectively. Lamination patterns consisting of decimeter thick layers in the CLB can be traced along slopes in the western portion of the study area. Gravels of the CLB unit are poorly sorted and generally cemented by salt
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(predominantly halite). In between the two gravel sequences a thick paleosol layer has been identified,
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indicating a significant temporal interruption of aggradation processes after the deposition of the Cañón
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del Loa beds, possibly between 31 and 26 Ma (Sepúlveda and Vásquez, 2019). Most N-S striking faults in the direct vicinity of the study area were likely established prior to the
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deposition of the Cañón del Loa beds (Sepulveda and Vásquez, 2018) and were still active, or
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reactivated, after the deposition of the Alto Hospicio gravels and coeval units in the Coastal Cordillera, as the sedimentary cover shows an offset in some places (e.g. González et al., 2003; Sepulveda and
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Vásquez, 2018). This holds also for a N-S striking splay fault of the AFS to the east of the study area,
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which could have been active since about 14 Ma (Ritter et al., 2018b). However, Pleistocene displacements along N-S faults belonging to the AFS have been reported further to the South (~23.5°S;
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González et al., 2006; Ritz et al., 2019; Vargas et al., 2011). Uplift has also been likely along the E-W striking Loa reverse fault, which is located to the east of our study area (Fig. 1). Vásquez et al. (2018) postulate that the Cañón del Loa beds were uplifted along this fault before they were covered by Alto Hospicio gravels. A striking feature of comparably young surface modification in the vicinity of the study area is the deeply incised gorge of the modern Río Loa, which has been estimated to have formed since 274 ± 74 ka, based on cosmogenic nuclide surface exposure dating (Ritter et al., 2018a). This age is further supported by a 40Ar/39Ar biotite age of 255 ± 17 ka, obtained from a tephra layer emplaced within an elevated river terrace (Vásquez et al., 2018). It is important to note that the massive incision into the bedrock to a depth of around 600 m below the surface represents the communication of an allochthonous climate signal originating from the elevated areas (High Andes, Precordillera) to the east. This is underlined by the fact 9
Journal Pre-proof that the entire drainage system on the northern and southern rim of the Loa canyon is hanging as it has been bisected by incision of the Río Loa.
2.3 Regional climate characteristics Apart from the general long-term hyperarid climate conditions caused by the circulation patterns as described above (sect. 1), the present-day local climate in the study area is strongly influenced by the abundance of coastal fog (locally known as Camanchaca). Cold and moist air originating from the
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Pacific is trapped underneath an inversion layer caused by the subsidence of warm continental air masses
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leading to the formation of fog (e.g. Weischet, 1966; Weischet, 1996). As the inversion layer in northern
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Chile is on average located at an elevation around 900 m, the Coastal Cordillera with an average height of 1000 m (Hartley et al., 2000) prevents the moist air from migrating inland (Miller, 1976). Thus, low-
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lying areas like the Loa canyon are used as pathways by the advective fog to reach the hinterland (cf.
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Cereceda et al., 2008a).
The present-day occurrence, distribution and precipitation yield from the coastal fog has been studied
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in detail by Cereceda et al. (2008a, b) for a N-S transect spanning over the littoral plains and the Coastal
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Cordillera in northern Chile, located ~70 km north of the study area of this work. The authors found a strong influence of elevation and distance from the coast on the amount of precipitation that they
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measured by using fog collectors. In general, fog precipitation is mostly focused on elevations between 650 and 1200 m and in the period between April to October at the western margin of the Coastal Cordillera (Cereceda et al., 2008a, b). The authors further showed that elevation gradients have a significant impact on dew yields, increasing from 0.44 l vertical m-2 day-1 at 650 m to 7.12 vertical m-2 day-1 at 850 m (base of the fog collector 2 m above the ground; measurement period 2001 to 2004). Further moisture supply can be initiated by special synoptic conditions, which can bring significant amounts of rain to the coastal range of northern Chile (Houston, 2006a). Such events caused floodings after several heavy rainfalls in the 20th century (Ortega et al., 2019; Vargas et al., 2006). Moisture sources originate mainly in the southern or northern Pacific (e.g. Bozkurt et al., 2016; Houston, 2006a; Jordan et al., 2019). Since 2015, several unusual and severe rain events that have caused floodings in 10
Journal Pre-proof the coastal zone of northern Chile have been documented and analyzed in detail (e.g. Bozkurt et al., 2016; Jordan et al., 2015; Jordan et al., 2019; Rondanelli et al., 2019; Wilcox et al., 2016). Spatially associated to the mid-latitude winter rainfall zone, modern rainfall in the coastal range of northern Chile is mainly focused on the months March to September, with intensified precipitation favored by El Niño (Houston, 2006a; Ortega et al., 2019) and prior-El Niño conditions (Vargas et al., 2006). This holds, for example, for the March and August 2015 rain events (Jordan et al., 2019). However, the El Niño Southern Oscillation (ENSO) is believed to be only one of many factors affecting the special synoptic
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conditions that lead to rainfall events in northern Chile (Houston, 2006a; Ortega et al., 2019; Rondanelli
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et al., 2019). Jordan et al. (2019) argue that the rainfall occurring during the June 2017 event was not
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related to El Niño-like conditions but could have occurred due to the advection of cold air masses from the Pacific, as described by Vuille and Ammann (1997). For the second half of the 20th century, Ortega
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et al. (2019) reported a significant increase in extreme rainfall events during the last warm phase of the
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Pacific Decadal Oscillation (PDO) between 26°S and 32°S, emphasizing another possible factor that might affect the precipitation patterns in northern Chile. A further variable characterizing the regional
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climate in northern Chile is the diurnal wind system. A thermal gradient between the cold Pacific and
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the heated Altiplano causes strong near-surface advection towards the hinterland, resulting in strong afternoon westerly winds between the coast and the Western Cordillera (Richter and Schmidt, 2002;
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Rutllant, 2003). During the night, this pattern is reversed with weaker oceanward winds prevailing (Richter and Schmidt, 2002).
3 Material and methods 3.1 Sampling strategy and sample analysis Concentrations of terrestrial cosmogenic nuclides (TCN), especially those of cosmogenic
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Be, have
routinely been measured to quantify catchment-mean erosion rates (e.g. Bierman and Steig, 1996; Binnie et al., 2006; Granger et al., 1996; Placzek et al., 2014; Placzek et al., 2010; Sosa Gonzalez et al., 2016). In this study, we take advantage of the fact that the concentration of a given radionuclide decreases following burial at depths of several meters and shielding from cosmic rays. The half-life of 10Be has been determined to be 1.387 ± 0.01 Myr (Chmeleff et al., 2010; Korschinek et al., 2010) and the half11
Journal Pre-proof life of 26Al is considered to be 705 kyr (Nishiizumi, 2004). Assuming the youngest age of ca. 16 Ma for the deposition of the Cañón del Loa beds and the Alto Hospicio gravels as found by Vásquez et al. (2018), would mean < 0.1% of any in situ-produced 10Be that was present in the gravels prior to their deposition will remain, and approximately half as much
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concentrations of 10Be in the Cañón del Loa beds and the Alto Hospicio gravels are a reflection of their post-depositional exposure history. Steady (hyperarid) climatic conditions over the long-term perspective (sect. 2.3) likely promote long-term steady-state erosion, a key assumption for the
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application of TCN to infer erosion rates (e.g. Dunai, 2010; Granger et al., 1996; Lal, 1991). By
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collecting pebbles, we exclude the possibility of sampling wind-blown material in an environment
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experiencing high wind velocities. Grain size-dependent differences in TCN concentrations as described by Codilean et al. (2014) should not be of great significance in our study area as the incision of the
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sampled catchments took place entirely into alluvium and not into vein-bearing bedrock. In addition to
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that, (sub-)catchment sizes are very small (0.0014 to 0.2619 km2), minimizing transfer times. This circumstance should also reduce the possibility of incorporating significant amounts of pebbles
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originating directly from gully-side slopes adjacent to the sampling sites. Placzek et al. (2014) have
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shown that TCN concentrations from channel sediments in the hyperarid Atacama Desert can be significantly affected by sediment contributions from local slopes, which tend to cause bias by
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increasing the overall nuclide TCN concentration in the channel sediments. However, such a mechanism seems predominantly applicable to grain sizes smaller than those sampled here (Placzek et al., 2014). Furthermore, amalgamating a sufficiently large number of clasts should ensure that the TCN concentration reflects the sediment contribution of the entire drainage system. Though it is not possible to mix the pebbles in proportion to their erosion rates, precluding assumptions of thorough sediment mixing in proportion to the erosion rates of the contributing area (e.g. Bierman and Steig, 1996; Binnie et al., 2006), the production rate differences within the catchments that would cause this to be problematic are minor. We sampled 13 catchments and subcatchments along the E-W topographic gradient illustrated in Figure 1. We aimed at collecting 30 individual quartz-bearing pebbles per catchment within less than 100 m of each other along the main channel of the respective catchment and amalgamated these samples 12
Journal Pre-proof (cf. Repka et al., 1997). We avoided sampling close to slope failures, directly below knickpoints or channel mouths and took care not to collect similar-looking fragments which could have originated from the same pebbles that were broken apart by thermal cracking (cf. Dunai, 2010) or due to salt and/or gypsum expansion, which are widespread phenomena in our study area (Fig. 10). Pebbles were only collected from the bottom and center of the channels and not from lateral stream banks. In the easternmost catchments, this restrictive sampling strategy led to lower yields of suitable pebbles as surface abundance of grain sizes larger than sand decreased markedly. We avoided the catchments that
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drain the elevated hinterland to the north to minimize differences in erosion-controlling variables
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(Fig. 3).
Figure 3: Oblique view of the study area (ArcScene 10.5.1) based on Pleíades 1B satellite multispectral imagery showing the location of the sampled catchments within the E-W transect (framed in red color) and the sampled bedrock knickpoints (red triangles). Channel networks of the catchments sampled are shown in green (including the channels into which the sampled catchments shed), all other large stream networks in blue. In the western portion, sample sites are located so as to reflect local climate conditions, as the drainage is not fed by the more elevated, larger upstream catchment areas to the north (thick black solid line for upstream basin limits and black dashed line indicating general direction of flow towards the south). This is applicable also for catchment BA15-016 in the eastern portion (group E2). All other sampled eastern catchments drain into the trunk streams of catchments that reach the 13
Journal Pre-proof elevated areas to the north, where bedrock is exposed. However, these subcatchments have their channel heads further south and they are entirely hosted in the gravels.
In order to assess the importance of proximity to the local base level on erosion we also decided to sample small N-S transects perpendicular to the Río Loa canyon in the western part of the study area (Group W2: BA15-001 to BA15-003). Inner-catchment variability of erosion rates was investigated by sampling subcatchments within the catchment BA15-004 to the west of the local ridge (Group W1:
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BA15-004 though BA15-007; BA15-010). In the eastern portion, all sampled catchments are N-S
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oriented. BA15-012 and BA16-001 (Group E1) are located in the transition zone towards the
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easternmost catchments that do not show any signs of sub-recent erosion (Group E2: BA15-013, BA15-014, BA15-016). In addition, two prominent bedrock knickpoints (BA15-008, BA15-009) were
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sampled to determine local bedrock erosion rates. The outcropping rocks belong to the Caleta Loa
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metamorphic complex (Csbcl, Figs. 3, 4) and represent the closest bedrock exposures with regard to the W1 catchments. At the knickpoints, the thalweg is offset vertically by ~4 m (BA15-008) and ~8 m
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(BA15-009), respectively.
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The samples were prepared as Accelerator Mass Spectrometry (AMS) targets in the laboratories of the Institute of Geology and Mineralogy, University of Cologne, Germany. Mean clast thicknesses of the
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amalgamated samples varied between 4.1 and 0.6 cm (measured over the shortest axis; ntotal = 335). The majority of clasts (75 %) could be described as “pure quartz” or as having “high (> 80%) quartz content”. Clast colors were highly variable but the majority were greyish without desert varnish, angular to rounded and of low to medium sphericity. The clasts were individually crushed and sieved; either the 250-710 µm (bedrock samples) or 500-710 µm (channel pebbles) grain size fractions were analyzed. The respective fractions from each clast were checked for impurities and hand-picked if necessary. An equal amount of sample material from each clast was amalgamated into one sample. Afterwards, the amalgamated samples and the bedrock samples were etched following the laboratory protocols based on Kohl and Nishiizumi (1992) and quartz purity was checked using an in-house ICP-OES. The etched quartz material was dissolved following spiking with certified, commercially available Be (Scharlab, 1000 mg/l) and Al (Scharlab, 1000 mg/l) carrier solutions. Subsequent preparation of the samples as Be 14
Journal Pre-proof and Al targets for AMS followed the stacked column approach detailed in Binnie et al. (2015), coprecipitating Al and Be hydroxides with Ag following Stone et al. (2004). Chemical blanks were prepared and measured in tandem with the samples. 10Be/9Be and 26Al/27Al values were measured on CologneAMS (Dewald et al., 2013). 10Be/9Be ratios were normalized to the ICN standard dilution series (Nishiizumi et al., 2007) and 26Al/27Al ratios were normalized to the 26Al AMS standards provided by Nishiizumi (2004). The stable Al contents of the dissolved, spiked samples were determined using ICPOES and standard addition (4 aliquots) in tandem with quality control measurements of NIST SRM165a.
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Reagent blank corrected concentrations of 10Be and 26Al were derived following Binnie et al. (2019).
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Uncertainties on the concentrations include the propagated uncertainties in the AMS ratios of the
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samples and blanks together with the estimated standard deviation of the amount of 9Be, or 27Al, the samples contained after spiking.
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The LSDn scaling scheme (Lifton et al., 2014) was chosen to calculate the erosion rates using the online
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calculator described in Balco et al. (2008) and formerly known as the CRONUS-Earth online calculator (version 3; https://hess.ess.washington.edu/math/v3/v3_erosion_in.html). The input used in the online
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calculator is shown in the supplement. We did not correct production rates for uplift as the likely bias
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this would induce is not significant for the erosion rates within the precision of our measurements (see supplement). To estimate the attenuation length of the cosmogenic radiation, and thus approximate the
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averaging time of our erosion rate measurements, we assumed that the clasts have a mean density of 2.7 g cm-3. The relative porosity of gravel comparable to that we find in our study area is likely to be around 30% (Hölting and Coldewey, 2013), resulting in a value of around 1.9 g cm-3 for the density of the conglomerates and ~85 cm for the attenuation path length (cf. Ewing et al., 2006; see supplement). The density of the bedrock we sampled at the two knickpoints was measured to be approximately 2.6 g cm-3, resulting in an attenuation path length of 60 cm. This approach only accounts for spallogenic nuclide production, which makes up the majority of the total 10Be and 26Al production in quartz, especially in relatively slowly eroding catchments (e.g. Dunai, 2010). Topographic shielding was calculated from a 1 m digital terrain model (DTM) using the ArcMap-Toolboxes provided by Li (2013, 2018). The DTM was obtained from a stereo pair of Pleíades 1B images acquired on 04/11/2014 by using the Geomatica 2017 software. 15
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3.2 DTM-based data and climate data analysis For the calculation of catchment-mean shielding, as well as for all other DTM-based data we consistently included only those areas in our analysis which could possibly contribute to the erosion rates that we measure (cf. Fryirs and Brierley, 2012). Thus, we used a 10x10 m moving window in ArcGIS (ver. 10.5.1) to exclude all flat areas that do not appear able to contribute sediments to the fluvial system. These areas were defined as those pixels where the surface does not change more than 1 m in elevation
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within 100 m2 around that pixel.
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Apart from delineating the respective catchment areas, mean elevations and gradients, as well as
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geographic centers, we also traced fault scarps and derived values for some geomorphological parameters from the digital terrain data. Vásquez et al. (2018) stated that fault activity strongly
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influences erosional processes in northern Chile. Thus, hill- and slopeshade maps were used to identify
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possible faults crossing the study area. Along the Loa canyon, outcrops and the level of the pre-incision surface were mapped using the DTM and the Pleíades images. Swath profiles were drawn to assess
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surface tilting. During a field trip conducted in October 2018, aerial imagery was acquired using a DJI
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Spark drone. From these images, we created orthomosaics of all catchments except for those of group E2 (BA15-013, -014, -016) using Agisoft Photoscan software (ver. 1.4.4). The mosaics (3-5 cm
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resolution), the raw aerial pictures that were available for all catchments and the orthorectified and pansharpened Pleíades images (~50 cm resolution) were used to manually extract stream networks and obtain the total surface stream length per catchment. Furthermore, we analyzed our dataset for any correlation between the widely used geomorphological parameters drainage density (Dd), mean relief, mean slope and mean hilltop curvature (CHT), and the erosion rates (e.g. Ahnert, 1970; Granger and Riebe, 2014; Hurst et al., 2013; Hurst et al., 2012). Since some subcatchments contain first-order streams, and thus do not contain any slopes except for the bounding slopes that define the local watershed and thus the subcatchment limit, we included the divides into the hilltop curvature analysis. For the hilltop curvature data, we inverted the elevations of the DTM and extracted the artificially low topographic networks using the standard flow direction/flow accumulation procedures in ArGIS. The lines of this network, representing the ridges and hilltops, were buffered to 2 m so that values from 16
Journal Pre-proof neighboring catchments sharing a ridge could be assigned to both catchments. In catchment BA15-002, we had to remove a group of ridges manually from an area that appeared as a flat surface in the inverted DTM. The buffered ridges were used to clip a plan curvature raster of the study area, from which the values used in this study were extracted. As local uplift and different catchment orientation likely have a strong influence on the geomorphic parameters, special focus was also paid to a comparison of the catchments facing similar conditions, i.e. group W1 vs. W2 and E1 vs. E2. To avoid a possible bias of artificial relief generation in the eastern
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catchments due to the general dip of the pre-incision surface towards the south, we also used the mean
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incision depth of each catchment as surrogate for relief to correlate to our erosion rates. We derived
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these values by measuring elevations along a swath profile oriented parallel to the long axis (in E-W or N-S direction) of the respective catchments. The swath width was defined by the drainage divides of the
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catchment and so varied along the path of the profile. The maximum differences in elevation collected
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along the path length of the swath profile at a 1 m resolution were finally averaged to calculate the mean incision depth.
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As well as the quantification of process rates and geomorphological parameters we also analyzed climate
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data to characterize present-day climatic conditions in our study area in detail. Hoffmeister (2018) and other researchers within the Collaborative Research Centre (CRC) 1211 (Earth – Evolution at the dry
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limit) now provide freely available datasets from multiple meteorological stations installed along E-W running transects in northern Chile (DOI: 10.5880/CRC1211DB.4). As our study area is located directly within one of these transects, climate datasets are available from the littoral plain ~8 km to the west, from the eastern margin of our study area and from the Cerros de Calate ~14 km to the east (stations #11, #12, #13; for location see Figs. 1 and 5). The range of comparable data spans October 2017 to January 2019 (~460 days). Climate stations #12 and #13 were also operating during the June 2017 rain event, which we thus included in our analysis. Furthermore, we generated a small dataset to obtain some insight into the present-day microclimatic conditions prevailing on the western and eastern slopes of the ridge separating the catchments of groups W1 and W2. During the field trip conducted in October 2018, we measured ground temperature and relative humidity using Voltcraft DL-210TH data loggers for a ~7-days period (10/18/2018, 07:00 pm to 10/26/2018, 08:00 am). We stationed 3 pairs of data loggers 17
Journal Pre-proof directly on the surface at ~710 to 720 m altitude (DTM elevation) on the opposing pairs of east- and west-facing ridges. The single units of each pair were placed in a similar distance from the ridge (see supplement). We placed a Kestrel 5500 Weather Meter on the central gypsum-covered peak on top of the ridge between the W1 and W2 catchments, to measure wind speeds 70 cm above the ground (732 m a.s.l).
