Stable oxygen isotope variability in two contrasting glacier river catchments in Greenland

Abstract

Analysis of stable oxygen isotope (¹⁸O) characteristics is a useful tool to investigate water provenance in glacier river systems. In order to attain knowledge on the diversity of ¹⁸O variations in Greenlandic rivers, we examined two contrasting glacierised catchments disconnected from the Greenland Ice Sheet (GrIS). At the Mittivakkat Gletscher river, a small river draining a local temperate glacier in southeast Greenland, diurnal oscillations in ¹⁸O occurred with a 3-h time lag to the diurnal oscillations in run-off. The mean annual ¹⁸O was -14.68-±-0.18-[per thousand] during the peak flow period. A hydrograph separation analysis revealed that the ice melt component constituted 82-±-5-% of the total run-off and dominated the observed variations during peak flow in August 2004. The snowmelt component peaked between 10:00 and 13:00 local time, reflecting the long travel time and an inefficient distributed subglacial drainage network in the upper part of the glacier. At the Kuannersuit Glacier river on the island Qeqertarsuaq in west Greenland, the ¹⁸O characteristics were examined after the major 1995-1998 glacier surge event. The mean annual ¹⁸O was -19.47-±-0.55-[per thousand]. Despite large spatial variations in the ¹⁸O values of glacier ice on the newly formed glacier tongue, there were no diurnal oscillations in the bulk meltwater emanating from the glacier in the post-surge years. This is likely a consequence of a tortuous subglacial drainage system consisting of linked cavities, which formed during the surge event. Overall, a comparison of the ¹⁸O compositions from glacial river water in Greenland shows distinct differences between water draining local glaciers and ice caps (between -23.0 and -13.7-[per thousand]) and the GrIS (between -29.9 and -23.2-[per thousand]). This study demonstrates that water isotope analyses can be used to obtain important information on water sources and the subglacial drainage system structure that is highly desired for understanding glacier hydrology.

Hydrol. Earth Syst. Sci., 20, 11971210, 2016 www.hydrol-earth-syst-sci.net/20/1197/2016/ doi:10.5194/hess-20-1197-2016 Author(s) 2016. CC Attribution 3.0 License.

Jacob C. Yde1, Niels T. Knudsen2, Jrgen P. Steffensen3, Jonathan L. Carrivick4, Bent Hasholt5,Thomas Ingeman-Nielsen6, Christian Kronborg2, Nicolaj K. Larsen2, Sebastian H. Mernild1,7, Hans Oerter8, David H. Roberts9, and Andrew J. Russell10

1Faculty of Engineering and Science, Sogn og Fjordane University College, Sogndal, Norway

2Department of Geoscience, University of Aarhus, Aarhus, Denmark

3Centre for Ice and Climate, University of Copenhagen, Copenhagen, Denmark

4School of Geography and water@leeds, University of Leeds, Leeds, UK

5Department of Geosciences and Natural Resource Management, University of Copenhagen, Copenhagen, Denmark

6Arctic Technology Centre, Technical University of Denmark, Kgs. Lyngby, Denmark

7Antarctic and Sub-Antarctic Program, Universidad de Magallanes, Punta Arenas, Chile

8Alfred Wegener Institute, Helmholtz Centre for Polar and Marine Research, Bremerhaven, Germany

9Department of Geography, University of Durham, Durham, UK

10School of Geography, Politics & Sociology, Newcastle University, Newcastle upon Tyne, UK Correspondence to: Jacob C. Yde ([email protected])

Received: 15 April 2015 Published in Hydrol. Earth Syst. Sci. Discuss.: 18 June 2015 Revised: 18 February 2016 Accepted: 29 February 2016 Published: 21 March 2016

Abstract. Analysis of stable oxygen isotope ( 18O) characteristics is a useful tool to investigate water provenance in glacier river systems. In order to attain knowledge on the diversity of 18O variations in Greenlandic rivers, we examined two contrasting glacierised catchments disconnected from the Greenland Ice Sheet (GrIS). At the Mittivakkat Gletscher river, a small river draining a local temperate glacier in southeast Greenland, diurnal oscillations in 18O occurred with a 3 h time lag to the diurnal oscillations in run-off. The mean annual 18O was 14.68 [notdef] 0.18 during the peak ow pe

riod. A hydrograph separation analysis revealed that the ice melt component constituted 82 [notdef] 5 % of the total run-off and

dominated the observed variations during peak ow in August 2004. The snowmelt component peaked between 10:00 and 13:00 local time, reecting the long travel time and an inefcient distributed subglacial drainage network in the upper part of the glacier. At the Kuannersuit Glacier river on the island Qeqertarsuaq in west Greenland, the 18O characteristics were examined after the major 19951998 glacier surge event. The mean annual 18O was 19.47 [notdef] 0.55 . Despite

large spatial variations in the 18O values of glacier ice on the newly formed glacier tongue, there were no diurnal os-

cillations in the bulk meltwater emanating from the glacier in the post-surge years. This is likely a consequence of a tortuous subglacial drainage system consisting of linked cavities, which formed during the surge event. Overall, a comparison of the 18O compositions from glacial river water in Greenland shows distinct differences between water draining local glaciers and ice caps (between 23.0 and 13.7 ) and

the GrIS (between 29.9 and 23.2 ). This study demon

strates that water isotope analyses can be used to obtain important information on water sources and the subglacial drainage system structure that is highly desired for understanding glacier hydrology.

1 Introduction

There is an urgent need for improving our understanding of the controls on water sources and ow paths in Greenland. As in other parts of the Arctic, glacierised catchments in Greenland are highly sensitive to climate change (Milner et al., 2009; Blaen et al., 2014). In recent decades freshwater run-off from the Greenland Ice Sheet (GrIS) to adjacent

Published by Copernicus Publications on behalf of the European Geosciences Union.

