Publisher: GSA
Journal: GEOL: Geology
Article ID: G33829
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Drainage capture and discharge variations driven by
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glaciation in the Southern Alps, New Zealand
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Ann V. Rowan1*, Mitchell A. Plummer2, Simon H. Brocklehurst1, Merren A. Jones1,
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and David M. Schultz1
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1
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Manchester M13 9PL, UK
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2
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*Current address: Institute of Geography and Earth Sciences, Aberystwyth University,
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Aberystwyth SY23 3DB, UK.
10
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School of Earth, Atmospheric and Environmental Sciences, University of Manchester,
Idaho National Laboratory, Idaho Falls, Idaho 83415-2107, USA
ABSTRACT
Sediment flux in proglacial fluvial settings is primarily controlled by discharge,
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which usually varies predictably over a glacial–interglacial cycle. However, glaciers can
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flow against the topographic gradient to cross drainage divides, reshaping fluvial
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drainage networks and dramatically altering discharge. In turn, these variations in
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discharge will be recorded by proglacial stratigraphy. Glacial-drainage capture often
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occurs in alpine environments where ice caps straddle range divides, and more subtly
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where shallow drainage divides cross valley floors. We investigate discharge variations
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resulting from glacial-drainage capture over the past 40 ka for the adjacent Ashburton,
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Rangitata, and Rakaia basins in the Southern Alps, New Zealand. Although glacial-
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drainage capture has previously been inferred in the range, our numerical glacier model
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provides the first quantitative demonstration that this process drives larger variations in
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discharge for a longer duration than those that occur due to climate change alone. During
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the Last Glacial Maximum, the effective drainage area of the Ashburton catchment
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increased to 160% of the interglacial value with drainage capture, driving an increase in
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discharge exceeding that resulting from glacier recession. Glacial-drainage capture is
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distinct from traditional (base level-driven) drainage capture and is often unrecognized in
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proglacial deposits, complicating interpretation of the sedimentary record of climate
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change.
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INTRODUCTION
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Over glacial–interglacial time scales, rates of climate-driven erosion and sediment
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transport scale with drainage area (e.g., Brocklehurst and Whipple, 2007), leading to the
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expectation that drainage capture will affect sediment transport and be recorded by
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proglacial stratigraphy. However, understanding is limited of the timing of sediment
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transport during the glacial–interglacial cycle (e.g., Dühnforth et al., 2008; Shulmeister et
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al., 2010). At thousand-year time scales, greater than that taken to reach a hypothetical
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steady-state (Cuffey and Paterson, 2010), glacial-drainage capture can change the
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effective drainage area of adjacent catchments resulting in variations in discharge. At
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hundred-year time scales, when glaciers are at a transient state of adjusting to a change in
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climate, water is released by or stored within the ice mass during recession and advance
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(Dühnforth et al., 2008).
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We investigate discharge variations resulting from variations in ice extent—both
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by glacial-drainage capture and advance or recession—over the past 40 k.y. for three
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adjacent catchments in the Southern Alps, New Zealand. The sedimentary record in New
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Zealand is an important archive of Southern Hemisphere climate change (Newnham et
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al., 1999), as the Southern Alps underwent multiple, extensive glaciations during the late
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Quaternary (e.g., Suggate, 1990). Gravel-bed braided rivers constructed a series of
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prograding, coalesced alluvial fans in three catchments—the Rakaia, Ashburton, and
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Rangitata—which are exposed at the coastal cliff of the Canterbury Plains (Fig. 1).
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Luminescence dating indicates that the majority of the visible stratigraphy was deposited
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during the Last Glacial Maximum (LGM, 24–18 ka) and a previous glacial maximum at
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37–31 ka (Rowan et al., 2012) (Fig. 1D). Are the Canterbury sediments primarily a
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record of ice advance and recession within the current catchments, or has glacial-drainage
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capture has a substantial impact on sediment flux and deposition?