4 Results
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4.1 Local tectonics
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Mapping of outcrops along the Loa canyon reveals that (half-)graben structures exist in the study area
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as previously described by Vásquez et al. (2018); the bedrock located below the local ridge possibly represents a paleohorst structure (A in Fig. 4). As illustrated at the eastern margin of the elevation profile
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in figure 4, relative uplift is accommodated along the splay fault described by Ritter et al. (2018b) and
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Vásquez et al. (2018). A paleosurface separating the CLB and AHG units below the eastern part of the study area is not embedded horizontally but dips about 1.7° to the east, coinciding with the general
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eastward dip of surfaces that lie to the east of the ridge. By contrast, the pre-incision surface dips about
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1° to the west. By slope and relief analysis, we were also able to trace tectonically induced offsets that likely represent further extensions of a NW-SE running normal fault and/or of the E-W running Loa
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reverse fault as mapped by Vásquez et al. (2018). Geometry of the fault planes appears to be such that they would meet at an area close to the local ridge to the north of groups W1 and W2, if further extrapolated (Fig. 1; supplement). Surface displacement along the extension of the Loa fault can also be inferred from the elevation profile (Fig. 4). In addition to that, a sharp contrast in incision along the E-W transect coincides with a large channel that by itself marks the outer limit of significant erosion to the east. As this pattern is consistent along the channel until it has crossed the elevated areas to the north, we assume that the channel could follow a fault scarp, which has promoted differential fluvial incision and could thus be responsible for the sharp incision contrast between the E1 and E2 catchments (Fig.1, B in Fig. 4; supplement).
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Figure 4: Swath profile (1 m resolution, derived from Pleiades 1B satellite images) and geological cross section (the origin is in the west) across the study area (for extent see Fig. 1). The mean elevation profile
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is indicated by the thick black line (± std. dev. in grey), the thin black line represents the minimum elevation which corresponds to the Loa canyon rim. Colored diamonds denote sampled mean catchment
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locations and -groups; green triangles represent the sampled bedrock knickpoints. The dashed orange
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line represents the linear fit of all calibration points (orange points) placed on the flat surface to the east of the study area, interpolating the possible dip of the pre-incision surface to the west. Geological units and faults were mapped based on Vásquez et al. (2018) and own analysis of remote sensing data [Csbcl – Caleta Loa complex; TrJsl(a) – Sierra de Lagunas strata; Jich – Chuculay fm; Kip – Paiquina complex; CLB – Cañón del Loa beds; AHG – Alto Hospicio gravels]. The bedrock contact and mapped faults indicate the presence of several (half-) graben structures, whereby all faults except for the two westernmost (undetermined) are normal faults. The easternmost fault corresponds to a splay of the Atacama Fault, as described by Ritter et al. (2018b) and Vásquez et al. (2018). The area located between sampling groups W1 and W2, spatially corresponding to the local ridge, might represent a paleohorst structure, indicated by a sequence of serrated bedrock outcrops (Point A). A channel that has been cut by the incision of the Loa river might follow a N-S striking fault scarp (Point B; Fig. 1). Outcrops of a 19
Journal Pre-proof paleosurface separating the AHG and CLB units are indicated by grey points; its eastward average dip below the study area indicated by a pink dashed line.
4.2 Local climate Analysis of the data provided by Hoffmeister (2018) revealed that station #12 recorded the lowest mean air temperature (13.4°C; 15.4°C during daytime 07:30 am to 07:30 pm) and soil surface temperatures (14.5°C; 20.4°C during daytime) of all three stations. Minimum soil surface temperatures at station #12
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decreased below 0°C but maximum temperatures reached almost 40°C during the timespan that was
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analyzed (Fig. 5).
Figure 5. Swath profile (30 m resolution, SRTM data) across the “center transect” of the Collaborative Research Centre (CRC) 1211 (for spatial extent see Fig. 1) including climate stations #11, #12 and #13 (red diamonds; Hoffmeister, 2018). Weather station #12 is located on the eastern margin of the study area of this work (red dashed polygon). A color code has been used to visualize differences in the climate variables air temperature (measured 2 m above the ground; T air), relative humidity (2 m; Φ), surface 20
Journal Pre-proof temperature (Tsurf.), soil temperature (-10 cm; Tsoil), wind direction (2 m; W.dir.), wind speed (2m; vwind), relative amount of time a leaf wetness sensor has been moistened (20 cm, Wet) and precipitation, as measured during the timespan 10/01/2017 to 01/10/2019 (10 min measurement interval). Sea sprayinduced salt encrusting might have influenced the measurements of the leaf wetness sensor at station #11 (marked with *). Red color indicates driest or hottest conditions in relation to the other two climate stations, blue color vice versa and yellow medium values. Values are expressed as average (𝑥̅ ), average during daytime (07:30 am to 07:30 pm; 𝑥̅ day), maximum, minimum and summarized (Total). For
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discussion see text.
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Average values for the relative air humidity of station #12 are the highest when compared to the other
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stations (75.9%; 68.5% during daytime), most likely a response to lower temperatures. Precipitation from rain did not differ significantly between the coastal plain (0.8 mm) and the eastern boundary of our
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study area (0.9 mm) during the analyzed period. In contrast to that, the station #13 at the Cerros de Calate (1148 m a.s.l.) recorded 14.1 mm precipitation, of which a great amount could have originated
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from heavy fog; a similar assumption might be valid for the other stations (cf. McKay et al., 2003). The
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maximum amount from a single event at station #12 was 0.5 mm of precipitation within 10 min on
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2/27/2018 in the afternoon (2.6 mm/10 min at station #13). The occurrence of water precipitating on the surfaces is generally highest close to the sea, where surfaces were moist ~57% of the measurement period (measurement height 20 cm above the ground). However, measurements close to the coast at station #11 have likely been affected by salt deposition on the sensor due to sea spray (D. Hoffmeister and J. Schween 2019, pers. comm.). At station #12, this value should strongly be linked to the occurrence of fog. Accordingly, the vicinity of the study area might have been about ~13% of the measurement period covered by fog. At the Cerros de Calate, this value decreases to ~6%. Similar patterns can be observed for the mean daytime occurrence of fog (#11: 29.9%, #12: 3.3%, #13: 0.8%). At station #12, these values would imply fog abundances during the daytime of about 16 days as well as during 46 nights. The recent rain events in northern Chile have had limited impact on surface modification in our study area. Based on our own observations made during three visits (02/2015, 11/2016, 10/2018) we can state 21
Journal Pre-proof that some surface modification occurred after the 2015 rain events which led to fluvial accumulation of clays in small depressions of the eastern flat surface as well as in channels to the east of the local ridge. On steep slopes, however, we did not find signs of significant downslope material transport, while the rain recorded in February 2018 did not leave significant signs of fluvial activity in the areas we explored. The same can be stated for the June 2017 event, when CRC 1211 climate stations #12 and #13 were already operating. On the eastern edge of the study area (station #12), 1.1 mm precipitated between 11:50 pm (06/06/2017) and 03:20 am (07/06/2017). Further east at station #13, 0.7 mm of precipitation
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was recorded between 9:10 pm and 02:20 am.
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The dataset obtained from the data loggers we placed along the western and eastern slopes of the ridge
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during the one-week-measurement in October 2018 indicates some microclimatic differences (Fig. 6). The east-facing devices recorded a mean daytime temperature (07:30 am to 07:30 pm) of 24.1°C, while
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their west-facing counterparts recorded 19.7°C. A similar surface temperature was measured by the
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CRC 1211 station #12 at the eastern limit of our study area (19.1°C; Hoffmeister, 2018). The difference in temperature across the ridge is most pronounced during the morning hours before noon but persists
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during the afternoon (Fig. 6A).
Figure 6. Variability of temperature (A) and relative humidity (B) at the ground as measured during a one-week-measurement (10/18/2018 to 10/25/2018) along the western (blue lines) and eastern (red lines) flanks of the ridge separating the W1 and W2 catchments. Thick solid lines represent average 22
Journal Pre-proof values (10 minutes interval, 3 data loggers) and thin dashed lines the standard deviation. The dashed black line in (B) indicates the 75% humidity level, above which halite becomes hygroscopic (Gu et al., 2017). This threshold has been exceeded every night during the measurement period.
In addition, relative humidity at the surface dropped from on average 52.5% on the western slopes to 42.5% to the east of the ridge (no comparable data available for CRC 1211 station #12). The daily variability of the relative humidity mirrors the daily course of the air temperature (Fig. 6B). Differences
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between east- and west-facing surfaces are most pronounced during daytime, while relative humidities
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of up to 100% were reached during the early morning hours on both flanks of the ridge. In general, the
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night was characterized by a relative humidity of > 75% during the measurement period. The wind gauge placed on top of the peak recorded a maximum wind speed of 13.7 m s-1 (mean: 4.3 m s-1,
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mean daytime 5.8 m s-1).
4.3 Nuclide concentrations, erosion rates and geomorphological parameters
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The reagent blank corrected concentrations of 10Be and 26Al of the amalgamated samples ranged from
blank subtractions for
10
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1.06 × 105 at g-1 to 3.29 × 106 at g-1 and 8.30 × 105 at g-1 to 1.71 × 107 at g-1, respectively. Respective Be and
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Al constituted < 7.2% and < 6.9% of the total number of atoms
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measured. With regard to the two bedrock samples,
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Be concentrations were 7.86 × 104 at g-1
(BA15-008) and 7.28 × 104 at g-1 (BA15-009). However, the lower
10
Be/9Be rate of BA15-009 has
meant a blank subtraction of 19.7%, arguably too high to confidently rely on this data. Nevertheless, the concentration is similar to that obtained from BA15-008 (6.5% 10Be blank subtraction) and fits to what we would expect for the given topographic setting. We therefore present the results of this sample for the sake of completeness but do not include it in later discussion. In general, the concentrations decrease from east to west along the transect. Plotting the 10Be and 26Al results on a complex exposure diagram indicates steady-state erosional conditions within 2σ for all samples except BA15-001, suggesting that within the precision of the measurements there is no evidence the samples have experienced complex burial histories (see supplement). The erosion rates we derive using the online calculator (Balco et al., 2008) are shown in tables 1 and 2. As the study area is spatially 23
Journal Pre-proof small with low relative relief and production rates between sites are similar, we compare our cosmogenic nuclide results with each other using the internal uncertainties reported by the online calculator (see supplement for external uncertainties).
Table 1. Erosion rates and erosional timescales obtained from the amalgamated samples. ID
Location
Elevation
a
10
26
Be erosion
a
10
Al erosion
-1 b
Be Tave
-1 b
Incision age
c
(Myr)d 8 2 2 2 >1 >1 >1 1 2 <6 < 10 <3 2
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(°) (m a.s.l.) (m Myr ) (m Myr ) (Myr) BA15-001 -21.4214 -69.9783 713 3.2 ± 0.3 4.2 ± 0.6 0.3 BA15-002 -21.4229 -69.9766 696 9.0 ± 0.7 10.5 ± 1.4 < 0.1 BA15-003 -21.4242 -69.9767 695 13.4 ± 1.1 15.6 ± 2.7 < 0.1 BA15-004 -21.4243 -69.9812 680 28.8 ± 3.3 27.2 ± 10.6 < 0.1 BA15-005 -21.4240 -69.9830 670 26.5 ± 2.4 26.9 ± 4.0 < 0.1 BA15-006 -21.4257 -69.9825 664 15.0 ± 1.5 17.6 ± 4.2 < 0.1 BA15-007 -21.4237 -69.9820 677 20.7 ± 1.8 21.3 ± 3.8 < 0.1 BA15-010 -21.4227 -69.9834 677 39.6 ± 4.7 34.8 ± 6.9 < 0.1 BA15-012 -21.4141 -69.9725 737 3.7 ± 0.3 4.1 ± 0.7 0.2 BA15-013 -21.4125 -69.9654 734 1.1 ± 0.1 1.1 ± 0.4 0.8 BA15-014 -21.4085 -69.9661 757 1.1 ± 0.1 1.3 ± 0.4 0.8 BA15-016 -21.4146 -69.9671 732 1.8 ± 0.3 2.3 ± 0.7 0.5 BA16-001 -21.4200 -69.9715 704 4.5 ± 0.7 4.9 ± 1.1 0.2 a Catchment-averaged. b LSDn scaling scheme used to derive production rates. All errors are 2σ (internal uncertainty). c For z = 85 cm, assuming steady state erosion on the order of 3-5 × Tave (Lal, 1991). d Approximation based on the mean incision depth and rate of incision.
The catchment-wide erosion rates obtained from the amalgamated samples reflect the overall visible
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incision patterns along the E-W transect (Fig. 7). Accordingly, the lowest 10Be and 26Al rates of around 1-2 m Myr-1, are found in the easternmost catchments (group E2). By contrast, the westernmost catchments erode at up to 39.6 ± 4.7 m Myr-1 (BA15-010), implying an east-west increase of erosion by one order of magnitude within the 2500 m transect. Similar rates are obtained from the two bedrock knickpoints to the west of the W1 catchments (Tab. 2).
Table 2. Erosion rates and erosional timescales obtained from the bedrock samples. ID BA15-008
Location (°) -21.4255
-69.9846
Elevation (m a.s.l.) 607
10
Be erosion (m Myr -1)a 36.3 c
±
5.1
Be Tave (Myr)b < 0.1
c
BA15-009 -21.4258 -69.9857 582 39.5 ± 11.2 a LSDn scaling scheme used to derive production rates. All errors are 2σ (internal uncertainty). b For z = 60 cm, assuming steady state erosion on the order of 3-5 × Tave (Lal, 1991). c Reduced reliability because of large blank subtraction, see text for discussion.
24
10
< 0.1c
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The timescales over which the erosion rates integrate also differ by one order of magnitude between fast and slowly eroding catchments, being ~20 kyr in catchment BA15-010 and about 0.8 Myr for BA15-014. The bedrock erosion integrates over ~20 kyr. The averaging time of the erosion rate measurements provides an estimate of the period over which the measurements themselves are valid, being the time to erode an average of one attenuation length, or in this case a depth of ~85 cm, assuming steady state erosion over time periods on the order of 3-5 × Tave (Lal, 1991). Comparing erosion rates
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integrated over such variable time periods presents a challenge in the assertion of driving factors, which
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might exert control over differing timescales (Schumm and Lichty, 1965). However, as the basins are
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located relatively close to each other, the large-scale environmental factors such as the uplift of the Coastal Cordillera or synoptic climate conditions (sect. 1, 2.2, 2.3) are likely to be very similar, including
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the long-term perspective. In addition, the topography of the western catchments has clearly required
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erosion over timescales much longer than the averaging time of the cosmogenic nuclide measurements. Assuming the erosion rates we derive in the west have been relatively constant over the averaging times
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of the eastern catchments, their comparison is tenable.
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Figure 7. Catchment-wide and bedrock 10Be erosion rates (in m Myr-1) and 10Be erosional timescales
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Tave (in brackets; Myr), illustrated on a Pleíades 1B pansharpened multispectral satellite image. The data indicate an approximately 40-fold E-W increase of erosion along a distance of ~2500 m. A N-S
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increase is also observed in groups W2, E1 and E2 (see also Fig. 8). Bedrock contact (Cañón del Loa
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beds – CLB and Caleta Loa complex – Csbcl) and limit of Alto Hospicio gravels (AHG) are represented by white dashed lines. The dashed red lines indicate tectonic displacements as inferred from the DTM analysis and observations made in the field. See text for discussion of the reliability of sample BA15-009 (*).
Apart from the general decrease in W-E direction, a downstream increase of erosion is also observed (Fig 8). Downstream of catchment BA15-004, the bedrock at knickpoint BA15-008 erodes at 36.3 ± 5.1 m Myr1. In catchment BA15-004, eroding at 28.8 ± 3.3 m Myr-1, subcatchment BA15-005 erodes faster (26.5 ± 2.4 m Myr1) than subcatchment BA15-007 (20.7 ± 1.8 m Myr-1), which is located further upstream (Fig. 8A). Both subcatchments are southeast-facing. Subcatchment BA15-006, situated directly opposite to BA15-005 (northwest-facing), erodes at 15.0 ± 1.5 m Myr-1. 26
Journal Pre-proof The east-facing catchments of group W2 are connected by a N-S running channel that drains into the Río Loa canyon (Figs. 3, 8B). The calculated catchment-wide erosion rates decrease in a northerly direction from 13.4 ± 1.1 m Myr-1 (BA15-003) to 9.0 ± 0.7 m Myr-1 (BA15-002) and 3.2 ± 0.3 m Myr-1 (BA15-001). The northernmost catchment of this triplet erodes at similar rates as the western N-S oriented catchments BA16-001 (4.5 ± 0.7 m Myr-1) and BA15-012 (3.7 ± 0.3 m Myr-1) with the former
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being located closer to the Río Loa canyon.
Figure 8. Inner-catchment E-W and N-S increase of erosion rates. Shown are normalized channel
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profiles (Demoulin, 1998) of the master drainage channel of BA15-001 through BA15-003 (group W2; B) and BA15-005 through BA15-007 (group W1; A), which is for the latter case the main channel of catchment BA15-004. The channel profiles extend from the respective channel head to the northern rim of the Loa canyon. Points of confluence and the sampling point of BA15-004 are illustrated by red dots; the locations of knickpoint bedrock samples are indicated by red triangles. The profile in (B) indicates that N-S running channels are influenced by the base level drop caused by the Río Loa incision. Thus, the N-S increase of erosion in the east-facing catchments can possibly be linked to upstream knickpoint migration and the upstream communication of the erosional signal perpendicular to the Río Loa. The inner-catchment variation of southeast- and northwest-facing subcatchment erosion in BA15-004 seems to be affected by other factors as the upstream migration of the erosion signal is impeded by the bedrock knickpoints located downstream of BA15-004 (A). 27
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Although the erosion rates clearly reflect the overall drainage patterns and the geomorphology of the study area, they do not correlate with the respective drainage densities. The majority of catchments show values for Dd between 52 and 67 km km-2 (Tab. 3). Not included into this group are the subcatchments BA15-005 through BA15-007. While the two southeast-facing subcatchments show values above 70 km km-2 for Dd, the opposite case applies to the northwest-facing subcatchment BA15-006 (~42 km km-2). The second-lowest value is reached by the slow-eroding catchment BA15-001. It is also
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both show similar stream lengths per area (~59 km km-2).
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noteworthy that the two fast-eroding catchments to the west of the ridge (BA15-004 and BA15-010)
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Journal Pre-proof Figure 9. Catchment-wide erosion rates plotted against local relief (A), mean incision (B), average catchment slope angle (C), and average catchment hilltop curvature (D). Colored numbers denote the coefficients of determination (R2) of the respective linear correlations. Such correlations (above 0.6) exist for all geomorphological parameters except for drainage densities. The westernmost catchments in (C) seem to have reached a threshold slope angle of ~30°; thus, a correlation is not made for this group. The parameters were derived from the eroding areas that contribute to the rates we measured
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(sect. 3.2). For further discussion, see text.
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A linear correlation of the erosion rates can be observed when related to local relief (R2 = 0.61; Fig. 9).
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However, while the correlation holds if only the western, i.e. catchments of groups W1 and W2, are considered (R 2 = 0.67), the correlation is very weak for the eastern catchments of groups E1 and E2
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(R2 = 0.04). For mean incision depth, taken as surrogate for local relief, the correlation of the entire
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dataset is slightly higher than for the relief (R2 = 0.65). However, if only those catchments are considered that are bounded by stable ridges, including the gypsum-covered ridge between groups W1 and W2, the
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linear correlation increases to R2 = 0.87. The highest values for mean incision (~50 m) are reached for
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the westernmost catchments BA15-004 and BA15-010, indicating a five-fold increase of incision when
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compared to the E1 and E2 catchments (Tab. 3).