Stable oxygen isotope variability in two contrasting glacier river catchments in Greenland

1198 J. C. Yde et al.: Stable oxygen isotope variability

seas has increased signicantly (Hanna et al., 2005, 2008;

Bamber et al., 2012; Mernild and Liston, 2012), and the total ice mass loss from the GrIS contributes with 0.33 mm sea level equivalent yr1 to global sea level rise (19932010;

Vaughan et al., 2013). In addition, ice mass loss from local glaciers (i.e. glaciers and ice caps peripheral to the GrIS;Weidick and Morris, 1998) has resulted in a global sea level rise of 0.09 mm sea level equivalent yr1 (19932010;

Vaughan et al., 2013). The changes in run-off are coupled to recent warming in Greenland (Hanna et al., 2012, 2013;Mernild et al., 2014), an increasing trend in precipitation and changes in precipitation patterns (Bales et al., 2009; Mernild et al., 2015a), and a decline in albedo (Bggild et al., 2010;Tedesco et al., 2011; Box et al., 2012; Yallop et al., 2012;Mernild et al., 2015b). Also, extreme surface melt events have occurred in recent years (Tedesco et al., 2008, 2011; van As et al., 2012), and in July 2012 more than 97 % of the GrIS experienced surface melting (Nghiem et al., 2012; Keegan et al., 2014). In this climate change context, detailed catchment-scale studies on water source and water ow dynamics are urgently needed to advance our knowledge of the potential consequences of future hydrological changes in Greenlandic river catchments.

Analysis of stable oxygen isotopes is a very useful technique to investigate water provenance in glacial river systems. Stable oxygen isotopes are natural conservative tracers in low-temperature hydrological systems (e.g. Moser and Stichler, 1980; Gat and Gonantini, 1981; Haldorsen et al., 1997; Kendall et al., 2014). Consequently, oxygen isotopes can be applied to determine the timing and origin of changes in water sources and ow paths because different water sources often have isotopically different compositions due to their exposure to different isotopic fractionation processes. Since the 1970s, this technique has been widely used for hydrograph separation (Diner et al., 1970). Most often a conceptual two-component mixing model is applied, where an old-water component (e.g. groundwater) is mixed with a new-water component (e.g. rain or snowmelt), assuming that both components have spatial and temporal homogeneous compositions. The general mixing model is given by the equation

QC = Q1C1 + Q2C2 + ..., (1) where the discharge Q and the isotopic value C are equal to the sum of their components. This simplied model has limitations when a specic precipitation event is analysed because the water isotope composition in precipitation (new water) may vary considerably during a single event (e.g.McDonnell et al., 1990) and changes in contributions from secondary old-water reservoirs may occur (e.g. Hooper and Shoemaker, 1986). Nevertheless, water isotope mixing models still provide valuable information on spatial differences in hydrological processes on diurnal to annual timescales (Kendall et al., 2014).

Figure 1. Location map (a) of the study areas at (b) the Mittivakkat Gletscher river, southeast Greenland (image from Landsat 8 OLI on3 September 2013), and at (c) the Kuannersuit Glacier river, west Greenland (image from Landsat 8 OLI on 8 July 2014).

In glacier-fed river systems, the principal water sources to bulk run-off derive from ice melt, snowmelt, rainfall and groundwater components. Depending on the objectives of the study and on the environmental setting, hydrograph separation of glacial rivers has been based on assumed endmember isotope mixing between two or three prevailing components (Behrens et al., 1971, 1978; Fairchild et al., 1999; Mark and Seltzer, 2003; Theakstone, 2003; Yde and Knudsen, 2004;Mark and McKenzie, 2007; Yde et al., 2008; Bhatia et al., 2011; Kong and Pang, 2012; Ohlanders et al., 2013; Blaen et al., 2014; Dahlke et al., 2014; Hindshaw et al., 2014; Meng et al., 2014; Penna et al., 2014; Rodriguez et al., 2014; Zhou et al., 2014). As glacierised catchments vary in size, altitudinal range, hypsometry, degree of glaciation, and thermal and morphological glacier types, isotope hydrograph separation often requires that the primary local controls on run-off generation are identied in order to analyse the variability in isotope time series. In detailed studies it may even be necessary to divide a main component, such as ice melt, into several ice facies sub-components (Yde and Knudsen, 2004).

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J. C. Yde et al.: Stable oxygen isotope variability 1199

However, in highly glacierised catchments the variability in oxygen isotope composition is generally controlled by seasonal snowmelt and ice melt with episodic inputs of rain-water, whereas contributions from shallow groundwater ow may become important in catchments, where glaciers comprise a small proportion of the total area (e.g. Blaen et al., 2014).

In this study, we examine the stable oxygen isotope composition in two Greenlandic glacier river systems, namely the Mittivakkat Gletscher river (13.6 km2), which drains a local non-surging glacier in southeast Greenland, and the Kuannersuit Glacier river (258 km2), which drains a local glacier on the island Qeqertarsuaq, west Greenland. The latter experienced a major glacier surge event in 19951998. Our aim is to gain insights into the variability and controls of the oxygen isotope composition in contrasting glacierised river catchments located peripheral to the GrIS (i.e. the river systems do not drain meltwater from the GrIS). Besides a study by Andreasen (1984) at the glacier Killersuaq in west Greenland, this is the rst study of oxygen isotope dynamics in rivers draining glacierised catchments peripheral to the GrIS.

2 Study sites

2.1 Mittivakkat Gletscher river, Ammassalik Island, southeast Greenland

Mittivakkat Gletscher (65 41[prime] N, 37 50[prime] W) is the largest glacier complex on Ammassalik Island, southeast Greenland (Fig. 1). The entire glacier covered an area of 26.2 km2 in 2011 (Mernild et al., 2012) and has an altitudinal range between 160 and 880 m a.s.l. (Mernild et al., 2013a). Bulk meltwater from the glacier drains primarily westwards to the proglacial Mittivakkat Valley and ows into the Sermilik Fjord. The sampling site is located at a hydrometric station 1.3 km down-valley from the main subglacial meltwater portal. The hydrological catchment has an area of 13.6 km2,

of which 9.0 km2 is glacierised (66 %). The maritime climate is Low Arctic with annual precipitation ranging from 1400 to 1800 mm water equivalent (w.e.) yr1 (19982006)

and a mean annual air temperature (MAAT) at 515 m a.s.l. of

2.2 C (19932011; updated from Mernild et al., 2008a).

There are no observations of contemporary permafrost in the area, and the proglacial vegetation cover is sparse.