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GEOMORPHIC EVIDENCE
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The modern Ashburton catchment has a significantly smaller drainage area (1239
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km2) than both the Rangitata (1549 km2) and Rakaia (2372 km2) (Fig. 1A). However, the
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Ashburton proglacial fan appears to represent a greater amount of vertical deposition than
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the Rangitata (~18 m cliff thickness compared to ~6 m) (Fig. 1D). Furthermore, the
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stratigraphy is subdivided into depositional storeys, of which there are fewer in the
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Rangitata section than those representing the Ashburton and Rakaia (Leckie, 2003;
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Rowan et al., 2012). The headwater geomorphology indicates that transfluent ice flowed
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from both the Rakaia and Rangitata into the Ashburton, dramatically increasing the
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effective drainage area of this catchment during glacials (Mabin, 1980; Evans, 2008;
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Shulmeister et al., 2010). There are two key regions of glacier transfluence (Figs. 1A and
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1C). First, southward-dipping lateral moraines at Prospect Hill and on either side of the
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Cameron Valley confirm that glacial occupancy reversed the Lake Stream tributary to
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flow south into the Ashburton (Pugh, 2008). Second, at Lake Clearwater, where Potts
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Stream currently flows into the Rangitata, part of the Pleistocene Rangitata Glacier
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diverted into the Ashburton, reversed the flow of the lower Potts Stream Glacier, and
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constructed extensive moraines and outwash fans (Mabin, 1980; Evans, 2008). The
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Canterbury bedrock is regionally homogeneous greywacke (Leckie, 2003) removing a
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lithologic control on drainage patterns. Uplift-driven drainage reversal seems highly
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unlikely. Specifically, reversal of Lake Stream by postglacial fluvial incision would have
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required ~200 m of incision within ~20 k.y., a mean incision rate an order of magnitude
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higher than the local rock uplift rate (Herman et al., 2009).
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METHODS
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To test the hypothesis that glacial-drainage capture resulted in discharge
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variations of sufficient magnitude to be reflected in the stratigraphic record, we used a
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glacier model (Plummer and Phillips, 2003) to simulate ice volumes since 40 ka.
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Advance and recession of mountain glaciers are controlled primarily by changes in
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temperature and mesoscale meteorology (Rother and Shulmeister, 2006; Anderson and
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Mackintosh, 2006). High precipitation levels in the Southern Alps increase glacier
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sensitivity to temperature (Golledge et al., 2012), which is the primary control on glacier
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mass balance (Anderson and Mackintosh, 2006). Previous glacier modeling (Golledge et
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al., 2012), climate modeling (Drost et al., 2007) and sea-surface temperature (SST) proxy
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records (Barrows et al., 2007) demonstrate that LGM temperatures in the eastern South
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Island were 6–8 °C cooler than present with little change in precipitation amount. We
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consider the uncertainty associated with LGM precipitation variability as within the
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overall uncertainty assigned to the relationship between simulated glacier extent and the
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primary driving force for glacial advance and recession (i.e., temperature change).
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The glacier model was implemented in a series of experiments that varied mean
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annual temperature from 8 to +2 °C relative to modern conditions in 0.5 °C increments
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(hereafter referred to as temperature change, T) to simulate steady-state ice volume and
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flow. We tested all available rainfall distributions for the Southern Alps and found the
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New Zealand National Institute of Water and Atmospheric Research Ltd. (NIWA) data
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(Tait et al., 2006) to be most appropriate (see the GSA Data Repository1). The ice
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volumes calculated for each T were given a chronology using the timing of changes in
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temperature relative to modern conditions indicated by SST data (Barrows et al., 2007).
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Simulations to describe steady-state at different T are achieved via transient ice-flow
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calculations that constrain the time-dependent transfer of mass through each glacier, the
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end point of which is when the system is balanced. Modeled glaciers for different T
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were compared to mapped terminal and lateral moraines (Mabin, 1980; Evans, 2008;
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Pugh, 2008; Shulmeister et al., 2010), and equilibrium-line altitude (ELA) depressions
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were calculated via the accumulation-area ratio method (Porter, 1975).