Table 3. Values of the geomorphological parameters obtained for the respective catchments. Size
Mean elevation
Dd
Mean slope
Mean CHT
(m2)a
(m a.s.l.)a
(km km-2)a
(m)a
(m)a
(°)a
(m-1)a
BA15-001
107646
713
52.28
66.43
24.82
24.62
0.11
BA15-002
36023
696
66.32
60.88
19.76
25.62
0.12
BA15-003
52562
695
60.55
79.90
23.33
25.77
0.12
BA15-004
162288
680
58.67
113.95
49.22
29.82
0.15
BA15-005
11373
670
79.43
70.50
14.42
31.65
0.14
BA15-006
7711
664
41.94
51.33
14.84
29.05
0.10
BA15-007
20999
677
70.91
75.42
22.18
31.41
0.16
BA15-010
154232
677
58.97
142.00
48.98
30.89
0.16
BA15-012
91605
737
56.07
73.16
7.70
16.60
0.07
BA15-013
76385
734
64.47
41.66
6.43
10.74
0.06
BA15-014
227647
757
56.94
76.96
9.95
12.54
0.06
BA15-016
153007
732
61.75
61.55
5.48
10.28
0.05
ID
Local relief Mean incision
29
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a
BA16-001 50500 Eroding areas.
704
59.55
41.52
9.10
17.79
0.09
A further separate analysis is also reasonable for the correlation of catchment-mean slopes and erosion, where linear correlation is especially strong for the eastern catchments (R2 = 0.84; Fig. 9). It is, however, striking that the mean slopes of the fast-eroding catchments of group W1 reach all values of around 30°. For mean hilltop curvature and erosion rates, linear correlation is high for the entire dataset (R2 = 0.76), with low values for CHT (~0.06-0.09 m-1) obtained for the eastern catchments. In the western portion,
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BA15-001 to -003 and BA15-006 reach mean hilltop convexities of about 0.10 to 0.12 m-1 (Tab. 3). The
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remaining catchments (BA15-004, BA15-005, BA15-007 and BA15-010) show values above 0.14 m-1.
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5 Discussion
5.1 Erosional stages of the study area: E-W and N-S contrasts
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The correlation patterns between the geomorphological parameters and the catchment-wide erosion rates point to the prevalence of two different erosional regimes along the E-W transect. The eastern
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catchments of groups E1 and E2 erode at lowest rates and mostly reach low values for the respective
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parameters (Tabs. 1 and 3). Among the parameters that are derived, drainage density 𝐷𝑑 , defined as the
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quotient of total stream length and catchment area, is commonly used to assess the degree of landscape dissection (e.g. Bierman and Montgomery, 2014; Collins and Bras, 2010). In arid landscapes without a soil-protecting vegetation cover, drainage density is believed to be strongly linked to the availability and variability of erosive precipitation inducing fluvial processes (Collins and Bras, 2010; Moglen et al., 1998). It has been shown that power law relationships between Dd and erosion exist in soil-mantled landscapes (Clubb et al., 2016). However, it appears that observed drainage densities in our study area do not reflect the different magnitude of erosion along the E-W transect. In part, this was caused by the fact that we could only extract two-dimensional stream networks due to poor DTM qualities generated from the drone pictures. In addition, Howard (1997) showed that values for Dd (1-D/2-D) decrease as incision increases in detachment-limited badlands dominated by threshold slopes. The drainage densities derived from our study area do indicate, however, that channel incision is more pronounced on the southeast-facing slopes than on their northwest-facing counterparts in catchment BA15-004. 30
Journal Pre-proof In theory, fluvial incision increases the local relief and hillslope gradients, which have shown to correlate linearly to erosion in low-relief and low-slope catchments and explained the response to a drop of the local base level (Ahnert, 1970; Montgomery and Brandon, 2002). Erosional steady-state can be assigned to those areas where long-term erosion rates are rather uniform and decoupled from relief or slope (Burbank, 2002). The lack of a simple linear correlation between erosion and local relief in the eastern catchments is biased by the general dip of the pre-incision surface towards the south, creating relief due to regional tectonic warping and pre-Middle Miocene alluvium aggradation, rather than being an
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indicator for local channel incision. As catchment BA16-001 has the smallest N-S extent, the dip of the
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surface causes an elevation gradient of about 10-20 m, while almost the entire relief of catchment
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BA15-016 (about 60 m) is created by the surface dip. This bias is to some extent reduced by applying mean incision depth as a surrogate for local relief. Using this parameter, it is possible to compare only
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those catchments that are incising into a stable, diffusive surface, which mostly applies to the E1, E2,
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and with some limitations also to the W1 and W2 catchments. To relate western and eastern catchments in our study area using the mean incision depth, it is an important prerequisite that large parts of the
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bounding ridges – defining the watershed of the respective catchments – have not been worn down
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significantly for timescales exceeding Tave. Based on DTM analysis and observations made in the field we believe this holds for all but the subcatchments BA15-005, BA15-006 and BA15-007. The latter are
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incising into a side slope of BA15-004, and the ridges defining their watersheds have been significantly lowered as a result of sediment removal along the main channel of BA15-004. Thus, the incision depth is significantly affected by the downwearing of the slopes and these subcatchments should be excluded from the correlation in such a small-scale analysis. However, a further constraint appears when this parameter is applied to the eastern catchments, as presumably tectonic movement has tilted the surfaces to the east (Fig. 4). This especially affects the E2 catchments, since incision is low and their E-W extents significant, artificially increasing the incision depth. Nevertheless, because averaging reduces this effect to some extent, they are included into the correlation. In the catchments of group E1, the eastward tilt does not affect the mean incision significantly, as E-W extents of the catchments are comparably low and the incision is generally deeper.
31
Journal Pre-proof Apart from these constraints, the high correlation of mean incision depth with the erosion rates indicates that the study area is adjusting to a drop of the erosional base level. This finding is supported by the linear correlation of hilltop curvature CHT with erosion. As slope and local relief are only reflecting the magnitude of erosional processes until a certain threshold is reached, beyond which a further increase is not possible (e.g. Binnie et al., 2007; Burbank, 2002; Montgomery and Brandon, 2002; von Blanckenburg, 2005), hilltop curvature, as a geomorphic parameter, represents the final stage of slope adjustment to adjacent channel incision in eroding landscapes (Hurst et al., 2012). This parameter is
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applicable to the entire dataset, as it records the slope morphology beyond the limitations of mean slope
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or relief (Hurst et al., 2013; Hurst et al., 2012). Accordingly, a linear correlation between erosion and
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CHT is observed. It should be noted that the kind of limitation of the methodology we applied and which affects the parameters local relief and mean incision could also affect the mean slope and CHT. Headward
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channel erosion in a catchment will impact on the bounding ridge between it and the neighboring
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catchment, affecting the ridge curvature which we included in our analysis (see supplement). However, as those catchments that share a drainage divide erode at similar rates and show similar values for the
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geomorphologic parameters, we assume that this methodological limitation does not change the results
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significantly. Thus, for the catchments of groups E1 and E2 it can be stated that adjustment processes are taking place very slowly in a transport-limited erosional regime (Binnie et al., 2007). Taking into
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account the possible relief of 400-600 m down to the Río Loa, the catchments reflect an early stage of erosional response to base level change. By contrast, the western catchments of group W1 and W2 show relatively high values for erosion and the respective geomorphic parameters. The W1 catchments most likely have reached a threshold angle, indicating that erosion has been decoupled from slope and is primarily governed by mass wasting (Burbank, 2002), a feature that is characteristic for a badland-like topography (Howard, 2009). So far, cosmogenic nuclide based erosion rate estimates have been somewhat spatially biased towards active mountainous regions (Portenga and Bierman, 2011), where the technique has been used to provide empirical evidence for theoretical geomorphological concepts, such as the attainment of threshold topography (e.g. Binnie et al., 2007; DiBiase et al., 2010; Ouimet et al., 2009). The notion of threshold topography, when applied at the broad scale suggests that, on average, hillslope gradients of mountains 32
Journal Pre-proof maintain some relationship with erosion rate that decouples as slopes reach their threshold angles of repose. Hillslopes that attain threshold gradients cannot steepen further and then erode by landsliding in proportion to the rate of incision of the valley trunk stream (Burbank, 2002; Burbank et al., 1996; Montgomery and Brandon, 2002). This has led to predictions of the emergence of ‘threshold topography’ after a particular rate of average erosion is reached, e.g. > ~200 m Myr-1 for catchments > 25 km2 at the eastern margin of the Tibetan Plateau (Ouimet et al., 2009); though the erosion rate and hillslope angle at which threshold topographic processes operate will also be a function of material
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strength (Palumbo et al., 2009; Schmidt and Montgomery, 1995) and climate (Gabet et al., 2004).
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This notion of ‘threshold topography’, or the threshold angles of hillslopes, was initially proposed to
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explain characteristic hillslope angles in soil studies concerned with local-scale slope processes operating on valley side slopes (e.g. Carson and Petley, 1970; see Montgomery, 2001). Hillslope
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modelling studies subsequently incorporated a critical slope gradient value to derive non-linear sediment
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transport laws, assuming processes remain dominantly transport-limited (e.g. Roering et al., 1999; 2007). Howard (1997) applied such modelling to a badland landscape, showing that parallel retreat of
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slopes follows a rapid incision, but also noting the lack of process rate information prevents more
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conclusive interpretations. Here we find that the notion of a limiting, or threshold, slope angle is apparent in our erosion rate data from a badland landscape, despite the relatively slow rates of erosion we
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measure. Nevertheless, the lack of developed gypsum crust cover and the evidence for mass wasting deposits in the channels at our fieldsite would suggest detachment-limited conditions prevail at this location and the reason for the attainment of threshold slope angles under such a low-energy environment likely relates to the lower material strength of the conglomeritic lithology. However, a linear correlation between erosion and local relief does exist in the western catchments, which can be explained by the fact that all catchments of group W1 and W2 except for BA15-005 and -007 have their basin head bounded by the local ridge or a plateau (BA15-006). The gypsum crust cover on the peaks of the ridge and its morphology indicate that it has been providing a long-term stable upper limit of the local relief, which has possibly not been worn down significantly for timescales exceeding Tave. Thus, local relief is only dependent on the magnitude of incision in the main channel of the respective catchments, which is governed by the erosion rates and in this special case a more 33
Journal Pre-proof meaningful parameter than mean slope. A similar argument is valid for the mean incision depth. However, as it is decoupled from the slope angle, erosion in the westernmost catchments is in a more developed stage and detachment-limited. This finding is corroborated by the high hilltop curvatures in these catchments. It is very likely that erosion rates would have been higher if the drainage had not incised into the bedrock a bit further downstream of the W1 catchments, slowing down the upstream migration of knickpoints (Fig. 8). These knickpoints act as local base levels for erosional processes further upstream, such that
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the bedrock erosion rate at BA15-008 (36.3 ± 5.1 m Myr-1) sets the pace of erosion in catchment
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BA15-004 (catchment-mean erosion rate of 28.8 ± 3.3 m Myr-1), which is located ~100 m further
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upstream of the knickpoint. If the bedrock knickpoints would not have acted as local erosional base levels, the ridge separating W1 and W2 catchments might have been already eroded away and the W2
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catchments captured. Elsewhere, rapid incision and subsequent base level stability has been found to be
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an important factor for the formation and preservation of badland topography (Howard, 1997). There is no indication that the W1 streams are significantly capturing the W2 watersheds, although its general
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Whipple, 2018).
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morphology indicates than an eastward shifting of the ridge probably has been initiated (cf. Forte and
Apart from the W-E contrast of erosion, manifested in detachment-limited vs. transport-limited
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erosional regimes, the results also highlight the strong dependency of erosion on the spatial distance to the Loa canyon. In all catchments of groups W2, E1 and E2 the northern catchments erode at lower rates than those that are located closer to the canyon. This is again most likely due to upstream knickpoint and/or knickzone migration perpendicular to the Río Loa (Fig. 8). However, given the low erosion rates in group E2, the communication of the erosional signal of the latest Río Loa incision has certainly not reached these catchments yet. The fact that the catchments BA15-013, -014 and those of group E1 are draining into larger channels, which in turn drain the elevated areas to the north, seems not to have any significant impact on the erosion rates. As we did not observe knickpoints on the confluence points of the analyzed catchments with the respective master channels, we argue that the W-E decrease in erosion also affects the elevated areas to the north.
34
Journal Pre-proof 5.2 Mechanisms causing the E-W erosion gradient across the ridge Special attention should be paid to a catchment that seems to contradict the general trends observed for our dataset. Subcatchment BA15-006 shows the lowest erosion rates of all W1 catchments (15.0 ± 1.5 m Myr-1), although it is located opposite to subcatchment BA15-005, eroding at 26.5 ± 2.4 m Myr-1. Furthermore, its drainage density is the lowest of the entire dataset. This could be, however, attributed to the high sensitivity of that parameter to minor changes in the stream network in small catchments. That such catchments sometimes underestimate Dd can be inferred from the datasets
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presented by Clubb et al. (2016). By contrast, the other two small subcatchments BA15-005 and
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BA15-007 have the highest drainage densities of all catchments. In addition to that, all other
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geomorphological parameters for BA15-006 are the smallest within group W1 (Tab. 3). Thus, the anomalous values obtained for this particular catchment are most likely caused by special conditions
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that can be found in its catchment head. This area can be described as a small, uneven plateau, covered
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by a gypsum-rich layer of varying thickness of 0 to > 30 cm and/or millimeter to centimeter-thick crusts forming centimeter-scale polygons (Fig. 2F). Gypsum crust, of the type observed in the eastern
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catchments, is almost absent, and the surface further covered by clasts predominantly of centimeter to
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decimeter size. Despite of the missing gypsum crust on this plateau (which has been excluded from the derivation of the geomorphic parameters, sect. 3.2), the surface does not show signs of significant fluvial
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erosion towards BA15-006. The unevenness of the terrain seems to be related to a weathering and/or deflation of former gypsum polygons, which can still be traced on the surface. The shallow subsurface is salt-cemented (mostly halite), which has been observed to be a wide-spread phenomenon in the study area (Fig. 10). At some places on top of the plateau, salt-cemented sediment forms a competent cap which is pervaded by few centimeter-wide and > 1 m deep cracks. At such a feature we found a small salt stalactite with the morphology of a small droplet (~1.5 cm along the long axis; Fig. 10F), indicating long-term, slow but steady downward (per descensum sensu Watson, 1985) transport of salts, possibly due to diurnal fog-precipitation. The salt-rich layer is being eroded at the channel heads of the subcatchments incising into the plateau, most likely impeding the erosion to keep the pace of the overall erosion in BA15-004.
35
Journal Pre-proof Moving across the ridge that separates the W1 and W2 catchments, the salt-cemented sediments are additionally covered by a decimeter-thick gypsum crust (costra, sensu Ericksen, 1981). Diurnal fogmoisture cannot permeate this gypsum cover to reach the halite (Ericksen, 1981) and leach it from the uppermost sediment layers as we observed it in the western catchments. The gypsum crust itself is a formidable protection of the surface against individual, presumably brief, rain events. It has a high physical cohesion, and gypsum, while being semi-soluble in water (Goudie and Viles, 2015), is kinetically sluggish to dissolve (Warren, 2016). Lichens are occasionally macroscopically visible but,
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however, are rare. The physical cohesion of the gypsum crust in the study is therefore presumably
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dominated by abiotic factors. Commensurate with its physical properties, the gypsum crust-covered
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areas show few signs of fluvial erosion. Visible traces of run-off erosion, such as accumulations of detrital material, are mostly limited to the boundaries of polygonal structures in the crust. Gypsum-
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covered hillslopes appear to erode from the bottom, in response to channel incision/back cutting.
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Eroding slopes have a pronounced break in slope where the crust terminates (the lower slopes are commonly at the angle of repose). These features are best preserved in the transitional group of
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catchments just east of the ridge (W2; Fig. 7 and supplement).
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With key physical parameters being similar across the study area (mean elevation, substrate, presumably long-term mean annual precipitation from rain) the presence of gypsum crust and the ‘health’ of the
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halite-cement in the top-layer of the sediment appears to be a good predictor of the order of magnitude of erosion rates; 15-40 m Myr-1 in areas without crust and leached halite, 1-2 m Myr-1 with an intact gypsum crust cover (Fig. 7).
36
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Figure 10. Pictures taken in the field illustrating some characteristic features of the surfaces in the study area. Salt outcrops along the northern slopes of the Loa canyon indicate the widespread presence of subsurface salt-cementation (white arrow in A; red arrow pointing at the local ridge separating W1 and W2 catchments). In the eastern parts, the salt-cemented gravels (III in B) are covered by gypsum-rich layers (II in B) and gypsum crust (I in B). Note that the crust is cohesive enough to form a topographic hanging edge (Picture taken in BA15-003). Thick, encrusted gypsum cover (costra sensu Ericksen, 1981) is abundant on the crest of the main ridge (C), indicating long-term erosional quiescence and surface 37
Journal Pre-proof stability. The high abundance of disintegrated clasts on the surfaces indicates long-term surface and/or subsurface stability (D). This statement is further supported by a fossil-bearing limestone, whose exposed east-facing surface has been differentially weathered, exposing shell fragments (white arrow in E). The weathered surfaces are restricted to the part of the rock sticking out of the crust (to the right of dashed white line); the east-facing area is not affected by weathering. The downward migration of salts formed a small (~1.5 cm) stalactite in BA15-006 (F).
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The factors that lead to the formation and/or preservation of the gypsum crust are therefore crucial in
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forming the landscape as we find it at the present day. Given the subdued topography and small lateral
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extent of the study area, microclimatic conditions are probably key factors in determining the spatial presence or absence of gypsum crusts. Long-term stable and arid to hyperarid conditions must prevail
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for a gypsum layer of a significant thickness to form as the precipitation of gypsum crystals is prevented
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in moist soils (Heine and Walter, 1996; Retallack and Huang, 2010). The coastal fog has been identified as an important source for the transfer of pedogenic salts like Ca and S from the Pacific Ocean and their
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subsequent per descensum deposition in the Coastal Cordillera at altitudes below 1200 m a.s.l. (Rech et
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al., 2003). As minor abundances of gypsum crust to the west of the local ridge on top of the plateau above BA15-006 and on a small slope within BA15-004 indicate (Fig. 2), gypsum deposition must have
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occurred in the entire study area at some point but was later removed to the west of the ridge. Microclimatically, there are three important differences between the east- and west-facing drainage bounded by the ridge (Fig. 11). Firstly, the W1 catchments face the strong afternoon inland winds (sect. 2.3) which can exceed 13 m s-1 as measured by the wind gauge placed on the peak between the W1 and W2 catchments. Wind exposure could be a limiting factor for the formation and preservation of gypsum in our study area, as extreme wind scour could impede its accumulation (Miller, 2008) or erode these surfaces (Eckardt and Spiro, 1999).
38
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Figure 11. Sketch illustrating possible mechanisms and microclimatic conditions that could be
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responsible for the E-W contrast of erosion in along the main ridge. The western slopes face strong
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daytime winds, cooling down the surface and possibly impeding atmospheric gypsum deposition. Furthermore, the steep slopes could act as a kind of natural fog collector, promoting long-term
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downward migration of salts which could occasionally release clasts, initiating downward chain
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reactions of clasts following gravity. The crust of the east-facing slopes remains comparably unaffected. However, the fact that chuca is mostly absent also points to long-term quiescence of gypsum deposition.
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During a rare rain event, the cemented subsurface of the west-facing slopes, although likely stabilizing the slopes, might promote surface runoff, while to the east the gypsum cover promotes infiltration.