The glacier has undergone continuous recession since the end of the Little Ice Age (Knudsen et al., 2008; Mernild et al., 2011). In recent decades the recession has accelerated and the glacier has lost approximately 29 % of its volume between 1994 and 2012 (Yde et al., 2014), and surface mass balance measurements indicate a mean thinning rate of 1.01 m w.e. yr1 between 19951996 and 20112012 (Mernild et al., 2013a). Similar to other local glaciers in the Ammassalik region, Mittivakkat Gletscher is severely out of contemporary climatic equilibrium (Mernild et al., 2012,

Table 1. Summary of 18O mean and range in bulk water samples at the Mittivakkat Gletscher river.

Year Campaign period n 18Omean 18Omax 18Omin

2003 1113 Aug 4 14.42 14.30 14.65

2004 822 Aug 103 14.55 14.19 14.91

2005 30 May12 Jun 29 14.71 14.35 15.16

2326 Jul 19 14.10 13.74 14.41

1119 Aug 44 14.73 14.13 16.43

2006 1116 Aug 11 14.85 14.26 15.42

2007 210 Aug 17 14.69 14.07 15.11

2008 29 May11 Jun 28 16.92 15.92 17.35

1016 Aug 15 14.84 14.47 15.20

2009 816 Aug 17 14.88 14.56 15.13

Collected at a sampling site ca. 500 m closer to the glacier front.

2013b) and serves as a representative location for studying the impact of climate change on glacierised river catchments in southeast Greenland (e.g. Mernild et al., 2008b, 2015b;Brcena et al., 2010, 2011; Kristiansen et al., 2013; Lutz et al., 2014).

2.2 Kuannersuit Glacier river, Qeqertarsuaq, west Greenland

Kuannersuit Glacier (69 46[prime] N, 53 15[prime] W) is located in central Qeqertarsuaq (formerly Disko Island), west Greenland (Fig. 1). It is an outlet glacier descending from the Sermersuaq ice cap and belongs to the QeqertarsuaqNuussuaq surge cluster (Yde and Knudsen, 2007). In 1995, the glacier started to surge down the Kuannersuit Valley with a frontal velocity up to 70 m per day (Larsen et al., 2010). By the end of 1998 or beginning of 1999, the surging phase terminated and the glacier went into its quiescent phase, which is presumed to last more than 100 years (Yde and Knudsen, 2005a). The 19951998 surge of Kuannersuit Glacier is one of the largest land-terminating surge events ever recorded; the glacier advanced 10.5 km down-valley, and approximately 3 km3 of ice was moved to form a new glacier tongue (Larsen et al., 2010).

The Kuannersuit Glacier river originates from a portal at the western side of the glacier terminus, and the sampling site is located 200 m down-stream (Yde et al., 2005a). The catchment area has an altitude range of 1001650 m a.s.l. and covers 258 km2, of which Kuannersuit Glacier constitutes 103 km2 of the total glacierised area of 168 km2 (Yde and Knudsen, 2005a). The valley oor consists of unvegetated outwash sediment; dead-ice deposits; and ice-cored, vegetated terraces. The proglacial area of the catchment is situated in the continuous permafrost zone (Yde and Knudsen, 2005b), and the climate is polar continental (Humlum, 1999).There are no meteorological observations from the area, but at the coastal town of Qeqertarsuaq (formerly Godhavn), located 50 km to the southwest, the MAAT was 2.7 and 1.7 C in 2011 and 2012, respectively (Cappelen, 2013).

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1200 J. C. Yde et al.: Stable oxygen isotope variability

Table 2. Summary of 18O mean and range in bulk water samples at the Kuannersuit Glacier river.

Year Campaign n 18Omean 18Omax 18Omin period

2000 2427 Jul 21 19.80 19.47 19.97

2001 1431 Jul 109 19.25 17.82 19.55

2002 1415 Jul 21 19.01 18.75 19.39

2003 1826 Jul 27 20.43 19.03 21.88

2005 1924 Jul 2 19.42 19.32 19.51

3 Methods

3.1 Sampling protocol and isotope analyses

In total, 287 oxygen isotope samples were collected from the Mittivakkat Gletscher river during the years 20032009 (Table 1). Most of the sampling campaigns were conducted in August at the end of the peak ow period (i.e. the summer period with relatively high run-off). The most intensively sampled period was from 8 to 22 August 2004, where sampling was conducted with a 4 h frequency supplemented by short periods of higher frequency sampling. In the years 2005 and 2008, meltwater was also collected during the early melt season (i.e. the period before the subglacial drainage system is well established) to evaluate the seasonal variability in the 18O signal. An additional 40 river samples were collected for multi-sampling tests.

During ve eld seasons in July 2000, 2001, 2002, 2003 and 2005, a total of 180 oxygen isotope samples were collected from the Kuannersuit Glacier river (Table 2), and another 44 river samples were collected for multi-sampling tests. In addition, 13 ice samples were obtained along a longitudinal transect at the centreline of the newly formed glacier tongue with 500 m sampling increments in July 2001, and 23 ice samples were collected along a transverse transect with 50 m sampling increments in July 2003. The transverse tran-sect crossed the longitudinal transect at a distance of 3250 m from the glacier front. Seven samples of rainwater were collected in a Hellmann rain gauge located in the vicinity of the glacier terminus in July 2002.

All water samples were collected manually in 20 mL vials.

Ice samples were collected in 250 mL polypropylene bottles or plastic bags before being slowly melted and decanted to 20 mL vials. The vials were stored in cold ( 5 C) and dark

conditions to avoid fractionation related to biological activity.

The relative deviations ( ) of water isotope compositions (18O/16O) were expressed in per mil () relative to Vienna Standard Mean Ocean Water (0 ; Coplen, 1996). The stable oxygen isotope analyses were performed at the Niels Bohr Institute, University of Copenhagen, Denmark, using mass spectrometry with an instrumental precision of [notdef]0.1 in

the oxygen isotope ratio ( 18O) value.

Figure 2. 18O time series of meltwater draining Mittivakkat

Gletscher in (a) 2005 and (b) 2008.

The oxygen isotope data from this study are available in the Supplement (Tables S1S6).