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The glacier model is two-dimensional in the quasi-horizontal plane and uses an
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energy balance calculation to drive ice flow. Temperature was calculated using a lapse
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rate of –6 °C km-1. The ice-flow model provides the primary control on discharge
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variations associated with glaciation. At steady state, discharge is equal to precipitation
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and scales with effective drainage area, modulated by the non-uniform precipitation
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distribution. For glacial-drainage capture to affect proglacial discharge requires that
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either subglacial meltwater flow follows the direction of ice flow across catchment
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boundaries or glacial melt occurs downstream of transfluence. Effective drainage areas
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were defined from subglacial meltwater flow directions for a given T. Above the ELA,
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minimal melting is assumed. Below the ELA, the direction of subglacial meltwater flow
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(the hydraulic potential) is controlled by the surface slope of the ice rather than the
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topographic surface at the base of the glacier. The magnitude of the reverse bed slope
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must be greater than the magnitude of the glacier surface slope by a factor of 11 for
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meltwater to not follow the glacier flow direction (Shreve, 1972) (a description of this
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model is provided in the Data Repository).
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Glacier advance reduces discharge as water is taken into storage as ice, whereas
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recession produces brief discharge peaks as extra meltwater is released. The change in
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discharge volume decays exponentially over the response time (Cuffey and Paterson,
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2010). Transient ice-flow calculations show that steady-state is reached within 400 yr of a
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specified perturbation, consistent with the response time estimated from analytical
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solutions (Jóhannesson et al., 1989) and that measured for the modern Tasman Glacier
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(20–200 yr) (Herman et al., 2011). The transient data provide a means of estimating the
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change in discharge that would occur over short (hundred-year) time scales. The
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discharge resulting from the transient glacier response to climate change was quantified
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from the change in balance within the effective drainage area for a given T, and
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combined with steady-state discharge for the duration of the glacier response to climate
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change. Sediment flux was estimated using the power-law relationship derived for Ivory
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Glacier in the central Southern Alps (Hicks et al., 1990).
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RESULTS AND DISCUSSION
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Simulated glaciers provided a good estimate of mapped ice extents (Fig. 2),
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although we note a 0.25 °C offset between T required for the Rakaia to reach LGM
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limits (at T = 6.5 °C) relative to the Ashburton and Rangitata Glaciers (at T = 6.25
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°C) (quantification of the model-data fit is given in the Data Repository). Based on
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modeled ice extents, T of at least 5 °C (Fig. 2A) was required for glacial-drainage
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capture at Lake Clearwater (Fig. 2B), where part of the Pleistocene Rangitata Glacier
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diverted into the Ashburton (Fig. 3). With T = 5.5 °C (Fig. 2C), a lobe of the Rakaia
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Glacier diverted to the south at Prospect Hill through Lake Stream, reversing the flow of
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the lower Cameron Glacier (Fig. 1A). Under LGM conditions (T > 6.0 °C), ice
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occupying Lake Stream blocked further ice flow southward and the Rakaia Glacier
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reverted to flow exclusively along the Rakaia Valley, effectively “switching off” drainage
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capture (Fig. 2D). Simultaneously, a much larger volume of ice diverted into the
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Ashburton at Lake Clearwater (Fig. 3), dramatically increasing effective drainage area at
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the expense of the Rangitata. Animations of model output illustrating glacial-drainage
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capture are presented in the Data Repository. LGM drainage capture increased the
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effective drainage area of the Ashburton relative to the modern drainage area from 1239
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km2 to 1981 km2 (to 160% of the modern drainage area), at the expense of the Rangitata,
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which decreased from 1549 km2 to 975 km2 (to 63%), making the Ashburton the larger
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basin (Fig. 4B). The Rakaia catchment had a slight decrease in effective drainage area
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from 2372 km2 to 2204 km2 (to 93%), but was the largest basin through the LGM and
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postglacial.