Secondly, almost the entire moisture supply reaching the study area originates from the Pacific Ocean (sect. 2.3) and it is unlikely that the present-day advection patterns differ significantly from those prevailing over the past millions of years (sect. 1). While it is reasonable to assume that the rare rain events are spatially uniform across the study area (due to similar and low elevation as compared to cloud base of rain-clouds), diurnal moisture deposition from fog (i.e. ground-hugging clouds) will not be uniform. Fog-moisture preferentially deposits on wind-facing surfaces, in our case the west-facing slopes, with accompanying wind/fog-shadow effects. Furthermore, dew yields are highly sensitive to altitude (see sect. 2.3). At the Alto Patache Climate station, located ~70 km to the north of our study 39
Journal Pre-proof area at an elevation of 850 m a.s.l., Cereceda et al. (2008a) measured an average precipitation of 7.0 l m-2 day-1 from fog between 1998 and 2005. The yield of fog water is highly focused on the winter/spring months, peaking in September (exceeding 20 l m-2 day-1), while lowest yields are recorded for February and March (Cereceda et al., 2008a). We hypothesize that the west-facing slopes along the moderate ridge that separates areas of different erosional regime act as a fog collector. Significant aspect-related contrasts in fog precipitation are supported by observations in catchment BA15-002 to the east of the main ridge. Situated on a ~25° slope, 2 m below the hilltop, we found a 30x30x15 cm
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fossil-bearing limestone sticking out of the gypsum crust. The west-facing side of the rock was strongly
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weathered, with shell fragments protruding about 2 cm from the weathered rock surface. The eastern
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side lacked weathering features, indicating long-term surface stability and that moisture deposition is sufficient to differentially dissolve limestone. Fog-precipitation is evidently sufficient to mobilize
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surface halite (halite stalactite, Fig. 10F). It has been shown that in arid desert environments, the highly
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soluble halite (which is hygroscopic at relative humidities > 75%, Gu et al., 2017; commonly exceeded in our study area, see also sect. 4.2) is segregated from the sulphates and washed down during periods
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of higher moisture supply (Eckardt and Spiro, 1999), cementing the subsurface. Such a mechanism has
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already been reported from the hyperarid deserts of Tunisia and Namibia, resulting in the formation of two-tiered crusts (Watson, 1985; Fig. 11). Apart from stabilizing the slope (cf. Watson, 1985; Wilcox
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et al., 2016), the cementation promotes surface runoff during rain events (Yang et al., 2016). Thus, on hillslopes on which halite is remobilized into the shallow subsurface (catchment group W1; Fig. 7), erosion of the then uncemented surface material is facilitated by its lack of cohesion and increased surface runoff. Thirdly, fog-moisture is exposed to sunlight after sunrise to the east of the ridge while in west-facing catchments the surfaces are shaded during the early morning hours. Water cannot infiltrate as deep as on the west-facing surfaces, where evaporation rates pick up later. Thus, on the long-term perspective, the east-facing slopes tend to be warmer/drier, likely promoting the formation of near-surface pedogenic alabastrine gypsum through increased evaporation (e.g. Eckardt et al., 2001; Watson, 1985). Taken together, the erosional gradient that we can observe today might have been promoted by a weakly developed surface crust to the west of the ridge, which is likely also subject to wind erosion. Especially 40
Journal Pre-proof on the flat areas in the eastern portion of the study area, the crust-covered surfaces and gypsum-rich subsurface layers likely soak up the majority of precipitation during a rare rain event (Ericksen, 1981). Such a behavior has already been observed for flat to gently dipping surfaces in northern Chile during the 2015 rain event (Jordan et al., 2015). Davis et al. (2010) have shown that a single rainfall event of 3.55 mm, as well as a rainfall of 7.14 mm cumulated over 5 consecutive days, did not produce any surface runoff or ponding on a flat gypsum-covered surface in the Yungay area. By contrast, the possibly weakly developed and wind-eroded crust in the west-facing catchments could have been more easily
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fluvially eroded, whereby overland flow is likely being accelerated when the erosion has reached the
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subsurface due to its strong cementation reducing the infiltration rates (Yang et al., 2016). Regularly-
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occurring moistening of the exposed subsurface could lead to a slow but steady downwash of the halite and downward shift of the focus of cementation. On threshold slopes, this could occasionally release
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previously cemented clasts which would move downslope following gravitation, possibly hitting other
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clasts, and initiating a downslope chain reaction. However, the subsurface salt layers do stabilize the slopes (cf. Wilcox et al., 2016) and slow down their incision (BA15-006). Incision eventually
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5.3 Evolution of drainage
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-006 (Figs. 2, 10, 11).
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undermines any remnants of intact crust covering the hilltops, as observed in catchments BA15-003 and
The analysis of tectonic and geological features (sect. 4.1) reveals that N-S striking (half-)graben structures exist in our study area, covered by the alluvial gravels. The bedrock outcrops as mapped along the Loa canyon indicate that the area around the ridge that is influencing the erosional behavior is placed on a paleohorst structure. This structure possibly underwent relative uplift after it was covered by the alluvial sediments, as gravels belonging to the CLB are found at higher elevations than the adjacent younger AHG pre-incision surface, thereby forming the present-day ridge separating the western catchments (Fig. 4). Tectonic movement after the deposition of the CLB and before the aggradation of the Alto Hospicio gravels has been reported for the Loa fault (Vásquez et al., 2018), which could have led to a canyon-parallel uplift of the study area on the order of tens of meters or less. However, a more important role could have been played by N-S oriented, meter-scaled tectonic vertical displacements 41
Journal Pre-proof and tilting in the eastern section of the study area, associated with the assumed N-S fault scarp (B in Fig. 4) and/or along hypothetically existing N-S faults that have produced the serrated profile shape of the paleohorst (A in Fig. 4). This view is supported by a ~300 m long, horizontally embedded layer of well-recognizable gravels, which are exposed along the Loa canyon below the E1 catchments and contact the eastern flank of the assumed paleo-horst. However, the easternmost ~20 m of this strata are dislocated downwards by about 90 m, roughly below the assumed N-S scarp between the E1 and E2 catchments. Further indications of fault movement, which also occurred after the deposition of the AHG
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unit, can be inferred from slope analysis and are also traceable in the field (Fig. 1; supplement). We
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found no indication that post-Miocene tectonic movement has influenced the overall drainage pattern
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and channel evolution. The tectonic influence on drainage, i.e. the tilting and uplift of parts of an Early Miocene depositional surface, apparently occurred mostly prior to the incision of the catchments
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investigated.
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It is likely that the N-S running eastern catchments of group E2 represent initial drainage conditions, incising into the Early Miocene depositional surface (sect. 2.2). Dividing the mean incision depths by
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the erosional time scales of the respective catchments might give a rough approximation about the
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incision history in our study area (Tab. 1). This approach is highly speculative, as long-term steady state erosion is assumed beyond the timescales the erosion rates integrate over. Furthermore, such a
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calculation is not applicable for the E2 catchments, as the mean incision values obtained from the DTM analysis is biased (sect. 5.1). If we assume that catchment BA15-001 (W2) reflects long-term steady erosional conditions of a pre-Río Loa incision landscape, the ridge in which the catchment incises into could have been formed, or incision could have commenced, around 8 Ma (Tab. 1). This age is in the range of the time-windows given for the uplift of the Chipana-Adamito block and the intensification of tectonic activity along the splay fault to the east (Ritter et al., 2018b). Furthermore, the onset of incision of large paleochannels into alluvial fan surfaces on the opposite side of the canyon has been dated to ~9 Ma by cosmogenic exposure dating, and fluvial incision in these paleochannels ended latest by 2 Ma due to tectonic truncation upstream (Ritter et al., 2018b). Assuming that these paleochannels continued westwards, probably along the trace of the present-day Río Loa Canon, their incision could have provided a significant base level drop on the order of > 20-30 m for all catchments in our study area at 42
Journal Pre-proof some time between 9 and 2 Myr ago. This pre-Río Loa incision provided more than the required base level drop for the eastern catchment groups (E1 and E2) to incise their present-day thalwegs. Furthermore, it provided most of the required base level drop for catchment group W2 and about 50% of the base level drop required for catchment group W1. Latest with the incision of the Rio Loa Canyon 250-300 kyr ago (Ritter et al., 2018b; Vásquez et al., 2018), the latter received more than the required base level lowering to form their present-day thalwegs. The reconstructed time required to incise channels to their present depth, assuming the current erosion
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rates as being constant through time, is on the order of 1-2 Myr for catchments groups W1, W2 and E1.
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This could indicate intensified erosion during the Quaternary or the base level lowering provided by the
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pre-Río Loa drainage. However, this chronological reconstruction remains speculative and is not entirely consistent, since remnants of gypsum crust cover are present on hillslopes of all catchment groups. The
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gypsum crusts on the hillslopes are indicative of a polyphase (at least two-phase) incision of the
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channels. A significant period of hillslope stability is required to accumulate decimeter-scale gypsum crusts. Deriving the time required for the formation for such a gypsum layer is not straightforward. Rech
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et al. (2003) estimated that in the westernmost parts of Coastal Cordillera and at altitudes of about 700 m,
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about 1 m of soil gypsum could be accumulated on a gently sloping surface within ~1.9 Myr. For the Namib Desert, it has been reported that clearly visible gypsum crystals are formed within 10 kyr, while
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the formation of mature gypcretes of about 1 m thickness possibly takes around 800 kyr (Heine and Walter, 1996). While the corresponding accumulation rates are highly dependent on local conditions, these numbers indicate that long timespans are required to form a thick gypsum cover. We infer that the hillslope-response to channel incision in the study area had at least one long-term hiatus on the order of several 100 kyr or more.
6 Conclusion The topography of the study area reflects a long-term erosional gradient between a detachment-limited regime of relatively rapidly eroding drainage systems in the western portion and a transport-limited regime of extremely slow erosion in the eastern section. Cosmogenic nuclide-derived erosion rates increase more than 40-fold from east to west over a distance of about 2.5 km. The erosion regime and 43
Journal Pre-proof rates appear to be influenced by the presence and state of a gypsum cover and/or the state of the pervasive halite-cement in the otherwise unconsolidated continental sediments into which the drainage has incised. Contrasts in the state of the gypsum cover and halite cement are driven by microclimatic conditions: (1) directional erosion by strong westerly winds, (2) aspect-dependent fog precipitation and (3) aspectdependent evapotranspiration. Tectonic movements were likely important to create the modest topography (i.e. a 20-30 m high N-S ridge) that spatially coincides with the transition of erosion regimes and state of the ground cover (gypsum crust). The functional relationship of the ridge to the fluvial
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evolution of the study area is that of a wind-break for areas to the east and increased fog precipitation
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for areas to the west. In general, the state of the ground cover – and thus the erosional regime – appears
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to be predominantly governed by abiotic factors.
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Acknowledgements
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Thanks go to the field campaign teams in 2015 and 2016 and to all colleagues and students who helped collecting and preparing the samples. Elena Voronina and Tomasz Goral made a great job processing
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the samples in the lab. A very special thank goes to Diego Jaldin of the Universidad Catolica del Norte
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(UCN), who was a more than valuable help during the fieldwork conducted in October 2018 and who provided important scientific inputs. Moreover, we would like to thank Eduardo Campos and colleagues
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at the UCN in Antofagasta for their support. Finally, we thank two anonymous reviewers whose suggestions and comments have greatly improved this manuscript. The research was part of a pilot study for a coordinated research centre funded by the Deutsche Forschungsgemeinschaft (DFG, German Research Foundation) – Projektnummer 268236062 – SFB 1211.
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Ritz, J.-F., Braucher, R., Bourlès, D., Delouis, B., Marquardt, C., Lavenu, A., Philip, H., and Ortlieb, L.: Slip rate of trench-parallel normal faulting along the Mejillones Fault (Atacama Fault System): Relationships with the northern Chile subduction and implications for seismic hazards, Terra Nova, 31, 390-404, 2019.
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Journal Pre-proof Schumm, S. A. and Lichty, R. W.: Time, space, and causality in geomorphology, American Journal of Science, 263, 110-119, 1965. Sepulveda, F. and Vásquez, P.: Estratos del Cañón del Loa: nuevo registro sedimentario paleógeno en la cordillera de la Costa del norte de Chile (21°S), XV Congreso Geológico Chileno, 2018. Sepúlveda, F. A. and Vásquez, P.: Paleogene (Eocene?-Oligocene) sedimentation and associated tectonic setting in the Coastal Cordillera of northern Chile (21°-21°30’ S), Geophysical Research Abstracts, 21, 2019.
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Wilcox, A. C., Escauriaza, C., Agredano, R., Mignot, E., Zuazo, V., Otárola, S., Castro, L., Gironás, J., Cienfuegos, R., and Mao, L.: An integrated analysis of the March 2015 Atacama floods, Geophysical Research Letters, 43, 8035-8043, 2016. Yang, J.-L., Zhang, G.-L., Yang, F., Yang, R.-M., Yi, C., Li, D.-C., Zhao, Y.-G., and Liu, F.: Controlling effects of surface crusts on water infiltration in an arid desert area of Northwest China, Journal of Soils and Sediments, 16, 2408-2418, 2016.
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Journal Pre-proof Highlights We investigate erosion patterns in the hyperarid core of the Atacama Desert
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Erosion rates increase steadily by one order of magnitude within 2.5 km
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DTM analysis reveals the existence of two different erosional regimes
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The presence of gypcrete in the study area is linked to microclimatic conditions
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Gypsum surface cover affects surface processes
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Journal Pre-proof Declaration of interests
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☒ The authors declare that they have no known competing financial interests or personal relationships that could have appeared to influence the work reported in this paper.
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Joel Mohren, Steven A. Binnie, Benedikt Ritter, Tibor J. Dunai PII:
S0921-8181(19)30553-3
DOI:
https://doi.org/10.1016/j.gloplacha.2019.103068
Reference:
GLOBAL 103068
To appear in:
Global and Planetary Change
Received date:
29 May 2019
Revised date:
17 October 2019
Accepted date:
25 October 2019
Please cite this article as: J. Mohren, S.A. Binnie, B. Ritter, et al., Development of a steep erosional gradient over a short distance in the hyperarid core of the Atacama Desert, northern Chile, Global and Planetary Change(2019), https://doi.org/10.1016/ j.gloplacha.2019.103068
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© 2019 Published by Elsevier.
Journal Pre-proof Development of a steep erosional gradient over a short distance in the hyperarid core of the Atacama Desert, northern Chile Joel Mohren*1, Steven A. Binnie1, Benedikt Ritter1, Tibor J. Dunai1
1
Institute of Geology and Mineralogy, University of Cologne, Germany
Zülpicher Str. 49b, 50674 Cologne
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*Corresponding author. Email: [email protected]
Abstract
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Although generally considered to exist under hyperarid conditions over the long term, landscapes in
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many parts of the Coastal Cordillera of northern Chile have undergone fluvial erosion. Small-scale drainage systems in this mountain range are mostly isolated from river networks and associated
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processes originating in the Precordillera or the High Andes to the east, thus providing natural
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laboratories to investigate the interplay between erosion, atmospheric deposition, tectonics and local (micro-) climatic conditions. In this study, we present a set of cosmogenic
10
Be and
26
Al derived,
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catchment-wide erosion rates along a short (2.5 km) E-W transect on the northern rim of the Río Loa Canyon in the Coastal Cordillera of northern Chile (latitude 21.4°S). Here, a flat sedimentary gravel surface, which was deposited before the Middle Miocene, becomes increasingly dissected and changes into a badland-like topography to the west. The 10Be erosion rates increase by approximately an order of magnitude from east to west, reflecting (1) localized tectonic movements, (2) geologically recent base level lowering, (3) time-integrated (micro-) climatic gradients and (4) the presence/absence of gypcrete. These findings are corroborated by analysis of geomorphologic parameters, which point towards the presence of two fundamentally different erosional regimes in this small study area. These regimes are sharply delineated along a topographically modest tectonic ridge. To the west, a detachment-limited erosion regime prevails, while in the east a transport-limited regime is dominant. The presence or absence of gypcrete, whose prevalence is governed by (micro-) climatic conditions, generally appears 1
Journal Pre-proof to reflect the respective erosional regimes. The erosion rates we infer point to a long-term process of differential drainage evolution in the study area, likely on timescales of millions of years.
Keywords: Geomorphology; Gypsum; Erosion; Crust; Cosmogenic nuclides; Atacama Desert
1 Introduction The formation of topography is directly linked to a variety of endogenic and exogenic processes, such
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as tectonics, erosion, aggradation and weathering. The efficiency of exogenic processes is influenced by parameters such as rock type, climate and vegetation cover and their likely interplay, but the differing
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timescales over which such Earth-shaping processes integrate and variations in their influence over time
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can blur our understanding of long-term topography formation. Thus, for the study of the interactions between Earth-shaping processes, environmental conditions and their impact on topography formation,
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it is often desirable to isolate the influence of individual variables, while maintaining consistency in other factors over the timescale that is being analyzed. For timescales on the order of > 104 yrs,
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encompassing climatic cycles, this requirement is rarely met. On the other hand, the question about how
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changing environmental conditions affect topography formation is – also in the context of modern
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climate change and its impact on future conditions – one of the most important issues in the field of geomorphology, underlining the relevance of the long-term perspective. In this study we analyze a unique landscape located within the hyperarid core of the Atacama Desert in northern Chile (cf. Ritter et al., 2018b; Ritter et al., 2019). The onset and secular variation of aridity in the Atacama Desert, as well as spatial variability, are still a matter of scientific debate. Estimates for the onset of hyperaridity range between the Pleistocene and the Oligocene (e.g. Alpers and Brimhall, 1988; Amundson et al., 2012; Betancourt et al., 2000; Carrizo et al., 2008; de Porras et al., 2017; Dunai et al., 2005; Evenstar et al., 2009; Evenstar et al., 2017; Gayo et al., 2012; Hartley and Chong, 2002; Jordan et al., 2014; Nishiizumi et al., 2005; Rech et al., 2006; Reich et al., 2009; Ritter et al., 2018b; Ritter et al., 2019; Sáez et al., 2012; Sáez et al., 1999; Sillitoe and McKee, 1996; Veit et al., 2016; Wang et al., 2015). The geological record points to predominantly arid conditions in the region since 150 Ma (Hartley et al., 2005). Long-term stable climatic conditions in the Coastal Cordillera of northern Chile 2
Journal Pre-proof are achieved by (1) the steady latitudinal location of the South American continent since the MidCretaceous (Somoza and Zaffarana, 2008); (2) the rainshadow effect of the High Andes blocking westward moisture transport originating from the Amazonian region (e.g. Houston and Hartley, 2003; Weischet, 1966); (3) the cold Peru-Chile Current cooling the air masses and thus reducing the possible amount of absolute humidity to be transported with onshore winds along the western coast of South America (e.g. Houston, 2006a; Stoertz and Ericksen, 1974), and (4) the perennial atmospheric subsidence caused by the Hadley circulation (e.g. Houston, 2006b; Houston and Hartley, 2003) and the
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Rutllant circulation (Rutllant, 2003). Taken together, these factors lead to an excess of mean annual
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potential evaporation by two to three orders of magnitude when compared to mean annual precipitation
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in the Atacama Desert (Houston, 2006a). Some areas within the hyperarid core of the desert, which is located between ~19°S and ~22°S and includes the Coastal Cordillera and the Central Depression (cf.
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Ritter et al., 2018b), currently receive less than 2 mm yr-1 of precipitation from rainfall (e.g. Houston,
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2006a). In these areas, signs of significant fluvial erosion are rare, and most of the landscape appears to be fossilized under a blanket of gypsum dust and gypcrete (Ericksen, 1981; Mortensen, 1929; Mortimer,
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1980). Exposure ages in these ‘fossilized’ areas are predominantly of Miocene age (cf. Carrizo et al.,
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2008; Dunai et al., 2005; Evenstar et al., 2017; Nishiizumi et al., 2005; Ritter et al., 2018b). Further to the west, within the Coastal Cordillera in the vicinity of the coast, apparently well-developed
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drainage networks indicate a fluvial response to base level change and/or to increasing precipitation. However, significant fluvial erosion in this area (~19°S to ~22°S) is restricted to (1) the Río Loa, a major drainage system crossing the entire Coastal Cordillera from the High Andes to the Pacific Ocean and (2) local catchments, i.e. catchments hosted entirely within the Coastal Cordillera. In the latter, only a few incise down to the base level (sea-level or coastal abrasion platform) as many of these local drainages terminate as hanging valleys high on the coastal cliff, illustrating that incision into the face of the cliff is slower than cliff retreat (Mortimer, 1980). The catchments that do incise to base level have very low denudation rates, ~2 m Myr-1 for catchments near our study area, in spite of their steep gradients (~400 to ~700 m mean local relief on 5 km radius; Starke et al., 2017). Our study area incorporates a zone of fluvially shaped topography near the coast, in the immediate vicinity of the northern rim of the Rio Loa canyon. This particular setting allows the investigation of 3
Journal Pre-proof spatially resolved erosion rates and surface processes along a west-east transect that displays a progressive reduction in the amount of fluvial incision despite experiencing a common base-level (i.e. the Río Loa, Fig. 1). The catchments are small (< 0.3 km2; Figs. 1, 2) and so record the affects of longterm local climatic conditions, rather than an allochthone climate signal (cf. Ritter et al., 2018b; Ritter et al., 2019). Furthermore, the catchments in the study area incise into a rather homogeneous and uniform substrate, consisting of two sedimentary units which have been successively deposited in the time period between Eocene?-Upper Oligocene and Miocene (Sepúlveda and Vásquez, 2019; Vásquez et al., 2018).