3.2 Multi-sample tests

In the Mittivakkat Gletscher river, we conducted three multi-sample tests at 14:00 local time on 9, 15 and 21 August 2004 to determine the combined uncertainty related to sampling and analytical error. During the multi-sample tests samples were collected simultaneously (within 3 min). The tests show standard deviations of 0.08 (n = 25), 0.06 (n = 5) and

0.04 (n = 10), respectively, which are lower than the in

strumental precision ([notdef]0.1 ).

In the Kuannersuit Glacier river, multi-sample tests were conducted in 2001, 2002 and 2003, showing a standard deviation of [notdef]0.16 (n = 5), [notdef]0.13 (n = 17) and [notdef]0.44

(n = 22), respectively. The multi-sample test in 2003 showed

a standard deviation signicantly larger than the instrumental precision ([notdef]0.1 ). This deviation cannot be explained by

the presence of a few high 18O values. The most plausible explanation is that the glacier run-off was not well mixed in 2003, possibly because different parts of the drainage system merged close to the glacier portal.

3.3 Run-off measurements

Stagedischarge relationships were used to determine runoff at each study site. The accuracy of individual run-off measurements is within [notdef]7 % (e.g. Herschy, 1999). For

details on run-off measurements we refer to Hasholt and

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J. C. Yde et al.: Stable oxygen isotope variability 1201

-14.0

-14.2

-14.4

dO ()

-14.6

-14.8

-15.0

12

Discharge (ms)

10

8

6

4

2

16

Air temperature (C)

0

12

8

4

Figure 4. Time series of 18O (red curve) and discharge (black curve) in the Kuannersuit Glacier river during the period 1431 July 2001.

between the early ablation seasons in 2005 and 2008 is consistent with the meteorological record from Tasiilaq, which shows that the region received a large amount of precipitation in May 2008 (140 mm) compared to a dry May 2005 (17 mm;Cappelen, 2013). Episodic effects on 18O by precipitation seem common throughout the ablation season. For instance, another short-term change occurred on 1415 August 2005 (Fig. 2a), when a negative peak in 18O of 2 coincided

with a snowfall event (14 mm in Tasiilaq; Cappelen, 2013) and subsequent elevated contribution from snowmelt.

During the peak ow periods, the mean annual 18O

was 14.68 [notdef] 0.18 (Table 1). We use the 2004 time se

ries to assess oxygen isotope dynamics in the Mittivakkat Gletscher river during the peak ow period when the sub-glacial drainage system is assumed to be well established, transporting the majority of meltwater in a channelised network (Mernild, 2006). In Fig. 3, the 2004 18O time series

is shown together with run-off (at the hydrometric station), air temperature (at a nunatak at 515 m a.s.l.) and electrical conductivity (at the hydrometric station; corrected to 25 C).

There was no precipitation during the entire sampling period, except for some drizzle on 8 August prior to the collection of the rst sample. The time series shows characteristic diurnal variations in 18O composition, e.g. on 910 and 1618 August 2004. However, the diurnal pattern was severely disturbed at around 03:00 on 11 August 2004. The hydro-graph shows that during the falling limb the diurnal trend in run-off was interrupted, coinciding with an air temperature increase and a change in 18O from decreasing to slightly increasing values. The run-off stayed almost constant until a rapid 39 % increase in run-off occurred at 13:00 on 12 August 2004, accompanied by an increase in 18O and decrease

in electrical conductivity. Thereafter, run-off remained at an elevated level for more than 2 days before returning to a diurnal oscillation of run-off. Hydrograph separation of water sources is a helpful tool to elucidate the details of this event (see Sect. 4.3).

In the Kuannersuit Glacier river, the sample-weighted mean annual 18O was 19.47 [notdef] 0.55 during the peak

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0

20

Conductivity (S cm)

16

12

8

4

0

08.08. 10.08. 12.08. 14.08. 16.08. 18.08. 20.08. 22.08. Date (2004)

Figure 3. Time series of 18O, discharge, air temperature and electric conductivity in meltwater draining Mittivakkat Gletscher during the period 821 August 2004.

Mernild (2006) for the Mittivakkat Gletscher river and Yde et al. (2005a) for the Kuannersuit Glacier river. In short, at the Mittivakkat Gletscher river the run-off measurements were conducted at a hydrometric monitoring station located after the braided river system had changed into a single river channel about 500 m from the river outlet. The station was installed in August 2004 and recorded water stage every 10 min during the peak ow period. At the Kuannersuit Glacier river the run-off measurements were obtained at a hydro-metric monitoring station installed in July 2001 at a location where the river merges to a single channel. Water stage was recorded every hour during the peak ow period. The station was destroyed during the spring river break-up in 2002.

4 Results

4.1 18O characteristics

At the Mittivakkat Gletscher river, the early melt season is characterised by an increasing trend in 18O. In 2005 the

18O values in the early melt season were coincident with the 18O values during the peak ow period (Fig. 2a; Table 1). This indicates that the onset of ice melt commenced before the early melt season sampling campaign. In contrast, the 2008 onset of ice melt was delayed, and snowmelt totally dominated the bulk composition of the river water except on30 May 2008, when a rainfall event (19 mm in the nearby town of Tasiilaq, located 10 km to the southeast of the Mittivakkat Gletscher river catchment; Cappelen, 2013) caused a positive peak in 18O of 1 (Fig. 2b). This difference

1202 J. C. Yde et al.: Stable oxygen isotope variability

Figure 5. Diurnal 18O variations in the Kuannersuit Glacier river on studied days in July in the post-surge years 20002003. Multi-sample tests conducted in 2001, 2002 and 2003 showed standard deviations of [notdef]0.16, [notdef]0.13 and [notdef]0.44 , respectively.

ow period (a sample-weighted value is applied because the number of samples per year deviated between 2 and 109). In Fig. 4, the variations in 18O are presented together with run-off for the period 1431 July 2001. The 2001 run-off measurements showed diurnal oscillations with minimums around 10:0012:00 and maximums at 19:0020:00, correlating well with reversed oscillations in solutes (Yde et al., 2005a) and poorly with suspended sediment concentrations (Knudsen et al., 2007). However, the variability of 18O did

not correlate with run-off or any of these variables. While some of the episodic damming and meltwater release events appear as peaks on the run-off time series, the peaks in the 18O time series coincided with rainfall events (e.g. on the nights of 21 and 29 July 2001). Besides these episodic peaks, a lack of diurnal uctuations in 18O characterised the 18O time series.