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Recession from the LGM was rapid: the Rakaia terminus receded from the range
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front to Prospect Hill with warming (from T = 6.5 °C to 3 °C) (Fig. 2) indicated by
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SST to have occurred within four thousand years of the LGM peak (Barrows et al., 2007)
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and in agreement with results from terrestrial cosmogenic nuclide dating of moraines in
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the Rakaia Valley (Shulmeister et al., 2010) when calibrated using the local production
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rate (Putnam et al., 2010). Moreover, relatively minor climate variations drove dramatic
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drainage capture during the glacial: oscillations in temperature of 0.5 °C occurred over
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~500 yr (Barrows et al., 2007) (Fig. 4A), driving variations in effective drainage area of
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~20% between the Ashburton and Rangitata. Temperature change from 4.5 °C to 6.5
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°C drove transfluent ice flow at Lake Clearwater (Fig. 3). The domains captured by the
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Ashburton include areas close to the main drainage divide where precipitation levels are
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highest, so the relationship between drainage area and discharge is not linear. Therefore,
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discharge from the Ashburton increased more rapidly than effective drainage area during
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the LGM to 218% of modern discharge, whereas discharge from the Rangitata and
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Rakaia decreased to 53% and 95% (Fig. 4C).
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Results for calculations of the transient glacier response to temperature change
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indicate short-lived peaks in discharge corresponding to the reduction in glacier volume
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during recession. During larger advances, discharge was effectively reduced to zero for
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short intervals (Fig. 4E). However, precipitation falling as rain would not contribute to
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glacier mass gain, and, during advances seasonal discharge due to melting was still
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considerable as indicated by outwash deposits (Shulmeister et al., 2010). While the
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magnitude of peaks produced by the post-LGM loss of ice mass was greater than
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discharge variations resulting from drainage capture (Fig. 4E), the duration of discharge
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variations due to the transient response did not exceed 400 yr and the volume of water
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released during recession was less. When integrated over the LGM and the post-LGM
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discharge peak (24–15.5 ka), the Ashburton discharge is nearly doubled relative to the
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interglacial value due to drainage capture (from 1.3 to 2.2 × 109 m3 a-1), whereas the
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increase in discharge due to loss of ice volume was 112% (to 1.4 × 109 m3 a-1) (Fig. 4E).
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Hence, over longer periods than the response time, glacial-drainage capture exerts a
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greater control on the magnitude of discharge variations than recession.
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Sediment production rates in the Southern Alps are amongst the highest on Earth
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(Hales and Roering, 2005). Therefore, if the proglacial rivers are transport-limited, one
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would expect glacial-drainage capture, rather than recession, to control sediment
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transport in the Ashburton and Rangitata. The existing geochronology indicates that
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deposition of the fluvial sediments at the coast occurred from 40 to 18 ka. Over this
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interval, the total sediment flux to the Canterbury Plains was 26% from the Ashburton
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and 20% from the Rangitata; the remainder was sourced from the Rakaia (54%) (Fig.
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4D). However, if we integrate sediment volumes for the LGM (24–18 ka), the
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contribution from the Ashburton (29%) outweighs that from the Rangitata (18%),
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although an opposing trend for the penultimate glacial maximum (37–31 ka) gives a
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greater total sediment flux from the Rangitata (27%) than the Ashburton (17%). When
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LGM sediments were deposited at the modern coastline (Rowan et al., 2012), aggradation
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rates in the Ashburton were higher than during the previous glacial maximum due to
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drainage capture, as observed by the greater sedimentary thickness at the coast in this
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catchment relative to the Rangitata (Fig. 1D). The lesser sedimentary thickness in the
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Rangitata represents either deposition during only the penultimate glacial maximum
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when drainage capture was less extensive or low rates of sediment flux coincident with
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two phases of deposition in the Ashburton and Rakaia.