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Irrespective of their age, both of these alluvial layers presumably behave similarly to fluvial erosion,
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meaning that substrate/lithological parameters can be considered constant throughout the area studied.
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As such, our study location provides a unique natural laboratory for the investigation of the temporal
the hyperarid core of the Atacama Desert.
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and spatial variability of precipitation and its impacts on the formation of topography on the fringe of
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Our results from cosmogenic nuclide measurements and digital terrain model analysis confirm the presence of a time-integrated gradient in erosional style and rates. We will show that the likely causes
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for this gradient relate to: moisture sources (Pacific); the interception of that moisture by subtle,
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2 Regional setting
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intervening, fault-driven relief; and the resulting presence/preservation of gypcrete.
2.1 Topographic characteristics
The study area is a spatially short (~2500 m) E-W transect located on the northern rim of the Río Loa canyon, at an elevation of about 700 m a.s.l. within the Coastal Cordillera (21.42° S, 69.98° W; Fig. 1). Approximately 8 km to the west, the Río Loa drains into the Pacific Ocean; the closest cities are Iquique to the north (~130 km), Quillagua to the southeast (~50 km) and Tocopilla to the south (~80 km).
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Figure 1. Sentinel 1B true color image (L2A product) showing the location of the study area in the
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Coastal Cordillera at the northern rim of the Río Loa canyon, northern Chile. Fault systems in the vicinity of the study area are taken from the geological map of this region (Vásquez et al., 2018) or
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based on own observations and data analysis. A strong E-W gradient in fluvial erosion is apparent, with
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pronounced incision in the west and limited incision to the east. The study area straddles this gradient where it is most pronounced. Three climate stations (#11, #12, #13), from which meteorological data
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has been analyzed for this study (Hoffmeister, 2018), are indicated by white stars.
We decided to conduct our work in this area because the topography and surface patterns change remarkably along the short spatial transect. In the eastern portion, the landscape can generally be characterized as smooth, with gypsum-rich chuca and costra (sensu Ericksen, 1981) covering almost the entire surface, which is gently dipping to the southwest. The small-scale drainage systems (< 0.3 km2) are predominantly N-S oriented and terminate as hanging valleys in the rim of the much larger and more deeply incised Río Loa canyon. Local relief is generally low and slopes show high convexities (Fig. 2). A significant proportion of the eastern sector does not show any signs of fluvial erosion as surfaces are flat and encrusted, with costra being exposed. Gypsum crust has formed polygonal structures and channel bottoms are partly covered by chuca. In some places, especially along 5
Journal Pre-proof slopes and a few meters below the rim of the northern flank of the Loa canyon, a > 20 cm thick salt layer
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mainly consisting of halite and sometimes forming nodules, crops out (Fig. 10).
Figure 2. Handheld camera (A-B) and DJI Spark drone (C-F) pictures illustrating the geomorphological E-W contrasts in the study area. The channels of the western catchments (groups W1 and W2; see Fig. 3) have incised relatively deeply into the gravels (A, C). In contrast to that, the relief in the eastern catchments (groups E1 and E2) is much lower (B, D with BA15-016 in the foreground). A local ridge 6
Journal Pre-proof sharply separates encrusted (east) from non-encrusted surfaces (west; E). However, small patches of gypsum crust can be found in the western catchment BA15-004, covering a small slope and a plateau above subcatchment BA15-006 (dashed polygons in C and F). In some areas, salt crops out along slopes (white arrows in C and F; see also Fig. 10).
By contrast, fluvial erosion has formed a well-developed drainage system in the western portion of the transect. In comparison to the eastern drainages, incision appears to be much greater. Signs of sub-recent
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discharge, such as minor accumulations of clay in some channels and centimeter-scaled incision into
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this layer are observable, as are scars and deposits from slope failures. Although the local erosional base
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level is determined by the incision of the Río Loa to the south, the drainage patterns are mainly W-E/EW oriented. This is due to the presence of a N-S oriented local ridge into which the catchments have
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incised. The ridge is partly covered by a thick gypsum crust, up to ~40 cm on the ridge peaks and
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spatially coinciding with the sharply defined and easy-to-trace westward limit of gypsum surface cover (Fig. 2). To the west of the ridge, gypsum encrustation is absent, apart from some highly localized
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patches, and bare gravel is exposed at the surface. Many surface cracks, similar to those described by
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eastern sectors.
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Allmendinger and González (2010, and references therein), can be observed in both the western and the
2.2 Geological and geomorphological background Large parts of the Coastal Cordillera have been strongly affected by extensional tectonics since the Oligocene (Hartley et al., 2000). The tectonic regime is likely caused by the strong plate coupling between the Nazca Plate and the South American Plate during the subduction process and triggers crustal shortening (e.g. Allmendinger and González, 2010). Numerous N-S striking normal faults belonging to the Atacama Fault System (AFS; e.g. González et al., 2003), visible all over the Coastal Cordillera, are evidence for stress release. The faults promoted the formation of horst and graben structures, which can be identified along the Río Loa Canyon (González et al., 2003; Sepulveda and Vásquez, 2018, 2019; Vásquez et al., 2018). The bedrock below the alluvial surface of the study area is exposed along the ~600 m deep Loa Canyon and mainly comprises of Carboniferous metamorphic rocks (Caleta Loa 7
Journal Pre-proof complex, Csbcl) and the Upper Triassic to Lower Jurassic Sierra de Lagunas strata, mainly composed of andesites (TrJsl; Vásquez et al., 2018). Uplift of these rocks accommodated by the Atacama Fault has possibly been initiated during the Eocene and continues until the present day (Cosentino and Jordan, 2017; Juez-Larré et al., 2010). During the Early Miocene, tectonic uplift isolated large areas of the Coastal Cordillera from the influence of Precordilleran sediment aggradation (Evenstar et al., 2017; Mortimer et al., 1974). In a narrow corridor between 19° and 21.6°S latitude, trench-parallel shortening also causes E-W oriented stress release along reverse faults with pure dip slip movement (Allmendinger
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and González, 2010; Allmendinger et al., 2005). The E-W striking reverse fault systems are believed to
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have been active at least since Late Miocene and some could have remained active until the present day
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(Allmendinger and González, 2010; Allmendinger et al., 2005; Carrizo et al., 2008; González et al., 2015; Ritter et al., 2018b). The study area is located between two of these E-W striking reverse faults
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systems, namely the Chipana fault in the north and the Adamito fault to the south (Skármeta and
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Marinovic, 1981; Vásquez et al., 2018; Fig. 1). These faults are believed to define the southern and northern limits of the Chipana-Adamito block, which has possibly undergone uplift since 16-4 Ma
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(Vásquez et al., 2018) or, in the case of the Adamito fault, since 14-10 Ma (Ritter et al., 2018b). The
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uplift likely enhanced the isolation of the area from sediment sources to the north, south and east (cf. Ritter et al., 2018b; Vásquez et al., 2018). This interpretation is supported by the presence of two
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chronostratigraphic units that cover the underlying bedrock in the study area by a thickness of up to 300 m, as recently mapped by Sepúlveda and Vásquez (2019) and Vásquez et al. (2018). According to these authors, rounded to well-rounded conglomerates with a variable degree of consolidation and of possibly Eocene to Upper Oligocene age (Cañón del Loa beds, CLB) are partially covered by generally angular and moderately consolidated gravels with a maximum age of 24.69 ± 0.09 Ma (Alto Hospico gravels, AHG). On the side of the Loa canyon opposite to our study area, i.e. on its southern slopes, a tephra layer located within this unit 30 to 50 m below the surface has been dated to 21.4 ± 0.5 Ma, while a second tephra layer located close to the base of the AHG yielded an age of 23.50 ± 0.16 Ma (Vásquez et al., 2018). The minimum age obtained from this unit on the Chipana-Adamito block is 16.14 ± 0.02 Ma (all ages derived from 40Ar/39Ar dating in biotite; Vásquez et al., 2018), indicating an Early Miocene age for the entire AHG unit in the vicinity of our study area. It is likely that the majority 8
Journal Pre-proof of these gravels were deposited in an alluvial paleoenvironment under arid paleoclimatic conditions (Sepúlveda and Vásquez, 2019; Vásquez et al., 2018). The gravels of both strata are of varying lithology, including volcanic, metamorphic, intrusive and sedimentary rocks (cf. Sepulveda and Vásquez, 2018). Clast sizes as observed in the field cover the entire gravel and boulder spectra; Sepulveda and Vásquez (2018) report mean clast sizes of 6-25 cm for the CLB and 1-2 cm for the AHG, respectively. Lamination patterns consisting of decimeter thick layers in the CLB can be traced along slopes in the western portion of the study area. Gravels of the CLB unit are poorly sorted and generally cemented by salt
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(predominantly halite). In between the two gravel sequences a thick paleosol layer has been identified,
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indicating a significant temporal interruption of aggradation processes after the deposition of the Cañón
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del Loa beds, possibly between 31 and 26 Ma (Sepúlveda and Vásquez, 2019). Most N-S striking faults in the direct vicinity of the study area were likely established prior to the
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deposition of the Cañón del Loa beds (Sepulveda and Vásquez, 2018) and were still active, or
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reactivated, after the deposition of the Alto Hospicio gravels and coeval units in the Coastal Cordillera, as the sedimentary cover shows an offset in some places (e.g. González et al., 2003; Sepulveda and
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Vásquez, 2018). This holds also for a N-S striking splay fault of the AFS to the east of the study area,
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which could have been active since about 14 Ma (Ritter et al., 2018b). However, Pleistocene displacements along N-S faults belonging to the AFS have been reported further to the South (~23.5°S;
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González et al., 2006; Ritz et al., 2019; Vargas et al., 2011). Uplift has also been likely along the E-W striking Loa reverse fault, which is located to the east of our study area (Fig. 1). Vásquez et al. (2018) postulate that the Cañón del Loa beds were uplifted along this fault before they were covered by Alto Hospicio gravels. A striking feature of comparably young surface modification in the vicinity of the study area is the deeply incised gorge of the modern Río Loa, which has been estimated to have formed since 274 ± 74 ka, based on cosmogenic nuclide surface exposure dating (Ritter et al., 2018a). This age is further supported by a 40Ar/39Ar biotite age of 255 ± 17 ka, obtained from a tephra layer emplaced within an elevated river terrace (Vásquez et al., 2018). It is important to note that the massive incision into the bedrock to a depth of around 600 m below the surface represents the communication of an allochthonous climate signal originating from the elevated areas (High Andes, Precordillera) to the east. This is underlined by the fact 9
Journal Pre-proof that the entire drainage system on the northern and southern rim of the Loa canyon is hanging as it has been bisected by incision of the Río Loa.
2.3 Regional climate characteristics Apart from the general long-term hyperarid climate conditions caused by the circulation patterns as described above (sect. 1), the present-day local climate in the study area is strongly influenced by the abundance of coastal fog (locally known as Camanchaca). Cold and moist air originating from the
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Pacific is trapped underneath an inversion layer caused by the subsidence of warm continental air masses
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leading to the formation of fog (e.g. Weischet, 1966; Weischet, 1996). As the inversion layer in northern
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Chile is on average located at an elevation around 900 m, the Coastal Cordillera with an average height of 1000 m (Hartley et al., 2000) prevents the moist air from migrating inland (Miller, 1976). Thus, low-
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lying areas like the Loa canyon are used as pathways by the advective fog to reach the hinterland (cf.
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Cereceda et al., 2008a).
The present-day occurrence, distribution and precipitation yield from the coastal fog has been studied
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in detail by Cereceda et al. (2008a, b) for a N-S transect spanning over the littoral plains and the Coastal
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Cordillera in northern Chile, located ~70 km north of the study area of this work. The authors found a strong influence of elevation and distance from the coast on the amount of precipitation that they
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measured by using fog collectors. In general, fog precipitation is mostly focused on elevations between 650 and 1200 m and in the period between April to October at the western margin of the Coastal Cordillera (Cereceda et al., 2008a, b). The authors further showed that elevation gradients have a significant impact on dew yields, increasing from 0.44 l vertical m-2 day-1 at 650 m to 7.12 vertical m-2 day-1 at 850 m (base of the fog collector 2 m above the ground; measurement period 2001 to 2004). Further moisture supply can be initiated by special synoptic conditions, which can bring significant amounts of rain to the coastal range of northern Chile (Houston, 2006a). Such events caused floodings after several heavy rainfalls in the 20th century (Ortega et al., 2019; Vargas et al., 2006). Moisture sources originate mainly in the southern or northern Pacific (e.g. Bozkurt et al., 2016; Houston, 2006a; Jordan et al., 2019). Since 2015, several unusual and severe rain events that have caused floodings in 10
Journal Pre-proof the coastal zone of northern Chile have been documented and analyzed in detail (e.g. Bozkurt et al., 2016; Jordan et al., 2015; Jordan et al., 2019; Rondanelli et al., 2019; Wilcox et al., 2016). Spatially associated to the mid-latitude winter rainfall zone, modern rainfall in the coastal range of northern Chile is mainly focused on the months March to September, with intensified precipitation favored by El Niño (Houston, 2006a; Ortega et al., 2019) and prior-El Niño conditions (Vargas et al., 2006). This holds, for example, for the March and August 2015 rain events (Jordan et al., 2019). However, the El Niño Southern Oscillation (ENSO) is believed to be only one of many factors affecting the special synoptic
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conditions that lead to rainfall events in northern Chile (Houston, 2006a; Ortega et al., 2019; Rondanelli
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et al., 2019). Jordan et al. (2019) argue that the rainfall occurring during the June 2017 event was not
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related to El Niño-like conditions but could have occurred due to the advection of cold air masses from the Pacific, as described by Vuille and Ammann (1997). For the second half of the 20th century, Ortega
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et al. (2019) reported a significant increase in extreme rainfall events during the last warm phase of the
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Pacific Decadal Oscillation (PDO) between 26°S and 32°S, emphasizing another possible factor that might affect the precipitation patterns in northern Chile. A further variable characterizing the regional
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climate in northern Chile is the diurnal wind system. A thermal gradient between the cold Pacific and
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the heated Altiplano causes strong near-surface advection towards the hinterland, resulting in strong afternoon westerly winds between the coast and the Western Cordillera (Richter and Schmidt, 2002;
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Rutllant, 2003). During the night, this pattern is reversed with weaker oceanward winds prevailing (Richter and Schmidt, 2002).
3 Material and methods 3.1 Sampling strategy and sample analysis Concentrations of terrestrial cosmogenic nuclides (TCN), especially those of cosmogenic
10
Be, have
routinely been measured to quantify catchment-mean erosion rates (e.g. Bierman and Steig, 1996; Binnie et al., 2006; Granger et al., 1996; Placzek et al., 2014; Placzek et al., 2010; Sosa Gonzalez et al., 2016). In this study, we take advantage of the fact that the concentration of a given radionuclide decreases following burial at depths of several meters and shielding from cosmic rays. The half-life of 10Be has been determined to be 1.387 ± 0.01 Myr (Chmeleff et al., 2010; Korschinek et al., 2010) and the half11
Journal Pre-proof life of 26Al is considered to be 705 kyr (Nishiizumi, 2004). Assuming the youngest age of ca. 16 Ma for the deposition of the Cañón del Loa beds and the Alto Hospicio gravels as found by Vásquez et al. (2018), would mean < 0.1% of any in situ-produced 10Be that was present in the gravels prior to their deposition will remain, and approximately half as much
26
Al. Thus, it is very likely that the
concentrations of 10Be in the Cañón del Loa beds and the Alto Hospicio gravels are a reflection of their post-depositional exposure history. Steady (hyperarid) climatic conditions over the long-term perspective (sect. 2.3) likely promote long-term steady-state erosion, a key assumption for the
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application of TCN to infer erosion rates (e.g. Dunai, 2010; Granger et al., 1996; Lal, 1991). By
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collecting pebbles, we exclude the possibility of sampling wind-blown material in an environment
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experiencing high wind velocities. Grain size-dependent differences in TCN concentrations as described by Codilean et al. (2014) should not be of great significance in our study area as the incision of the
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sampled catchments took place entirely into alluvium and not into vein-bearing bedrock. In addition to
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that, (sub-)catchment sizes are very small (0.0014 to 0.2619 km2), minimizing transfer times. This circumstance should also reduce the possibility of incorporating significant amounts of pebbles
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originating directly from gully-side slopes adjacent to the sampling sites. Placzek et al. (2014) have
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shown that TCN concentrations from channel sediments in the hyperarid Atacama Desert can be significantly affected by sediment contributions from local slopes, which tend to cause bias by
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increasing the overall nuclide TCN concentration in the channel sediments. However, such a mechanism seems predominantly applicable to grain sizes smaller than those sampled here (Placzek et al., 2014). Furthermore, amalgamating a sufficiently large number of clasts should ensure that the TCN concentration reflects the sediment contribution of the entire drainage system. Though it is not possible to mix the pebbles in proportion to their erosion rates, precluding assumptions of thorough sediment mixing in proportion to the erosion rates of the contributing area (e.g. Bierman and Steig, 1996; Binnie et al., 2006), the production rate differences within the catchments that would cause this to be problematic are minor. We sampled 13 catchments and subcatchments along the E-W topographic gradient illustrated in Figure 1. We aimed at collecting 30 individual quartz-bearing pebbles per catchment within less than 100 m of each other along the main channel of the respective catchment and amalgamated these samples 12
Journal Pre-proof (cf. Repka et al., 1997). We avoided sampling close to slope failures, directly below knickpoints or channel mouths and took care not to collect similar-looking fragments which could have originated from the same pebbles that were broken apart by thermal cracking (cf. Dunai, 2010) or due to salt and/or gypsum expansion, which are widespread phenomena in our study area (Fig. 10). Pebbles were only collected from the bottom and center of the channels and not from lateral stream banks. In the easternmost catchments, this restrictive sampling strategy led to lower yields of suitable pebbles as surface abundance of grain sizes larger than sand decreased markedly. We avoided the catchments that
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drain the elevated hinterland to the north to minimize differences in erosion-controlling variables
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(Fig. 3).
Figure 3: Oblique view of the study area (ArcScene 10.5.1) based on Pleíades 1B satellite multispectral imagery showing the location of the sampled catchments within the E-W transect (framed in red color) and the sampled bedrock knickpoints (red triangles). Channel networks of the catchments sampled are shown in green (including the channels into which the sampled catchments shed), all other large stream networks in blue. In the western portion, sample sites are located so as to reflect local climate conditions, as the drainage is not fed by the more elevated, larger upstream catchment areas to the north (thick black solid line for upstream basin limits and black dashed line indicating general direction of flow towards the south). This is applicable also for catchment BA15-016 in the eastern portion (group E2). All other sampled eastern catchments drain into the trunk streams of catchments that reach the 13
Journal Pre-proof elevated areas to the north, where bedrock is exposed. However, these subcatchments have their channel heads further south and they are entirely hosted in the gravels.
In order to assess the importance of proximity to the local base level on erosion we also decided to sample small N-S transects perpendicular to the Río Loa canyon in the western part of the study area (Group W2: BA15-001 to BA15-003). Inner-catchment variability of erosion rates was investigated by sampling subcatchments within the catchment BA15-004 to the west of the local ridge (Group W1:
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BA15-004 though BA15-007; BA15-010). In the eastern portion, all sampled catchments are N-S
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oriented. BA15-012 and BA16-001 (Group E1) are located in the transition zone towards the
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easternmost catchments that do not show any signs of sub-recent erosion (Group E2: BA15-013, BA15-014, BA15-016). In addition, two prominent bedrock knickpoints (BA15-008, BA15-009) were
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sampled to determine local bedrock erosion rates. The outcropping rocks belong to the Caleta Loa
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metamorphic complex (Csbcl, Figs. 3, 4) and represent the closest bedrock exposures with regard to the W1 catchments. At the knickpoints, the thalweg is offset vertically by ~4 m (BA15-008) and ~8 m
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(BA15-009), respectively.