Figure 5 shows the diurnal 18O variations during 4 days in July without rainfall in the years 20002003. There were no diurnal oscillations in 2000, 2001 and 2002. In 2003, the uctuations were much larger than in the preceding years, but the highest 18O (19.03 ) was measured at 21:00, and

low 18O prevailed during the night ( 21.0 ). This diur

nal variability was also reected in the standard deviations of the measurements taken over the 24 h periods, which increased from [notdef]0.07 in 2000 to [notdef]0.11, [notdef]0.23 and [notdef]0.70

in 2001, 2002 and 2003, respectively. The corresponding diurnal amplitudes for 20002003 were 0.28, 0.42, 0.64 and2.85 , respectively. Although these measurements from a single day each year are insufcient to represent the conditions for the entire peak ow period, they may indicate post-surge changes in the structure of subglacial hydrological system which are worth addressing in detail in future studies of the hydrological system of surging glaciers.

4.2 18O endmember components

On Mittivakkat Gletscher, three snow pits (0.1 m sampling increments) were excavated at different altitudes

in May 1999, showing a mean 18O composition of

16.5 [notdef] 0.6 (hereafter the uncertainty of 18O is given

by the standard deviation) in winter snow (Dissing, 2000).

The range of individual samples in each snow pit varied between 14.5 and 19.5 (269 m a.s.l.; mean 18O

=16.24 [notdef] 1.35; n = 36), 13.8 and 21.2 (502 m a.s.l.;

mean 18O = 17.11 [notdef] 2.13; n = 21), and 11.9 and 21.6 (675 m a.s.l.; mean 18O = 16.18 [notdef] 2.70; n = 26;

Dissing, 2000). Also, two ice-surface 18O records of 2.84 and 1.05 km in length (10 m sampling increments) were obtained from the glacier terminus towards the equilibrium line (Boye, 1999). The glacier ice 18O ranged between 15.0

and 13.3 with a mean 18O of 14.1 (Boye, 1999),

and the theoretical altitudinal effect (Dansgaard, 1964) of higher 18O towards the equilibrium line altitude (ELA) was not observed. The reasons for an absence of a 18O lapse rate are most likely the limited size and altitudinal range (160 880 m a.s.l.) of Mittivakkat Gletscher, but ice dynamics, ice age and meteorological conditions such as frequent inversion (Mernild and Liston, 2010) may also have an impact.The 18O of summer rain has not been determined in this region, but at the coastal village of Ittoqqortoormiit, located

840 km to the north of Mittivakkat Gletscher, observations show monthly mean 18O in rainwater of 12.8, 9.1 and 8.8 in June, July and August, respectively (data available from the International Atomic Energy Agency database WISER). Based on these observations it is evident that end-member snowmelt has a relatively low 18O compared to endmember ice melt and that these two water source components can be separated. Contributions from rainwater will likely result in episodic increase in the 18O of bulk meltwa

ter.

In the Kuannersuit Glacier river system, the glaciological setting differed from the Mittivakkat Gletscher river system.During the surge event of Kuannersuit Glacier, the glacier front advanced from 500 down to 100 m a.s.l., while a

signicant part of the glacier surface in the accumulation area was lowered by more than 100 m to altitudes below the ELA ( 11001300 m a.s.l.). A helicopter survey in July

2002 revealed that the post-surge accumulation area ratio was less than 20 % (Yde et al., 2005a). Hence, we assume that the primary post-surge water source during the peak ow period is ice melt, particularly from ablation of the new glacier tongue. The mean 18O value of glacier ice collected along the longitudinal and transverse transects was

20.5 [notdef] 1.0 (n = 36). This is consistent with 18O values

of glacier ice located near the glacier front, showing mean 18O of 19.4 [notdef] 0.9 (n = 20) in a section with debris lay

ers formed by thrusting and 19.8 [notdef] 1.1 (n = 37) in a

section without debris layers (Larsen et al., 2010). In contrast to the setting at the Mittivakkat Gletscher river, it was likely that another ice melt component in bulk run-off from Kuannersuit Glacier comprised water from several ice facies sub-component sources with various 18O values and spatial variability. During the surge event, a thick debris-rich basal

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Figure 6. Hydrograph showing the separation of the discharge in the Mittivakkat Gletscher river (black curve) into an ice melt component (red curve) and a snowmelt component (blue curve) during the period 821 August 2004. The error of the ice melt and snowmelt components depends on the constant endmember estimates and the cubic spline interpolation. The arrow indicates the onset of the abrupt change in discharge.

ice sequence was formed beneath the glacier and exposed along the glacier margins and at the glacier terminus (Yde et al., 2005b; Roberts et al., 2009; Larsen et al., 2010). The basal ice consisted of various genetic ice facies, where different isotopic fractionation processes during the basal ice formation resulted in variations in the 18O composition. The 18O in massive stratied ice was 16.6 [notdef] 1.9 (n = 10);

in laminated stratied ice it was 19.6 [notdef] 0.7 (n = 9); and

in dispersed ice it was 18.8 [notdef] 0.6 (n = 41; Larsen et al.,

2010). Also, during the termination of the surge event in winter 19981999, proglacial naled was stacked into 3 m

thick sections of thrust-block naled at the glacier front, as the glacier advanced into the naled (Yde and Knudsen, 2005b; Yde et al., 2005b; Roberts et al., 2009). Naled is an extrusive ice assemblage formed in front of the glacier by rapid freezing of winter run-off and/or proglacial upwelling water mixed with snow. A prole in a thrust-block naled section showed a 18O of 20.1 [notdef] 0.5 (n = 60; excluding an out

lier polluted by rainwater; Yde and Knudsen, 2005b). With regard to the endmember compositions of snowmelt and rain-water at the Kuannersuit Glacier river, it was not possible to access snow on the upper part of the glacier, so no 18O val

ues on snowmelt were measured. Rainwater was collected during rainfall events in July 2002, showing a wide range in 18O between 18.78 and 6.57 and a median 18O of 10.32 [notdef] 4.49 (n = 7; Table S6).