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CONCLUSION
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Glacial-drainage capture resulted in a sustained increase in discharge from the
Ashburton during the LGM, which outweighed the short-lived increase in discharge due
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to glacier recession post-LGM, and maintained rates of sediment transport when glacier
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expansion reduced the transport capacity in adjacent basins. We conclude that effective
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drainage areas varied substantially during the glacial maxima and thus drove changes in
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discharge and sediment flux. While variations in the extent of glaciation alter the annual
210
discharge within each catchment, the magnitude of the variations in discharge produced
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by glacial-drainage capture have a greater effect on long-term discharge and deposition in
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affected basins, and complicate the interpretation of proglacial fluvial stratigraphy as an
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archive of climate change.
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ACKNOWLEDGMENTS
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This work was undertaken while Rowan was in receipt of Natural
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Environment Research Council studentship NE/F008295/1. Additional funding was
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gratefully received from the Dudley Stamp Memorial Fund, the Quaternary Research
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Association New Research Workers Grant, and the British Sedimentological
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Research Group Harwood Fund. The National Institute of Water and Atmospheric
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Research Ltd. (NIWA) provided the gridded rainfall data and Land Information New
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Zealand provided the 50-m digital elevation model. Three anonymous reviewers
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provided constructive comments.
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FIGURE CAPTIONS
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Figure 1. A: Location map (South Island, New Zealand) showing the modern drainage
312
divides (thick black lines; dashes indicate interpretation), mapped ice extents for the Last
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Glacial Maximum (LGM), late LGM and late glacial, LGM ice extent (white shading)
314
(Newnham et al., 1999), and surface trace of the Alpine fault (black dotted line). B: LGM
315
ice cover of the Southern Alps (white shading), and the LGM coastline (dashed line)
316
(Newnham et al., 1999). C: Locations of sites named in the text. D: Coastal cross-section
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of the Canterbury Plains (datum is NZ1949). Vertical exaggeration is × 1000. Inset shows
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Journal: GEOL: Geology
Article ID: G33829
change in sea-surface temperature (SST) (Barrows et al., 2007), optically stimulated
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luminescence ages (circles) (Rowan et al., 2012) for these sediments, and two glacial
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maxima (gray shading).
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Figure 2. Simulated ice volumes overlain on a shaded relief map of the model domain
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showing mapped ice extents (as in Fig. 1). Ice thickness (m; blue shading) and drainage
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divides (red lines) are shown for T of: 4.5 °C (A), 5.0 °C (B), 5.5 °C (C), and 6.5
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°C (D). White stars indicate changes in drainage area. LGM—Last Glacial Maximum.
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Figure 3. Simulated ice-flow velocity–vector maps for the Lake Clearwater region of the
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South Island, New Zealand. A: T = 4.5 °C, Potts Stream Glacier flows into the
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Rangitata, and no drainage capture takes place. B: T = 6.5 °C, part of the Pleistocene
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Rangitata Glacier reverses the lower Potts Stream Glacier and overruns Lake Clearwater,
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flowing into the Ashburton. Ice-flow vectors (black arrows) are scaled for flow velocity.
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Ice thickness (m; blue shading), and drainage divides (red lines) are shown. Location is
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indicated in Figure 1.
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Figure 4. A: Foraminiferal 18O sea-surface temperature (SST) relative to modern values
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from marine core DSDP-597 from 400 km offshore of eastern South Island (Barrows et
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al., 2007). B: Simulated effective drainage area. C: Simulated steady-state annual
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discharge. D: Simulated relative annual bedload sediment flux from model. E: Simulated
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annual discharge resulting from a combination of the steady-state and transient responses.
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Publisher: GSA
Journal: GEOL: Geology
Article ID: G33829
Results for the Rakaia (black dashed line), Ashburton (black solid line), and Rangitata
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(black dotted line). LGM—Last Glacial Maximum.
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1
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available online at www.geosociety.org/pubs/ft2012.htm, or on request from
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editing@geosociety.org or Documents Secretary, GSA, P.O. Box 9140, Boulder, CO
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80301, USA.
GSA Data Repository item 2012xxx, describing the application of the glacier model is
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