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The samples were prepared as Accelerator Mass Spectrometry (AMS) targets in the laboratories of the Institute of Geology and Mineralogy, University of Cologne, Germany. Mean clast thicknesses of the
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amalgamated samples varied between 4.1 and 0.6 cm (measured over the shortest axis; ntotal = 335). The majority of clasts (75 %) could be described as “pure quartz” or as having “high (> 80%) quartz content”. Clast colors were highly variable but the majority were greyish without desert varnish, angular to rounded and of low to medium sphericity. The clasts were individually crushed and sieved; either the 250-710 µm (bedrock samples) or 500-710 µm (channel pebbles) grain size fractions were analyzed. The respective fractions from each clast were checked for impurities and hand-picked if necessary. An equal amount of sample material from each clast was amalgamated into one sample. Afterwards, the amalgamated samples and the bedrock samples were etched following the laboratory protocols based on Kohl and Nishiizumi (1992) and quartz purity was checked using an in-house ICP-OES. The etched quartz material was dissolved following spiking with certified, commercially available Be (Scharlab, 1000 mg/l) and Al (Scharlab, 1000 mg/l) carrier solutions. Subsequent preparation of the samples as Be 14
Journal Pre-proof and Al targets for AMS followed the stacked column approach detailed in Binnie et al. (2015), coprecipitating Al and Be hydroxides with Ag following Stone et al. (2004). Chemical blanks were prepared and measured in tandem with the samples. 10Be/9Be and 26Al/27Al values were measured on CologneAMS (Dewald et al., 2013). 10Be/9Be ratios were normalized to the ICN standard dilution series (Nishiizumi et al., 2007) and 26Al/27Al ratios were normalized to the 26Al AMS standards provided by Nishiizumi (2004). The stable Al contents of the dissolved, spiked samples were determined using ICPOES and standard addition (4 aliquots) in tandem with quality control measurements of NIST SRM165a.
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Reagent blank corrected concentrations of 10Be and 26Al were derived following Binnie et al. (2019).
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Uncertainties on the concentrations include the propagated uncertainties in the AMS ratios of the
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samples and blanks together with the estimated standard deviation of the amount of 9Be, or 27Al, the samples contained after spiking.
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The LSDn scaling scheme (Lifton et al., 2014) was chosen to calculate the erosion rates using the online
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calculator described in Balco et al. (2008) and formerly known as the CRONUS-Earth online calculator (version 3; https://hess.ess.washington.edu/math/v3/v3_erosion_in.html). The input used in the online
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calculator is shown in the supplement. We did not correct production rates for uplift as the likely bias
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this would induce is not significant for the erosion rates within the precision of our measurements (see supplement). To estimate the attenuation length of the cosmogenic radiation, and thus approximate the
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averaging time of our erosion rate measurements, we assumed that the clasts have a mean density of 2.7 g cm-3. The relative porosity of gravel comparable to that we find in our study area is likely to be around 30% (Hölting and Coldewey, 2013), resulting in a value of around 1.9 g cm-3 for the density of the conglomerates and ~85 cm for the attenuation path length (cf. Ewing et al., 2006; see supplement). The density of the bedrock we sampled at the two knickpoints was measured to be approximately 2.6 g cm-3, resulting in an attenuation path length of 60 cm. This approach only accounts for spallogenic nuclide production, which makes up the majority of the total 10Be and 26Al production in quartz, especially in relatively slowly eroding catchments (e.g. Dunai, 2010). Topographic shielding was calculated from a 1 m digital terrain model (DTM) using the ArcMap-Toolboxes provided by Li (2013, 2018). The DTM was obtained from a stereo pair of Pleíades 1B images acquired on 04/11/2014 by using the Geomatica 2017 software. 15
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3.2 DTM-based data and climate data analysis For the calculation of catchment-mean shielding, as well as for all other DTM-based data we consistently included only those areas in our analysis which could possibly contribute to the erosion rates that we measure (cf. Fryirs and Brierley, 2012). Thus, we used a 10x10 m moving window in ArcGIS (ver. 10.5.1) to exclude all flat areas that do not appear able to contribute sediments to the fluvial system. These areas were defined as those pixels where the surface does not change more than 1 m in elevation
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within 100 m2 around that pixel.
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Apart from delineating the respective catchment areas, mean elevations and gradients, as well as
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geographic centers, we also traced fault scarps and derived values for some geomorphological parameters from the digital terrain data. Vásquez et al. (2018) stated that fault activity strongly
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influences erosional processes in northern Chile. Thus, hill- and slopeshade maps were used to identify
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possible faults crossing the study area. Along the Loa canyon, outcrops and the level of the pre-incision surface were mapped using the DTM and the Pleíades images. Swath profiles were drawn to assess
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surface tilting. During a field trip conducted in October 2018, aerial imagery was acquired using a DJI
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Spark drone. From these images, we created orthomosaics of all catchments except for those of group E2 (BA15-013, -014, -016) using Agisoft Photoscan software (ver. 1.4.4). The mosaics (3-5 cm
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resolution), the raw aerial pictures that were available for all catchments and the orthorectified and pansharpened Pleíades images (~50 cm resolution) were used to manually extract stream networks and obtain the total surface stream length per catchment. Furthermore, we analyzed our dataset for any correlation between the widely used geomorphological parameters drainage density (Dd), mean relief, mean slope and mean hilltop curvature (CHT), and the erosion rates (e.g. Ahnert, 1970; Granger and Riebe, 2014; Hurst et al., 2013; Hurst et al., 2012). Since some subcatchments contain first-order streams, and thus do not contain any slopes except for the bounding slopes that define the local watershed and thus the subcatchment limit, we included the divides into the hilltop curvature analysis. For the hilltop curvature data, we inverted the elevations of the DTM and extracted the artificially low topographic networks using the standard flow direction/flow accumulation procedures in ArGIS. The lines of this network, representing the ridges and hilltops, were buffered to 2 m so that values from 16
Journal Pre-proof neighboring catchments sharing a ridge could be assigned to both catchments. In catchment BA15-002, we had to remove a group of ridges manually from an area that appeared as a flat surface in the inverted DTM. The buffered ridges were used to clip a plan curvature raster of the study area, from which the values used in this study were extracted. As local uplift and different catchment orientation likely have a strong influence on the geomorphic parameters, special focus was also paid to a comparison of the catchments facing similar conditions, i.e. group W1 vs. W2 and E1 vs. E2. To avoid a possible bias of artificial relief generation in the eastern
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catchments due to the general dip of the pre-incision surface towards the south, we also used the mean
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incision depth of each catchment as surrogate for relief to correlate to our erosion rates. We derived
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these values by measuring elevations along a swath profile oriented parallel to the long axis (in E-W or N-S direction) of the respective catchments. The swath width was defined by the drainage divides of the
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catchment and so varied along the path of the profile. The maximum differences in elevation collected
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along the path length of the swath profile at a 1 m resolution were finally averaged to calculate the mean incision depth.
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As well as the quantification of process rates and geomorphological parameters we also analyzed climate
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data to characterize present-day climatic conditions in our study area in detail. Hoffmeister (2018) and other researchers within the Collaborative Research Centre (CRC) 1211 (Earth – Evolution at the dry
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limit) now provide freely available datasets from multiple meteorological stations installed along E-W running transects in northern Chile (DOI: 10.5880/CRC1211DB.4). As our study area is located directly within one of these transects, climate datasets are available from the littoral plain ~8 km to the west, from the eastern margin of our study area and from the Cerros de Calate ~14 km to the east (stations #11, #12, #13; for location see Figs. 1 and 5). The range of comparable data spans October 2017 to January 2019 (~460 days). Climate stations #12 and #13 were also operating during the June 2017 rain event, which we thus included in our analysis. Furthermore, we generated a small dataset to obtain some insight into the present-day microclimatic conditions prevailing on the western and eastern slopes of the ridge separating the catchments of groups W1 and W2. During the field trip conducted in October 2018, we measured ground temperature and relative humidity using Voltcraft DL-210TH data loggers for a ~7-days period (10/18/2018, 07:00 pm to 10/26/2018, 08:00 am). We stationed 3 pairs of data loggers 17
Journal Pre-proof directly on the surface at ~710 to 720 m altitude (DTM elevation) on the opposing pairs of east- and west-facing ridges. The single units of each pair were placed in a similar distance from the ridge (see supplement). We placed a Kestrel 5500 Weather Meter on the central gypsum-covered peak on top of the ridge between the W1 and W2 catchments, to measure wind speeds 70 cm above the ground (732 m a.s.l).
4 Results
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4.1 Local tectonics
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Mapping of outcrops along the Loa canyon reveals that (half-)graben structures exist in the study area
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as previously described by Vásquez et al. (2018); the bedrock located below the local ridge possibly represents a paleohorst structure (A in Fig. 4). As illustrated at the eastern margin of the elevation profile
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in figure 4, relative uplift is accommodated along the splay fault described by Ritter et al. (2018b) and
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Vásquez et al. (2018). A paleosurface separating the CLB and AHG units below the eastern part of the study area is not embedded horizontally but dips about 1.7° to the east, coinciding with the general
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eastward dip of surfaces that lie to the east of the ridge. By contrast, the pre-incision surface dips about
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1° to the west. By slope and relief analysis, we were also able to trace tectonically induced offsets that likely represent further extensions of a NW-SE running normal fault and/or of the E-W running Loa
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reverse fault as mapped by Vásquez et al. (2018). Geometry of the fault planes appears to be such that they would meet at an area close to the local ridge to the north of groups W1 and W2, if further extrapolated (Fig. 1; supplement). Surface displacement along the extension of the Loa fault can also be inferred from the elevation profile (Fig. 4). In addition to that, a sharp contrast in incision along the E-W transect coincides with a large channel that by itself marks the outer limit of significant erosion to the east. As this pattern is consistent along the channel until it has crossed the elevated areas to the north, we assume that the channel could follow a fault scarp, which has promoted differential fluvial incision and could thus be responsible for the sharp incision contrast between the E1 and E2 catchments (Fig.1, B in Fig. 4; supplement).
18
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Figure 4: Swath profile (1 m resolution, derived from Pleiades 1B satellite images) and geological cross section (the origin is in the west) across the study area (for extent see Fig. 1). The mean elevation profile
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is indicated by the thick black line (± std. dev. in grey), the thin black line represents the minimum elevation which corresponds to the Loa canyon rim. Colored diamonds denote sampled mean catchment
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locations and -groups; green triangles represent the sampled bedrock knickpoints. The dashed orange
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line represents the linear fit of all calibration points (orange points) placed on the flat surface to the east of the study area, interpolating the possible dip of the pre-incision surface to the west. Geological units and faults were mapped based on Vásquez et al. (2018) and own analysis of remote sensing data [Csbcl – Caleta Loa complex; TrJsl(a) – Sierra de Lagunas strata; Jich – Chuculay fm; Kip – Paiquina complex; CLB – Cañón del Loa beds; AHG – Alto Hospicio gravels]. The bedrock contact and mapped faults indicate the presence of several (half-) graben structures, whereby all faults except for the two westernmost (undetermined) are normal faults. The easternmost fault corresponds to a splay of the Atacama Fault, as described by Ritter et al. (2018b) and Vásquez et al. (2018). The area located between sampling groups W1 and W2, spatially corresponding to the local ridge, might represent a paleohorst structure, indicated by a sequence of serrated bedrock outcrops (Point A). A channel that has been cut by the incision of the Loa river might follow a N-S striking fault scarp (Point B; Fig. 1). Outcrops of a 19
Journal Pre-proof paleosurface separating the AHG and CLB units are indicated by grey points; its eastward average dip below the study area indicated by a pink dashed line.
4.2 Local climate Analysis of the data provided by Hoffmeister (2018) revealed that station #12 recorded the lowest mean air temperature (13.4°C; 15.4°C during daytime 07:30 am to 07:30 pm) and soil surface temperatures (14.5°C; 20.4°C during daytime) of all three stations. Minimum soil surface temperatures at station #12
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decreased below 0°C but maximum temperatures reached almost 40°C during the timespan that was
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analyzed (Fig. 5).
Figure 5. Swath profile (30 m resolution, SRTM data) across the “center transect” of the Collaborative Research Centre (CRC) 1211 (for spatial extent see Fig. 1) including climate stations #11, #12 and #13 (red diamonds; Hoffmeister, 2018). Weather station #12 is located on the eastern margin of the study area of this work (red dashed polygon). A color code has been used to visualize differences in the climate variables air temperature (measured 2 m above the ground; T air), relative humidity (2 m; Φ), surface 20
Journal Pre-proof temperature (Tsurf.), soil temperature (-10 cm; Tsoil), wind direction (2 m; W.dir.), wind speed (2m; vwind), relative amount of time a leaf wetness sensor has been moistened (20 cm, Wet) and precipitation, as measured during the timespan 10/01/2017 to 01/10/2019 (10 min measurement interval). Sea sprayinduced salt encrusting might have influenced the measurements of the leaf wetness sensor at station #11 (marked with *). Red color indicates driest or hottest conditions in relation to the other two climate stations, blue color vice versa and yellow medium values. Values are expressed as average (𝑥̅ ), average during daytime (07:30 am to 07:30 pm; 𝑥̅ day), maximum, minimum and summarized (Total). For
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discussion see text.
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Average values for the relative air humidity of station #12 are the highest when compared to the other
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stations (75.9%; 68.5% during daytime), most likely a response to lower temperatures. Precipitation from rain did not differ significantly between the coastal plain (0.8 mm) and the eastern boundary of our
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study area (0.9 mm) during the analyzed period. In contrast to that, the station #13 at the Cerros de Calate (1148 m a.s.l.) recorded 14.1 mm precipitation, of which a great amount could have originated
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from heavy fog; a similar assumption might be valid for the other stations (cf. McKay et al., 2003). The
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maximum amount from a single event at station #12 was 0.5 mm of precipitation within 10 min on
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2/27/2018 in the afternoon (2.6 mm/10 min at station #13). The occurrence of water precipitating on the surfaces is generally highest close to the sea, where surfaces were moist ~57% of the measurement period (measurement height 20 cm above the ground). However, measurements close to the coast at station #11 have likely been affected by salt deposition on the sensor due to sea spray (D. Hoffmeister and J. Schween 2019, pers. comm.). At station #12, this value should strongly be linked to the occurrence of fog. Accordingly, the vicinity of the study area might have been about ~13% of the measurement period covered by fog. At the Cerros de Calate, this value decreases to ~6%. Similar patterns can be observed for the mean daytime occurrence of fog (#11: 29.9%, #12: 3.3%, #13: 0.8%). At station #12, these values would imply fog abundances during the daytime of about 16 days as well as during 46 nights. The recent rain events in northern Chile have had limited impact on surface modification in our study area. Based on our own observations made during three visits (02/2015, 11/2016, 10/2018) we can state 21
Journal Pre-proof that some surface modification occurred after the 2015 rain events which led to fluvial accumulation of clays in small depressions of the eastern flat surface as well as in channels to the east of the local ridge. On steep slopes, however, we did not find signs of significant downslope material transport, while the rain recorded in February 2018 did not leave significant signs of fluvial activity in the areas we explored. The same can be stated for the June 2017 event, when CRC 1211 climate stations #12 and #13 were already operating. On the eastern edge of the study area (station #12), 1.1 mm precipitated between 11:50 pm (06/06/2017) and 03:20 am (07/06/2017). Further east at station #13, 0.7 mm of precipitation
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was recorded between 9:10 pm and 02:20 am.
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The dataset obtained from the data loggers we placed along the western and eastern slopes of the ridge
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during the one-week-measurement in October 2018 indicates some microclimatic differences (Fig. 6). The east-facing devices recorded a mean daytime temperature (07:30 am to 07:30 pm) of 24.1°C, while
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their west-facing counterparts recorded 19.7°C. A similar surface temperature was measured by the
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CRC 1211 station #12 at the eastern limit of our study area (19.1°C; Hoffmeister, 2018). The difference in temperature across the ridge is most pronounced during the morning hours before noon but persists
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during the afternoon (Fig. 6A).
Figure 6. Variability of temperature (A) and relative humidity (B) at the ground as measured during a one-week-measurement (10/18/2018 to 10/25/2018) along the western (blue lines) and eastern (red lines) flanks of the ridge separating the W1 and W2 catchments. Thick solid lines represent average 22
Journal Pre-proof values (10 minutes interval, 3 data loggers) and thin dashed lines the standard deviation. The dashed black line in (B) indicates the 75% humidity level, above which halite becomes hygroscopic (Gu et al., 2017). This threshold has been exceeded every night during the measurement period.
In addition, relative humidity at the surface dropped from on average 52.5% on the western slopes to 42.5% to the east of the ridge (no comparable data available for CRC 1211 station #12). The daily variability of the relative humidity mirrors the daily course of the air temperature (Fig. 6B). Differences
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between east- and west-facing surfaces are most pronounced during daytime, while relative humidities
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of up to 100% were reached during the early morning hours on both flanks of the ridge. In general, the
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night was characterized by a relative humidity of > 75% during the measurement period. The wind gauge placed on top of the peak recorded a maximum wind speed of 13.7 m s-1 (mean: 4.3 m s-1,
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mean daytime 5.8 m s-1).
4.3 Nuclide concentrations, erosion rates and geomorphological parameters
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The reagent blank corrected concentrations of 10Be and 26Al of the amalgamated samples ranged from
blank subtractions for
10
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1.06 × 105 at g-1 to 3.29 × 106 at g-1 and 8.30 × 105 at g-1 to 1.71 × 107 at g-1, respectively. Respective Be and
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Al constituted < 7.2% and < 6.9% of the total number of atoms
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measured. With regard to the two bedrock samples,
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Be concentrations were 7.86 × 104 at g-1
(BA15-008) and 7.28 × 104 at g-1 (BA15-009). However, the lower
10
Be/9Be rate of BA15-009 has
meant a blank subtraction of 19.7%, arguably too high to confidently rely on this data. Nevertheless, the concentration is similar to that obtained from BA15-008 (6.5% 10Be blank subtraction) and fits to what we would expect for the given topographic setting. We therefore present the results of this sample for the sake of completeness but do not include it in later discussion. In general, the concentrations decrease from east to west along the transect. Plotting the 10Be and 26Al results on a complex exposure diagram indicates steady-state erosional conditions within 2σ for all samples except BA15-001, suggesting that within the precision of the measurements there is no evidence the samples have experienced complex burial histories (see supplement). The erosion rates we derive using the online calculator (Balco et al., 2008) are shown in tables 1 and 2. As the study area is spatially 23
Journal Pre-proof small with low relative relief and production rates between sites are similar, we compare our cosmogenic nuclide results with each other using the internal uncertainties reported by the online calculator (see supplement for external uncertainties).