4.3 Hydrograph separation

The conditions for conducting hydrograph separation during the peak ow period were different for the two study catchments. At the Mittivakkat Gletscher river it was possible to distinguish between the 18O values of endmember ice melt and snowmelt components, and there were diurnal oscillations in 18O. In contrast, the available data from the Kuan-

nersuit Glacier river did not allow hydrograph separation in the years following the surge event. Here, there were no diurnal oscillations in 18O, and the composition and importance of the snowmelt component were unknown. Hence, we will continue by using the 2004 time series to construct a two-component hydrograph separation (Eq. 1) during a period without precipitation for the Mittivakkat Gletscher river.

First, we apply time-series cubic spline interpolation to estimate 18O at 1 h time-step increments, matching the temporal resolution of the run-off observations. This approach allows a better assessment of the diurnal 18O signal. For instance, a best-t analysis shows that overall the 18O signal

lags 3 h behind run-off (r2 = 0.66; linear correlation without

lag shows r2 = 0.58), indicating the combined effect of the

two primary components, snowmelt and ice melt, on the 18O

variations. The diurnal amplitude in 18O ranged between0.11 (11 August 2004) and 0.49 (16 August 2004). However, there was no statistical relation between diurnal 18O

amplitude and daily air temperature amplitude (r2 = 0.28),

indicating that other forcings than variability in surface melting may have a more dominant effect on the responding variability in 18O.

Based on the assumption that snowmelt and ice melt reect their endmember 18O compositions (16.5 and 14.1 , respectively), a hydrograph showing contributions from snowmelt and ice melt is constructed for the 2004 sampling period (Fig. 6). The ice melt component constituted 82 [notdef] 5 % (where [notdef] indicates the standard deviation of the

hourly estimates) of the total run-off and dominated the ob-served variations in total run-off (r2 = 0.99). This is expected

late in the peak ow period, when the subglacial drainage mainly occurs in a channelised network in the lower part of the glacier (Mernild, 2006). The slightly decreasing trend in the daily snowmelt component was likely a consequence of the diminishing snow cover on the upper part of the glacier.The snowmelt component peaked around 10:0013:00 each day, reecting the long distance from the melting snowpack to the proglacial sampling site and the possible existence of an inefcient distributed subglacial drainage network in the upper part of the glacier.

The most likely reason for an abrupt change in glacial runoff, such as the one observed during the early morning of 11 August 2004 followed by the sudden release of water 34 h later, is a roof collapse causing ice-block damming of a major subglacial channel. The hydrograph separation (Fig. 6) shows that the proportion between ice melt and snowmelt remained almost constant after the event commenced, indicating that the bulk water derived from a well-mixed part of the drainage system, which was unaffected by the large diurnal variation in ice melt generation. This suggests that the functioning drainage network transported meltwater from the upper part of the glacier with limited connection to the drainage network in the lower part. Meanwhile, ice melt was stored in a dammed section of the subglacial network located in the lower part of the glacier and suddenly released when the

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1204 J. C. Yde et al.: Stable oxygen isotope variability

dam broke at 13:00 on 12 August (Fig. 6). In the following hours ice melt comprised up to 94 % of the total run-off. On13 August the snowmelt component peaked at noon but then dropped markedly, and in the evening it only constituted 4 % of the total run-off. On 14 August there were still some minor disturbances in the lower drainage network, but from 15 August the drainage system had stabilised and the characteristic diurnal glacionival oscillations had taken over (Figs. 3 and 6).

4.4 Uncertainties in 18O hydrograph separation models

The accuracy of endmember hydrograph separation models is limited by the uncertainties of the estimated values of each endmember component, the uncertainty of the cubic spline interpolation at each data point and the uncertainty of 18O

in the river. While the uncertainty of 18O in the river is likely to be relatively small, the uncertainties of each endmember component must be kept in mind (e.g. Cable et al., 2011;Arendt et al., 2015). The assumption of discrete values of each endmember component is unlikely to reect the spatial and temporal changes in bulk 18O of snowmelt, ice melt and rainwater. For instance, Raben and Theakstone (1998) found a seasonal increase in mean 18O in snow pits on Austre Okstindbreen, Norway, and episodic events such as passages of storms (e.g. McDonnell et al., 1990; Theakstone, 2008) or melting of fresh snow in the late ablation season may cause temporal changes in one component. Also, snow-packs have a non-uniform layered structure with heterogeneous 18O composition, and isotopic fractionation is likely to occur as melting progresses and the snowpack is mixed with rainwater (e.g. Raben and Theakstone, 1998; Lee et al., 2010). It is also difcult to assess how representative snow pits and ice transects are for the bulk 18O value of each component. Spatial differences in 18O may exist within and between snow pits, but the overall effect on the isotopic composition of the water leaving the melting snowpack at a given time is unknown.

4.5 Longitudinal and transverse 18O transects

Glacier ice samples were collected on the surface of Kuannersuit Glacier to gain insights into the spatial variability of 18O on the newly formed glacier tongue. Both the longitudinal and transverse transects showed large spatial uctuations in 18O (Fig. 7). The longitudinal transect was sampled along the centreline but showed unsystematic uctuations on a 500 m sampling increment scale. In contrast, the transverse transect, which was sampled 3250 m up-glacier with 50 m increments, showed a more systematic trend where relatively high 18O values were observed along both lateral margins. From the centre towards the western margin an increasing trend of 0.46 per 100 m prevailed, whereas the eastern central part showed large uctuations in 18O be-

Figure 7. Variations in 18O of glacier ice along a longitudinal tran-sect and a transverse transect on Kuannersuit Glacier. The trans-verse transect crosses the longitudinal transect at a distance of 3250 m from the glacier terminus.

tween 22.69 and 20.08 . The total range of measured

18O in glacier ice along the transverse transect was 4.14 .