Table 1. Erosion rates and erosional timescales obtained from the amalgamated samples. ID
Location
Elevation
a
10
26
Be erosion
a
10
Al erosion
-1 b
Be Tave
-1 b
Incision age
c
(Myr)d 8 2 2 2 >1 >1 >1 1 2 <6 < 10 <3 2
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(°) (m a.s.l.) (m Myr ) (m Myr ) (Myr) BA15-001 -21.4214 -69.9783 713 3.2 ± 0.3 4.2 ± 0.6 0.3 BA15-002 -21.4229 -69.9766 696 9.0 ± 0.7 10.5 ± 1.4 < 0.1 BA15-003 -21.4242 -69.9767 695 13.4 ± 1.1 15.6 ± 2.7 < 0.1 BA15-004 -21.4243 -69.9812 680 28.8 ± 3.3 27.2 ± 10.6 < 0.1 BA15-005 -21.4240 -69.9830 670 26.5 ± 2.4 26.9 ± 4.0 < 0.1 BA15-006 -21.4257 -69.9825 664 15.0 ± 1.5 17.6 ± 4.2 < 0.1 BA15-007 -21.4237 -69.9820 677 20.7 ± 1.8 21.3 ± 3.8 < 0.1 BA15-010 -21.4227 -69.9834 677 39.6 ± 4.7 34.8 ± 6.9 < 0.1 BA15-012 -21.4141 -69.9725 737 3.7 ± 0.3 4.1 ± 0.7 0.2 BA15-013 -21.4125 -69.9654 734 1.1 ± 0.1 1.1 ± 0.4 0.8 BA15-014 -21.4085 -69.9661 757 1.1 ± 0.1 1.3 ± 0.4 0.8 BA15-016 -21.4146 -69.9671 732 1.8 ± 0.3 2.3 ± 0.7 0.5 BA16-001 -21.4200 -69.9715 704 4.5 ± 0.7 4.9 ± 1.1 0.2 a Catchment-averaged. b LSDn scaling scheme used to derive production rates. All errors are 2σ (internal uncertainty). c For z = 85 cm, assuming steady state erosion on the order of 3-5 × Tave (Lal, 1991). d Approximation based on the mean incision depth and rate of incision.
The catchment-wide erosion rates obtained from the amalgamated samples reflect the overall visible
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incision patterns along the E-W transect (Fig. 7). Accordingly, the lowest 10Be and 26Al rates of around 1-2 m Myr-1, are found in the easternmost catchments (group E2). By contrast, the westernmost catchments erode at up to 39.6 ± 4.7 m Myr-1 (BA15-010), implying an east-west increase of erosion by one order of magnitude within the 2500 m transect. Similar rates are obtained from the two bedrock knickpoints to the west of the W1 catchments (Tab. 2).
Table 2. Erosion rates and erosional timescales obtained from the bedrock samples. ID BA15-008
Location (°) -21.4255
-69.9846
Elevation (m a.s.l.) 607
10
Be erosion (m Myr -1)a 36.3 c
±
5.1
Be Tave (Myr)b < 0.1
c
BA15-009 -21.4258 -69.9857 582 39.5 ± 11.2 a LSDn scaling scheme used to derive production rates. All errors are 2σ (internal uncertainty). b For z = 60 cm, assuming steady state erosion on the order of 3-5 × Tave (Lal, 1991). c Reduced reliability because of large blank subtraction, see text for discussion.
24
10
< 0.1c
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The timescales over which the erosion rates integrate also differ by one order of magnitude between fast and slowly eroding catchments, being ~20 kyr in catchment BA15-010 and about 0.8 Myr for BA15-014. The bedrock erosion integrates over ~20 kyr. The averaging time of the erosion rate measurements provides an estimate of the period over which the measurements themselves are valid, being the time to erode an average of one attenuation length, or in this case a depth of ~85 cm, assuming steady state erosion over time periods on the order of 3-5 × Tave (Lal, 1991). Comparing erosion rates
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integrated over such variable time periods presents a challenge in the assertion of driving factors, which
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might exert control over differing timescales (Schumm and Lichty, 1965). However, as the basins are
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located relatively close to each other, the large-scale environmental factors such as the uplift of the Coastal Cordillera or synoptic climate conditions (sect. 1, 2.2, 2.3) are likely to be very similar, including
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the long-term perspective. In addition, the topography of the western catchments has clearly required
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erosion over timescales much longer than the averaging time of the cosmogenic nuclide measurements. Assuming the erosion rates we derive in the west have been relatively constant over the averaging times
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of the eastern catchments, their comparison is tenable.
25
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Figure 7. Catchment-wide and bedrock 10Be erosion rates (in m Myr-1) and 10Be erosional timescales
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Tave (in brackets; Myr), illustrated on a Pleíades 1B pansharpened multispectral satellite image. The data indicate an approximately 40-fold E-W increase of erosion along a distance of ~2500 m. A N-S
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increase is also observed in groups W2, E1 and E2 (see also Fig. 8). Bedrock contact (Cañón del Loa
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beds – CLB and Caleta Loa complex – Csbcl) and limit of Alto Hospicio gravels (AHG) are represented by white dashed lines. The dashed red lines indicate tectonic displacements as inferred from the DTM analysis and observations made in the field. See text for discussion of the reliability of sample BA15-009 (*).
Apart from the general decrease in W-E direction, a downstream increase of erosion is also observed (Fig 8). Downstream of catchment BA15-004, the bedrock at knickpoint BA15-008 erodes at 36.3 ± 5.1 m Myr1. In catchment BA15-004, eroding at 28.8 ± 3.3 m Myr-1, subcatchment BA15-005 erodes faster (26.5 ± 2.4 m Myr1) than subcatchment BA15-007 (20.7 ± 1.8 m Myr-1), which is located further upstream (Fig. 8A). Both subcatchments are southeast-facing. Subcatchment BA15-006, situated directly opposite to BA15-005 (northwest-facing), erodes at 15.0 ± 1.5 m Myr-1. 26
Journal Pre-proof The east-facing catchments of group W2 are connected by a N-S running channel that drains into the Río Loa canyon (Figs. 3, 8B). The calculated catchment-wide erosion rates decrease in a northerly direction from 13.4 ± 1.1 m Myr-1 (BA15-003) to 9.0 ± 0.7 m Myr-1 (BA15-002) and 3.2 ± 0.3 m Myr-1 (BA15-001). The northernmost catchment of this triplet erodes at similar rates as the western N-S oriented catchments BA16-001 (4.5 ± 0.7 m Myr-1) and BA15-012 (3.7 ± 0.3 m Myr-1) with the former
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being located closer to the Río Loa canyon.
Figure 8. Inner-catchment E-W and N-S increase of erosion rates. Shown are normalized channel
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profiles (Demoulin, 1998) of the master drainage channel of BA15-001 through BA15-003 (group W2; B) and BA15-005 through BA15-007 (group W1; A), which is for the latter case the main channel of catchment BA15-004. The channel profiles extend from the respective channel head to the northern rim of the Loa canyon. Points of confluence and the sampling point of BA15-004 are illustrated by red dots; the locations of knickpoint bedrock samples are indicated by red triangles. The profile in (B) indicates that N-S running channels are influenced by the base level drop caused by the Río Loa incision. Thus, the N-S increase of erosion in the east-facing catchments can possibly be linked to upstream knickpoint migration and the upstream communication of the erosional signal perpendicular to the Río Loa. The inner-catchment variation of southeast- and northwest-facing subcatchment erosion in BA15-004 seems to be affected by other factors as the upstream migration of the erosion signal is impeded by the bedrock knickpoints located downstream of BA15-004 (A). 27
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Although the erosion rates clearly reflect the overall drainage patterns and the geomorphology of the study area, they do not correlate with the respective drainage densities. The majority of catchments show values for Dd between 52 and 67 km km-2 (Tab. 3). Not included into this group are the subcatchments BA15-005 through BA15-007. While the two southeast-facing subcatchments show values above 70 km km-2 for Dd, the opposite case applies to the northwest-facing subcatchment BA15-006 (~42 km km-2). The second-lowest value is reached by the slow-eroding catchment BA15-001. It is also
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both show similar stream lengths per area (~59 km km-2).
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noteworthy that the two fast-eroding catchments to the west of the ridge (BA15-004 and BA15-010)
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Journal Pre-proof Figure 9. Catchment-wide erosion rates plotted against local relief (A), mean incision (B), average catchment slope angle (C), and average catchment hilltop curvature (D). Colored numbers denote the coefficients of determination (R2) of the respective linear correlations. Such correlations (above 0.6) exist for all geomorphological parameters except for drainage densities. The westernmost catchments in (C) seem to have reached a threshold slope angle of ~30°; thus, a correlation is not made for this group. The parameters were derived from the eroding areas that contribute to the rates we measured
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(sect. 3.2). For further discussion, see text.
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A linear correlation of the erosion rates can be observed when related to local relief (R2 = 0.61; Fig. 9).
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However, while the correlation holds if only the western, i.e. catchments of groups W1 and W2, are considered (R 2 = 0.67), the correlation is very weak for the eastern catchments of groups E1 and E2
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(R2 = 0.04). For mean incision depth, taken as surrogate for local relief, the correlation of the entire
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dataset is slightly higher than for the relief (R2 = 0.65). However, if only those catchments are considered that are bounded by stable ridges, including the gypsum-covered ridge between groups W1 and W2, the
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linear correlation increases to R2 = 0.87. The highest values for mean incision (~50 m) are reached for
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the westernmost catchments BA15-004 and BA15-010, indicating a five-fold increase of incision when
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compared to the E1 and E2 catchments (Tab. 3).
Table 3. Values of the geomorphological parameters obtained for the respective catchments. Size
Mean elevation
Dd
Mean slope
Mean CHT
(m2)a
(m a.s.l.)a
(km km-2)a
(m)a
(m)a
(°)a
(m-1)a
BA15-001
107646
713
52.28
66.43
24.82
24.62
0.11
BA15-002
36023
696
66.32
60.88
19.76
25.62
0.12
BA15-003
52562
695
60.55
79.90
23.33
25.77
0.12
BA15-004
162288
680
58.67
113.95
49.22
29.82
0.15
BA15-005
11373
670
79.43
70.50
14.42
31.65
0.14
BA15-006
7711
664
41.94
51.33
14.84
29.05
0.10
BA15-007
20999
677
70.91
75.42
22.18
31.41
0.16
BA15-010
154232
677
58.97
142.00
48.98
30.89
0.16
BA15-012
91605
737
56.07
73.16
7.70
16.60
0.07
BA15-013
76385
734
64.47
41.66
6.43
10.74
0.06
BA15-014
227647
757
56.94
76.96
9.95
12.54
0.06
BA15-016
153007
732
61.75
61.55
5.48
10.28
0.05
ID
Local relief Mean incision
29
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a
BA16-001 50500 Eroding areas.
704
59.55
41.52
9.10
17.79
0.09
A further separate analysis is also reasonable for the correlation of catchment-mean slopes and erosion, where linear correlation is especially strong for the eastern catchments (R2 = 0.84; Fig. 9). It is, however, striking that the mean slopes of the fast-eroding catchments of group W1 reach all values of around 30°. For mean hilltop curvature and erosion rates, linear correlation is high for the entire dataset (R2 = 0.76), with low values for CHT (~0.06-0.09 m-1) obtained for the eastern catchments. In the western portion,
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BA15-001 to -003 and BA15-006 reach mean hilltop convexities of about 0.10 to 0.12 m-1 (Tab. 3). The
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remaining catchments (BA15-004, BA15-005, BA15-007 and BA15-010) show values above 0.14 m-1.
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5 Discussion
5.1 Erosional stages of the study area: E-W and N-S contrasts
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The correlation patterns between the geomorphological parameters and the catchment-wide erosion rates point to the prevalence of two different erosional regimes along the E-W transect. The eastern
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catchments of groups E1 and E2 erode at lowest rates and mostly reach low values for the respective
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parameters (Tabs. 1 and 3). Among the parameters that are derived, drainage density 𝐷𝑑 , defined as the
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quotient of total stream length and catchment area, is commonly used to assess the degree of landscape dissection (e.g. Bierman and Montgomery, 2014; Collins and Bras, 2010). In arid landscapes without a soil-protecting vegetation cover, drainage density is believed to be strongly linked to the availability and variability of erosive precipitation inducing fluvial processes (Collins and Bras, 2010; Moglen et al., 1998). It has been shown that power law relationships between Dd and erosion exist in soil-mantled landscapes (Clubb et al., 2016). However, it appears that observed drainage densities in our study area do not reflect the different magnitude of erosion along the E-W transect. In part, this was caused by the fact that we could only extract two-dimensional stream networks due to poor DTM qualities generated from the drone pictures. In addition, Howard (1997) showed that values for Dd (1-D/2-D) decrease as incision increases in detachment-limited badlands dominated by threshold slopes. The drainage densities derived from our study area do indicate, however, that channel incision is more pronounced on the southeast-facing slopes than on their northwest-facing counterparts in catchment BA15-004. 30
Journal Pre-proof In theory, fluvial incision increases the local relief and hillslope gradients, which have shown to correlate linearly to erosion in low-relief and low-slope catchments and explained the response to a drop of the local base level (Ahnert, 1970; Montgomery and Brandon, 2002). Erosional steady-state can be assigned to those areas where long-term erosion rates are rather uniform and decoupled from relief or slope (Burbank, 2002). The lack of a simple linear correlation between erosion and local relief in the eastern catchments is biased by the general dip of the pre-incision surface towards the south, creating relief due to regional tectonic warping and pre-Middle Miocene alluvium aggradation, rather than being an
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indicator for local channel incision. As catchment BA16-001 has the smallest N-S extent, the dip of the
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surface causes an elevation gradient of about 10-20 m, while almost the entire relief of catchment
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BA15-016 (about 60 m) is created by the surface dip. This bias is to some extent reduced by applying mean incision depth as a surrogate for local relief. Using this parameter, it is possible to compare only
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those catchments that are incising into a stable, diffusive surface, which mostly applies to the E1, E2,
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and with some limitations also to the W1 and W2 catchments. To relate western and eastern catchments in our study area using the mean incision depth, it is an important prerequisite that large parts of the
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bounding ridges – defining the watershed of the respective catchments – have not been worn down
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significantly for timescales exceeding Tave. Based on DTM analysis and observations made in the field we believe this holds for all but the subcatchments BA15-005, BA15-006 and BA15-007. The latter are
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incising into a side slope of BA15-004, and the ridges defining their watersheds have been significantly lowered as a result of sediment removal along the main channel of BA15-004. Thus, the incision depth is significantly affected by the downwearing of the slopes and these subcatchments should be excluded from the correlation in such a small-scale analysis. However, a further constraint appears when this parameter is applied to the eastern catchments, as presumably tectonic movement has tilted the surfaces to the east (Fig. 4). This especially affects the E2 catchments, since incision is low and their E-W extents significant, artificially increasing the incision depth. Nevertheless, because averaging reduces this effect to some extent, they are included into the correlation. In the catchments of group E1, the eastward tilt does not affect the mean incision significantly, as E-W extents of the catchments are comparably low and the incision is generally deeper.
31
Journal Pre-proof Apart from these constraints, the high correlation of mean incision depth with the erosion rates indicates that the study area is adjusting to a drop of the erosional base level. This finding is supported by the linear correlation of hilltop curvature CHT with erosion. As slope and local relief are only reflecting the magnitude of erosional processes until a certain threshold is reached, beyond which a further increase is not possible (e.g. Binnie et al., 2007; Burbank, 2002; Montgomery and Brandon, 2002; von Blanckenburg, 2005), hilltop curvature, as a geomorphic parameter, represents the final stage of slope adjustment to adjacent channel incision in eroding landscapes (Hurst et al., 2012). This parameter is
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applicable to the entire dataset, as it records the slope morphology beyond the limitations of mean slope
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or relief (Hurst et al., 2013; Hurst et al., 2012). Accordingly, a linear correlation between erosion and
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CHT is observed. It should be noted that the kind of limitation of the methodology we applied and which affects the parameters local relief and mean incision could also affect the mean slope and CHT. Headward
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channel erosion in a catchment will impact on the bounding ridge between it and the neighboring
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catchment, affecting the ridge curvature which we included in our analysis (see supplement). However, as those catchments that share a drainage divide erode at similar rates and show similar values for the
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geomorphologic parameters, we assume that this methodological limitation does not change the results
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significantly. Thus, for the catchments of groups E1 and E2 it can be stated that adjustment processes are taking place very slowly in a transport-limited erosional regime (Binnie et al., 2007). Taking into
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account the possible relief of 400-600 m down to the Río Loa, the catchments reflect an early stage of erosional response to base level change. By contrast, the western catchments of group W1 and W2 show relatively high values for erosion and the respective geomorphic parameters. The W1 catchments most likely have reached a threshold angle, indicating that erosion has been decoupled from slope and is primarily governed by mass wasting (Burbank, 2002), a feature that is characteristic for a badland-like topography (Howard, 2009). So far, cosmogenic nuclide based erosion rate estimates have been somewhat spatially biased towards active mountainous regions (Portenga and Bierman, 2011), where the technique has been used to provide empirical evidence for theoretical geomorphological concepts, such as the attainment of threshold topography (e.g. Binnie et al., 2007; DiBiase et al., 2010; Ouimet et al., 2009). The notion of threshold topography, when applied at the broad scale suggests that, on average, hillslope gradients of mountains 32
Journal Pre-proof maintain some relationship with erosion rate that decouples as slopes reach their threshold angles of repose. Hillslopes that attain threshold gradients cannot steepen further and then erode by landsliding in proportion to the rate of incision of the valley trunk stream (Burbank, 2002; Burbank et al., 1996; Montgomery and Brandon, 2002). This has led to predictions of the emergence of ‘threshold topography’ after a particular rate of average erosion is reached, e.g. > ~200 m Myr-1 for catchments > 25 km2 at the eastern margin of the Tibetan Plateau (Ouimet et al., 2009); though the erosion rate and hillslope angle at which threshold topographic processes operate will also be a function of material
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strength (Palumbo et al., 2009; Schmidt and Montgomery, 1995) and climate (Gabet et al., 2004).
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This notion of ‘threshold topography’, or the threshold angles of hillslopes, was initially proposed to
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explain characteristic hillslope angles in soil studies concerned with local-scale slope processes operating on valley side slopes (e.g. Carson and Petley, 1970; see Montgomery, 2001). Hillslope
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modelling studies subsequently incorporated a critical slope gradient value to derive non-linear sediment
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transport laws, assuming processes remain dominantly transport-limited (e.g. Roering et al., 1999; 2007). Howard (1997) applied such modelling to a badland landscape, showing that parallel retreat of
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slopes follows a rapid incision, but also noting the lack of process rate information prevents more
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conclusive interpretations. Here we find that the notion of a limiting, or threshold, slope angle is apparent in our erosion rate data from a badland landscape, despite the relatively slow rates of erosion we
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measure. Nevertheless, the lack of developed gypsum crust cover and the evidence for mass wasting deposits in the channels at our fieldsite would suggest detachment-limited conditions prevail at this location and the reason for the attainment of threshold slope angles under such a low-energy environment likely relates to the lower material strength of the conglomeritic lithology. However, a linear correlation between erosion and local relief does exist in the western catchments, which can be explained by the fact that all catchments of group W1 and W2 except for BA15-005 and -007 have their basin head bounded by the local ridge or a plateau (BA15-006). The gypsum crust cover on the peaks of the ridge and its morphology indicate that it has been providing a long-term stable upper limit of the local relief, which has possibly not been worn down significantly for timescales exceeding Tave. Thus, local relief is only dependent on the magnitude of incision in the main channel of the respective catchments, which is governed by the erosion rates and in this special case a more 33
Journal Pre-proof meaningful parameter than mean slope. A similar argument is valid for the mean incision depth. However, as it is decoupled from the slope angle, erosion in the westernmost catchments is in a more developed stage and detachment-limited. This finding is corroborated by the high hilltop curvatures in these catchments. It is very likely that erosion rates would have been higher if the drainage had not incised into the bedrock a bit further downstream of the W1 catchments, slowing down the upstream migration of knickpoints (Fig. 8). These knickpoints act as local base levels for erosional processes further upstream, such that
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the bedrock erosion rate at BA15-008 (36.3 ± 5.1 m Myr-1) sets the pace of erosion in catchment
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BA15-004 (catchment-mean erosion rate of 28.8 ± 3.3 m Myr-1), which is located ~100 m further
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upstream of the knickpoint. If the bedrock knickpoints would not have acted as local erosional base levels, the ridge separating W1 and W2 catchments might have been already eroded away and the W2
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catchments captured. Elsewhere, rapid incision and subsequent base level stability has been found to be
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an important factor for the formation and preservation of badland topography (Howard, 1997). There is no indication that the W1 streams are significantly capturing the W2 watersheds, although its general
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Whipple, 2018).
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morphology indicates than an eastward shifting of the ridge probably has been initiated (cf. Forte and
Apart from the W-E contrast of erosion, manifested in detachment-limited vs. transport-limited
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erosional regimes, the results also highlight the strong dependency of erosion on the spatial distance to the Loa canyon. In all catchments of groups W2, E1 and E2 the northern catchments erode at lower rates than those that are located closer to the canyon. This is again most likely due to upstream knickpoint and/or knickzone migration perpendicular to the Río Loa (Fig. 8). However, given the low erosion rates in group E2, the communication of the erosional signal of the latest Río Loa incision has certainly not reached these catchments yet. The fact that the catchments BA15-013, -014 and those of group E1 are draining into larger channels, which in turn drain the elevated areas to the north, seems not to have any significant impact on the erosion rates. As we did not observe knickpoints on the confluence points of the analyzed catchments with the respective master channels, we argue that the W-E decrease in erosion also affects the elevated areas to the north.
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Journal Pre-proof 5.2 Mechanisms causing the E-W erosion gradient across the ridge Special attention should be paid to a catchment that seems to contradict the general trends observed for our dataset. Subcatchment BA15-006 shows the lowest erosion rates of all W1 catchments (15.0 ± 1.5 m Myr-1), although it is located opposite to subcatchment BA15-005, eroding at 26.5 ± 2.4 m Myr-1. Furthermore, its drainage density is the lowest of the entire dataset. This could be, however, attributed to the high sensitivity of that parameter to minor changes in the stream network in small catchments. That such catchments sometimes underestimate Dd can be inferred from the datasets
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presented by Clubb et al. (2016). By contrast, the other two small subcatchments BA15-005 and
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BA15-007 have the highest drainage densities of all catchments. In addition to that, all other
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geomorphological parameters for BA15-006 are the smallest within group W1 (Tab. 3). Thus, the anomalous values obtained for this particular catchment are most likely caused by special conditions
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that can be found in its catchment head. This area can be described as a small, uneven plateau, covered
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by a gypsum-rich layer of varying thickness of 0 to > 30 cm and/or millimeter to centimeter-thick crusts forming centimeter-scale polygons (Fig. 2F). Gypsum crust, of the type observed in the eastern
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catchments, is almost absent, and the surface further covered by clasts predominantly of centimeter to
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decimeter size. Despite of the missing gypsum crust on this plateau (which has been excluded from the derivation of the geomorphic parameters, sect. 3.2), the surface does not show signs of significant fluvial
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erosion towards BA15-006. The unevenness of the terrain seems to be related to a weathering and/or deflation of former gypsum polygons, which can still be traced on the surface. The shallow subsurface is salt-cemented (mostly halite), which has been observed to be a wide-spread phenomenon in the study area (Fig. 10). At some places on top of the plateau, salt-cemented sediment forms a competent cap which is pervaded by few centimeter-wide and > 1 m deep cracks. At such a feature we found a small salt stalactite with the morphology of a small droplet (~1.5 cm along the long axis; Fig. 10F), indicating long-term, slow but steady downward (per descensum sensu Watson, 1985) transport of salts, possibly due to diurnal fog-precipitation. The salt-rich layer is being eroded at the channel heads of the subcatchments incising into the plateau, most likely impeding the erosion to keep the pace of the overall erosion in BA15-004.
35
Journal Pre-proof Moving across the ridge that separates the W1 and W2 catchments, the salt-cemented sediments are additionally covered by a decimeter-thick gypsum crust (costra, sensu Ericksen, 1981). Diurnal fogmoisture cannot permeate this gypsum cover to reach the halite (Ericksen, 1981) and leach it from the uppermost sediment layers as we observed it in the western catchments. The gypsum crust itself is a formidable protection of the surface against individual, presumably brief, rain events. It has a high physical cohesion, and gypsum, while being semi-soluble in water (Goudie and Viles, 2015), is kinetically sluggish to dissolve (Warren, 2016). Lichens are occasionally macroscopically visible but,
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however, are rare. The physical cohesion of the gypsum crust in the study is therefore presumably
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dominated by abiotic factors. Commensurate with its physical properties, the gypsum crust-covered
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areas show few signs of fluvial erosion. Visible traces of run-off erosion, such as accumulations of detrital material, are mostly limited to the boundaries of polygonal structures in the crust. Gypsum-
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covered hillslopes appear to erode from the bottom, in response to channel incision/back cutting.
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Eroding slopes have a pronounced break in slope where the crust terminates (the lower slopes are commonly at the angle of repose). These features are best preserved in the transitional group of
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catchments just east of the ridge (W2; Fig. 7 and supplement).
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With key physical parameters being similar across the study area (mean elevation, substrate, presumably long-term mean annual precipitation from rain) the presence of gypsum crust and the ‘health’ of the
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halite-cement in the top-layer of the sediment appears to be a good predictor of the order of magnitude of erosion rates; 15-40 m Myr-1 in areas without crust and leached halite, 1-2 m Myr-1 with an intact gypsum crust cover (Fig. 7).
36
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Figure 10. Pictures taken in the field illustrating some characteristic features of the surfaces in the study area. Salt outcrops along the northern slopes of the Loa canyon indicate the widespread presence of subsurface salt-cementation (white arrow in A; red arrow pointing at the local ridge separating W1 and W2 catchments). In the eastern parts, the salt-cemented gravels (III in B) are covered by gypsum-rich layers (II in B) and gypsum crust (I in B). Note that the crust is cohesive enough to form a topographic hanging edge (Picture taken in BA15-003). Thick, encrusted gypsum cover (costra sensu Ericksen, 1981) is abundant on the crest of the main ridge (C), indicating long-term erosional quiescence and surface 37
Journal Pre-proof stability. The high abundance of disintegrated clasts on the surfaces indicates long-term surface and/or subsurface stability (D). This statement is further supported by a fossil-bearing limestone, whose exposed east-facing surface has been differentially weathered, exposing shell fragments (white arrow in E). The weathered surfaces are restricted to the part of the rock sticking out of the crust (to the right of dashed white line); the east-facing area is not affected by weathering. The downward migration of salts formed a small (~1.5 cm) stalactite in BA15-006 (F).
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The factors that lead to the formation and/or preservation of the gypsum crust are therefore crucial in
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forming the landscape as we find it at the present day. Given the subdued topography and small lateral
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extent of the study area, microclimatic conditions are probably key factors in determining the spatial presence or absence of gypsum crusts. Long-term stable and arid to hyperarid conditions must prevail
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for a gypsum layer of a significant thickness to form as the precipitation of gypsum crystals is prevented
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in moist soils (Heine and Walter, 1996; Retallack and Huang, 2010). The coastal fog has been identified as an important source for the transfer of pedogenic salts like Ca and S from the Pacific Ocean and their
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subsequent per descensum deposition in the Coastal Cordillera at altitudes below 1200 m a.s.l. (Rech et
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al., 2003). As minor abundances of gypsum crust to the west of the local ridge on top of the plateau above BA15-006 and on a small slope within BA15-004 indicate (Fig. 2), gypsum deposition must have
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occurred in the entire study area at some point but was later removed to the west of the ridge. Microclimatically, there are three important differences between the east- and west-facing drainage bounded by the ridge (Fig. 11). Firstly, the W1 catchments face the strong afternoon inland winds (sect. 2.3) which can exceed 13 m s-1 as measured by the wind gauge placed on the peak between the W1 and W2 catchments. Wind exposure could be a limiting factor for the formation and preservation of gypsum in our study area, as extreme wind scour could impede its accumulation (Miller, 2008) or erode these surfaces (Eckardt and Spiro, 1999).
38
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Figure 11. Sketch illustrating possible mechanisms and microclimatic conditions that could be
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responsible for the E-W contrast of erosion in along the main ridge. The western slopes face strong
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daytime winds, cooling down the surface and possibly impeding atmospheric gypsum deposition. Furthermore, the steep slopes could act as a kind of natural fog collector, promoting long-term
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downward migration of salts which could occasionally release clasts, initiating downward chain
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reactions of clasts following gravity. The crust of the east-facing slopes remains comparably unaffected. However, the fact that chuca is mostly absent also points to long-term quiescence of gypsum deposition.
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During a rare rain event, the cemented subsurface of the west-facing slopes, although likely stabilizing the slopes, might promote surface runoff, while to the east the gypsum cover promotes infiltration.
Secondly, almost the entire moisture supply reaching the study area originates from the Pacific Ocean (sect. 2.3) and it is unlikely that the present-day advection patterns differ significantly from those prevailing over the past millions of years (sect. 1). While it is reasonable to assume that the rare rain events are spatially uniform across the study area (due to similar and low elevation as compared to cloud base of rain-clouds), diurnal moisture deposition from fog (i.e. ground-hugging clouds) will not be uniform. Fog-moisture preferentially deposits on wind-facing surfaces, in our case the west-facing slopes, with accompanying wind/fog-shadow effects. Furthermore, dew yields are highly sensitive to altitude (see sect. 2.3). At the Alto Patache Climate station, located ~70 km to the north of our study 39
Journal Pre-proof area at an elevation of 850 m a.s.l., Cereceda et al. (2008a) measured an average precipitation of 7.0 l m-2 day-1 from fog between 1998 and 2005. The yield of fog water is highly focused on the winter/spring months, peaking in September (exceeding 20 l m-2 day-1), while lowest yields are recorded for February and March (Cereceda et al., 2008a). We hypothesize that the west-facing slopes along the moderate ridge that separates areas of different erosional regime act as a fog collector. Significant aspect-related contrasts in fog precipitation are supported by observations in catchment BA15-002 to the east of the main ridge. Situated on a ~25° slope, 2 m below the hilltop, we found a 30x30x15 cm
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fossil-bearing limestone sticking out of the gypsum crust. The west-facing side of the rock was strongly
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weathered, with shell fragments protruding about 2 cm from the weathered rock surface. The eastern
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side lacked weathering features, indicating long-term surface stability and that moisture deposition is sufficient to differentially dissolve limestone. Fog-precipitation is evidently sufficient to mobilize
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surface halite (halite stalactite, Fig. 10F). It has been shown that in arid desert environments, the highly
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soluble halite (which is hygroscopic at relative humidities > 75%, Gu et al., 2017; commonly exceeded in our study area, see also sect. 4.2) is segregated from the sulphates and washed down during periods
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of higher moisture supply (Eckardt and Spiro, 1999), cementing the subsurface. Such a mechanism has
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already been reported from the hyperarid deserts of Tunisia and Namibia, resulting in the formation of two-tiered crusts (Watson, 1985; Fig. 11). Apart from stabilizing the slope (cf. Watson, 1985; Wilcox
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et al., 2016), the cementation promotes surface runoff during rain events (Yang et al., 2016). Thus, on hillslopes on which halite is remobilized into the shallow subsurface (catchment group W1; Fig. 7), erosion of the then uncemented surface material is facilitated by its lack of cohesion and increased surface runoff. Thirdly, fog-moisture is exposed to sunlight after sunrise to the east of the ridge while in west-facing catchments the surfaces are shaded during the early morning hours. Water cannot infiltrate as deep as on the west-facing surfaces, where evaporation rates pick up later. Thus, on the long-term perspective, the east-facing slopes tend to be warmer/drier, likely promoting the formation of near-surface pedogenic alabastrine gypsum through increased evaporation (e.g. Eckardt et al., 2001; Watson, 1985). Taken together, the erosional gradient that we can observe today might have been promoted by a weakly developed surface crust to the west of the ridge, which is likely also subject to wind erosion. Especially 40
Journal Pre-proof on the flat areas in the eastern portion of the study area, the crust-covered surfaces and gypsum-rich subsurface layers likely soak up the majority of precipitation during a rare rain event (Ericksen, 1981). Such a behavior has already been observed for flat to gently dipping surfaces in northern Chile during the 2015 rain event (Jordan et al., 2015). Davis et al. (2010) have shown that a single rainfall event of 3.55 mm, as well as a rainfall of 7.14 mm cumulated over 5 consecutive days, did not produce any surface runoff or ponding on a flat gypsum-covered surface in the Yungay area. By contrast, the possibly weakly developed and wind-eroded crust in the west-facing catchments could have been more easily
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fluvially eroded, whereby overland flow is likely being accelerated when the erosion has reached the
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subsurface due to its strong cementation reducing the infiltration rates (Yang et al., 2016). Regularly-
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occurring moistening of the exposed subsurface could lead to a slow but steady downwash of the halite and downward shift of the focus of cementation. On threshold slopes, this could occasionally release
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previously cemented clasts which would move downslope following gravitation, possibly hitting other
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clasts, and initiating a downslope chain reaction. However, the subsurface salt layers do stabilize the slopes (cf. Wilcox et al., 2016) and slow down their incision (BA15-006). Incision eventually
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5.3 Evolution of drainage
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-006 (Figs. 2, 10, 11).
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undermines any remnants of intact crust covering the hilltops, as observed in catchments BA15-003 and
The analysis of tectonic and geological features (sect. 4.1) reveals that N-S striking (half-)graben structures exist in our study area, covered by the alluvial gravels. The bedrock outcrops as mapped along the Loa canyon indicate that the area around the ridge that is influencing the erosional behavior is placed on a paleohorst structure. This structure possibly underwent relative uplift after it was covered by the alluvial sediments, as gravels belonging to the CLB are found at higher elevations than the adjacent younger AHG pre-incision surface, thereby forming the present-day ridge separating the western catchments (Fig. 4). Tectonic movement after the deposition of the CLB and before the aggradation of the Alto Hospicio gravels has been reported for the Loa fault (Vásquez et al., 2018), which could have led to a canyon-parallel uplift of the study area on the order of tens of meters or less. However, a more important role could have been played by N-S oriented, meter-scaled tectonic vertical displacements 41
Journal Pre-proof and tilting in the eastern section of the study area, associated with the assumed N-S fault scarp (B in Fig. 4) and/or along hypothetically existing N-S faults that have produced the serrated profile shape of the paleohorst (A in Fig. 4). This view is supported by a ~300 m long, horizontally embedded layer of well-recognizable gravels, which are exposed along the Loa canyon below the E1 catchments and contact the eastern flank of the assumed paleo-horst. However, the easternmost ~20 m of this strata are dislocated downwards by about 90 m, roughly below the assumed N-S scarp between the E1 and E2 catchments. Further indications of fault movement, which also occurred after the deposition of the AHG
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unit, can be inferred from slope analysis and are also traceable in the field (Fig. 1; supplement). We
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found no indication that post-Miocene tectonic movement has influenced the overall drainage pattern
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and channel evolution. The tectonic influence on drainage, i.e. the tilting and uplift of parts of an Early Miocene depositional surface, apparently occurred mostly prior to the incision of the catchments
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investigated.
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It is likely that the N-S running eastern catchments of group E2 represent initial drainage conditions, incising into the Early Miocene depositional surface (sect. 2.2). Dividing the mean incision depths by
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the erosional time scales of the respective catchments might give a rough approximation about the
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incision history in our study area (Tab. 1). This approach is highly speculative, as long-term steady state erosion is assumed beyond the timescales the erosion rates integrate over. Furthermore, such a
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calculation is not applicable for the E2 catchments, as the mean incision values obtained from the DTM analysis is biased (sect. 5.1). If we assume that catchment BA15-001 (W2) reflects long-term steady erosional conditions of a pre-Río Loa incision landscape, the ridge in which the catchment incises into could have been formed, or incision could have commenced, around 8 Ma (Tab. 1). This age is in the range of the time-windows given for the uplift of the Chipana-Adamito block and the intensification of tectonic activity along the splay fault to the east (Ritter et al., 2018b). Furthermore, the onset of incision of large paleochannels into alluvial fan surfaces on the opposite side of the canyon has been dated to ~9 Ma by cosmogenic exposure dating, and fluvial incision in these paleochannels ended latest by 2 Ma due to tectonic truncation upstream (Ritter et al., 2018b). Assuming that these paleochannels continued westwards, probably along the trace of the present-day Río Loa Canon, their incision could have provided a significant base level drop on the order of > 20-30 m for all catchments in our study area at 42
Journal Pre-proof some time between 9 and 2 Myr ago. This pre-Río Loa incision provided more than the required base level drop for the eastern catchment groups (E1 and E2) to incise their present-day thalwegs. Furthermore, it provided most of the required base level drop for catchment group W2 and about 50% of the base level drop required for catchment group W1. Latest with the incision of the Rio Loa Canyon 250-300 kyr ago (Ritter et al., 2018b; Vásquez et al., 2018), the latter received more than the required base level lowering to form their present-day thalwegs. The reconstructed time required to incise channels to their present depth, assuming the current erosion
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rates as being constant through time, is on the order of 1-2 Myr for catchments groups W1, W2 and E1.
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This could indicate intensified erosion during the Quaternary or the base level lowering provided by the
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pre-Río Loa drainage. However, this chronological reconstruction remains speculative and is not entirely consistent, since remnants of gypsum crust cover are present on hillslopes of all catchment groups. The
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gypsum crusts on the hillslopes are indicative of a polyphase (at least two-phase) incision of the
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channels. A significant period of hillslope stability is required to accumulate decimeter-scale gypsum crusts. Deriving the time required for the formation for such a gypsum layer is not straightforward. Rech
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et al. (2003) estimated that in the westernmost parts of Coastal Cordillera and at altitudes of about 700 m,
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about 1 m of soil gypsum could be accumulated on a gently sloping surface within ~1.9 Myr. For the Namib Desert, it has been reported that clearly visible gypsum crystals are formed within 10 kyr, while
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the formation of mature gypcretes of about 1 m thickness possibly takes around 800 kyr (Heine and Walter, 1996). While the corresponding accumulation rates are highly dependent on local conditions, these numbers indicate that long timespans are required to form a thick gypsum cover. We infer that the hillslope-response to channel incision in the study area had at least one long-term hiatus on the order of several 100 kyr or more.
6 Conclusion The topography of the study area reflects a long-term erosional gradient between a detachment-limited regime of relatively rapidly eroding drainage systems in the western portion and a transport-limited regime of extremely slow erosion in the eastern section. Cosmogenic nuclide-derived erosion rates increase more than 40-fold from east to west over a distance of about 2.5 km. The erosion regime and 43
Journal Pre-proof rates appear to be influenced by the presence and state of a gypsum cover and/or the state of the pervasive halite-cement in the otherwise unconsolidated continental sediments into which the drainage has incised. Contrasts in the state of the gypsum cover and halite cement are driven by microclimatic conditions: (1) directional erosion by strong westerly winds, (2) aspect-dependent fog precipitation and (3) aspectdependent evapotranspiration. Tectonic movements were likely important to create the modest topography (i.e. a 20-30 m high N-S ridge) that spatially coincides with the transition of erosion regimes and state of the ground cover (gypsum crust). The functional relationship of the ridge to the fluvial
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evolution of the study area is that of a wind-break for areas to the east and increased fog precipitation
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for areas to the west. In general, the state of the ground cover – and thus the erosional regime – appears
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to be predominantly governed by abiotic factors.
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Acknowledgements
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Thanks go to the field campaign teams in 2015 and 2016 and to all colleagues and students who helped collecting and preparing the samples. Elena Voronina and Tomasz Goral made a great job processing
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the samples in the lab. A very special thank goes to Diego Jaldin of the Universidad Catolica del Norte
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(UCN), who was a more than valuable help during the fieldwork conducted in October 2018 and who provided important scientific inputs. Moreover, we would like to thank Eduardo Campos and colleagues
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at the UCN in Antofagasta for their support. Finally, we thank two anonymous reviewers whose suggestions and comments have greatly improved this manuscript. The research was part of a pilot study for a coordinated research centre funded by the Deutsche Forschungsgemeinschaft (DFG, German Research Foundation) – Projektnummer 268236062 – SFB 1211.
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Journal Pre-proof Highlights We investigate erosion patterns in the hyperarid core of the Atacama Desert
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Erosion rates increase steadily by one order of magnitude within 2.5 km
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DTM analysis reveals the existence of two different erosional regimes
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The presence of gypcrete in the study area is linked to microclimatic conditions
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Gypsum surface cover affects surface processes
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Journal Pre-proof Declaration of interests
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☒ The authors declare that they have no known competing financial interests or personal relationships that could have appeared to influence the work reported in this paper.
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