A possible explanation of this marked spatial variability may be that the ice forming the new tongue derived from different pre-surge reservoirs on the upper part of the glacier. If so, it is very likely that the marginal glacier ice was formed at relatively low elevations (high 18O signal), whereas the glacier ice in the western central part mainly derived from high-elevation areas of Sermersuaq ice cap (low 18O sig

nal). At present, there are only a few comparable studies on transverse variations in 18O across glacier tongues. Epstein and Sharp (1959) found a decrease in 18O towards the margins of Saskatchewan Glacier, Canada. Hambrey (1974) measured a similar decrease in 18O towards the margins of Charles Rabots Bre, Norway, in an upper transect, whereas a lower transect showed wide unsystematic variations in 18O.

Hambrey (1974) concluded that in the upper transect the marginal ice derived from higher altitudes than ice in the centre, whereas in the lower transect the wide variations were related to structural complexity of the glacier. However, both

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Table 3. Maximum and minimum 18O in glacier rivers.

Site Sampling period Latitude Longitude Maximum () Minimum () Reference

GreenlandMittivakkat Gletscher (local glacier) 20032009 65 41[prime] N 37 50[prime] W 13.7 17.4 This paper

Kuannersuit Glacier (ice cap outlet) 20002005 69 46[prime] N 53 15[prime] W 17.8 21.9 This paper

Hobbs Gletscher (local glacier) 2004 65 46[prime] N 38 11[prime] W 14.7 15.1 Yde (unpublished data)

Imersuaq (GrIS outlet) 2000 66 07[prime] N 49 54[prime] W 24.3 29.9 Yde and Knudsen (2004)

Killersuaq (ice cap outlet) 19821983 66 07[prime] N 50 10[prime] W 19.5 23.0 Andreasen (1984)

Leverett Glacier (GrIS outlet) 2009 67 04[prime] N 50 10[prime] W 23.2 24.2 Hindshaw et al. (2014)

Isunnguata Sermia (GrIS outlet) 2008 67 11[prime] N 50 20[prime] W 26.2

a Yde (unpublished data) N Glacier (GrIS outlet) 2008 68 03[prime] N 50 16[prime] W 23.3 28.3 Bhatia et al. (2011)

Scandinavia and SvalbardAustre Okstindbreen, Norway 19801995 66 00[prime] N 14 10[prime] E 11.8 14.4 Theakstone (2003)

Storglaciren, Sweden 2004 & 2011 67 54[prime] N 18 38[prime] E 10.9 15.9 Dahlke et al. (2014)

Austre Grnfjordbreen, Svalbard 2009 77 56[prime] N 14 19[prime] E 11.2

a Yde et al. (2012)

Dryadbreen, Svalbard 2012 78 09[prime] N 15 27[prime] E 13.0 15.5 Hindshaw et al. (2016)

Longyearbreen, Svalbard 2004 78 11[prime] N 15 30[prime] E 12.3 16.7 Yde et al. (2008) European AlpsGlacier de Tsaneuron, Switzerland 1994 46 20[prime] N 07 15[prime] E 7.8 12.2 Fairchild et al. (1999)

Dammagletscher, Switzerland 2008 46 38[prime] N 08 27[prime] E 13.3 17.3 Hindshaw et al. (2011)

Hintereisferner, Austria 19691970 46 49[prime] N 10 48[prime] E 13.8 19.4 Behrens et al. (1971)

Kesselwandferner, Austria 19691970 46 50[prime] N 10 48[prime] E 14.8 18.1 Behrens et al. (1971) AndesCordillera Blanca catchments, Peru 20042006 910 S 7778 W 13.3 15.3 Mark and McKenzie (2007)

Juncal River, Chile 20112012 32 52[prime] S 70 10[prime] W 16.4 18.0 Ohlanders et al. (2013) AsiaHailuogou Glacier river, China 20082009 29 34[prime] N 101 59[prime] E 13.7 17.6 Meng et al. (2014)

Kumalak Glacier no. 72, China 2009 41 49[prime] N 79 51[prime] E 9.8

a Kong and Pang (2012)

Urumqi Glacier no. 1, China 2009 43 07[prime] N 86 48[prime] E 8.7

a Kong and Pang (2012)

a Single sample.

of these studies are based on few samples. Hence, it therefore remains unknown whether a high spatial variability in 18O is

a common phenomenon or related to specic circumstances such as surge activity or presence of tributary glaciers.

5 Discussion

5.1 Differences in 18O between the Mittivakkat Gletscher river and Kuannersuit Glacier river

A signicant difference between the 18O dynamics in the Mittivakkat Gletscher river and Kuannersuit Glacier river is the marked diurnal oscillations in the former and the lack of a diurnal signal in the latter during the peak ow period. At the Mittivakkat Gletscher river, the 2004 hydrograph separation analysis showed a 3 h lag of 18O to run-off caused by the difference in travel time for ice melt and snowmelt. Meltwater in the early melt season was dominated by snowmelt with relatively high 18O and weak diurnal oscillations; whereas diurnal oscillations with amplitudes between 0.11 and 0.49 existed during the peak ow period due to mixing of a dominant ice melt component and a secondary snowmelt component. Diurnal oscillations in 18O are common in meltwater from small, glacierised catchments; for instance, at Aus-

tre Okstindbreen, Norway, the average diurnal amplitude is approximately 0.2 (Theakstone, 1988, 2003; Theakstone and Knudsen, 1989, 1996a, b). The largest diurnal amplitudes in 18O (up to 4.3 ) have been observed in small-scale GrIS catchments, such as at Imersuaq and N Glacier, where large differences in 18O exist between various ice facies and snowmelt (Yde and Knudsen, 2004; Bhatia et al., 2011).

The lack of strong diurnal oscillations as observed in the post-surge years at the Kuannersuit Glacier river indicates a mono-source system, a well-mixed drainage network or a multi-source system, where the primary components have similar 18O compositions. The expected primary component, glacier ice melt, has lower 18O than bulk run-off, and there must be additional contributions from basal ice melt (similar 18O composition to run-off), snowmelt (unknown 18O composition) or rainwater (higher 18O composition than run-off). We therefore hypothesise that the presence of a well-mixed drainage network is the most likely reason for the observed 18O signal in the bulk run-off from Kuannersuit Glacier. During the surge event the glacier surface became heavily crevassed and the pre-existing drainage system collapsed (Yde and Knudsen, 2005a). It is a generally accepted theory that the drainage system of surging glaciers trans-

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1206 J. C. Yde et al.: Stable oxygen isotope variability

forms into a distributed network where meltwater is routed via a system of linked cavities (Kamb et al., 1985; Kamb, 1987), but little is known about how subglacial drainage systems evolve into discrete ow systems in the years following a surge event. In the initial quiescent phase at Kuannersuit Glacier, frequent loud noises interpreted as drainage system roof collapses were observed, in addition to episodic export of ice blocks from the portal, suggesting ongoing changes to the englacial and subglacial drainage system. A consequence of these processes is also visible on the glacier surface, where circular collapse chasms formed above marginal parts of the subglacial drainage system (Yde and Knudsen, 2005a).

Lack of diurnal oscillations in 18O has previously been related to other causes at non-surging glaciers. At Glacier de Tsaneuron, Switzerland, sampling in the late melt season (2327 August 1994) showed no diurnal variations in 18O, which was interpreted by Fairchild et al. (1999) as a consequence of limited altitudinal range (less than 500 m) of the glacier. An alternative explanation may be that snowmelt only constituted so small a proportion of the total run-off in the late melt season that discrimination between snowmelt and ice melt was impossible. At the glacier Killersuaq, an outlet glacier from the ice cap Amitsulooq in west Greenland, Andreasen (1984) found that diurnal oscillations in 18O were prominent during the relatively warm summer of 1982, whereas no diurnal 18O oscillations were observed in 1983 because the glacier was entirely snow-covered throughout the ablation season, due to low summer surface mass balance caused by the 1982 El Chichn eruption (Ahlstrm et al., 2007).

5.2 18O compositions in glacier rivers

It is clear from the studies at Mittivakkat Gletscher and Kuannersuit Glacier that glacier rivers have different 18O compo

sitions. The bulk meltwater from Mittivakkat Gletscher has a 18O composition similar to the water draining the nearby local glacier Hobbs Gletscher and to waters from studied valley and outlet glaciers in Scandinavia, Svalbard, the European Alps, the Andes and Asia (Table 3). The 18O composition of Kuannersuit Glacier is lower and similar to the 18O com

position of the glacier Killersuaq (Table 3). Currently, the lowest 18O compositions are found in bulk meltwater draining the GrIS in west Greenland (Table 3), but there is a lack of 18O data from Antarctic rivers. Estimations of 18O based on D measurements suggest 18O values of 32.1, 34.4

and 41.9 in waters draining Wilson Piedmont Glacier,

Rhone Glacier and Taylor Glacier, respectively (Henry et al., 1977).

The differences in 18O in glacial rivers are due to a combination of geographical effects related to altitude, continentality and latitude (Dansgaard et al., 1973) and temporal effects that work on various timescales and in specic environments. These temporal effects include a seasonal effect (Dansgaard, 1964), a monsoonal effect (Tian et al., 2001;

Kang et al., 2002), a precipitation amount effect (Holdsworth et al., 1991) and a palaeoclimatic effect (Reeh et al., 2002).For instance, the altitude and continentality effects cause low 18O in rivers draining the GrIS compared to rivers draining valley glaciers at similar latitudes (Table 3). More data on the 18O composition and dynamics in glacial rivers are needed to improve the understanding of how the relative inuence of geographical and temporal effects varies on local and regional scales.

6 Conclusions

In this study, we have examined the oxygen isotope hydrology in two of the most studied glacierised river catchments in Greenland to improve our understanding of the prevailing differences between contrasting glacial environments. This study has provided insights into the variability and composition of 18O in river water draining glaciers and ice caps adjacent to the GrIS.

The following results were found:

The Mittivakkat Gletscher river on Ammassalik Island, southeast Greenland, has a mean annual 18O of

14.68 [notdef] 0.18 during the peak ow period, which

is similar to the 18O composition in glacier rivers in Scandinavia, Svalbard, the European Alps, the Andes and Asia. The Kuannersuit Glacier river on the island Qeqertarsuaq, west Greenland, has a lower mean annual 18O of 19.47 [notdef] 0.55 , which is similar to the 18O

composition in bulk meltwater draining an outlet glacier from the ice cap Amitsulooq but higher than the 18O

composition in bulk meltwater draining the GrIS.

In the Mittivakkat Gletscher river the diurnal oscillations in 18O were conspicuous. This was due to the presence of an efcient subglacial drainage system and diurnal variations in the ablation rates of snow and ice that had distinguishable oxygen isotope compositions.The diurnal oscillations in 18O lagged behind the diurnal oscillations in run-off by approximately 3 h. A hydrograph separation analysis revealed that the ice melt component constituted 82 [notdef] 5 % of the total run-off and

dominated the observed variations in total run-off during the peak ow period in 2004. The snowmelt component peaked between 10:00 and 13:00, reecting the long travel time and a possibly inefciently distributed subglacial drainage network in the upper part of the glacier.

In contrast to the Mittivakkat Gletscher river, the Kuannersuit Glacier river showed no diurnal oscillations in 18O. This is likely a consequence of glacier surging. In the years following a major surge event, where Kuannersuit Glacier advanced 10.5 km, meltwater was routed through a tortuous subglacial conduit network of linked

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J. C. Yde et al.: Stable oxygen isotope variability 1207

cavities, mixing the contributions from glacier ice, basal ice, snow and rainwater.

This study has shown that environmental and physical contrasts in glacier river catchments inuence the spatio-temporal variability of the 18O compositions. In Greenlandic glacier rivers, the variability in 18O com

position is much higher than previously known ranging from relatively high 18O values in small-scale coastal glacierised catchments to relatively low 18O values in GrIS catchments. This study demonstrates that water isotope analyses can be used to obtain important information on water sources and subglacial drainage system structure that is highly desired for understanding glacier hydrology.

The Supplement related to this article is available online at http://dx.doi.org/10.5194/hess-20-1197-2016-supplement

Web End =doi:10.5194/hess-20-1197-2016-supplement .

Acknowledgements. We thank all the students who have participated in the eldwork over the years. We are also grateful to the University of Copenhagen for allowing us to use the facilities at the Arctic Station and Sermilik Station and to the Niels Bohr Institute, University of Copenhagen, for processing the isotope samples. We thank Andreas Peter Bech Mikkelsen and four reviewers for valuable comments on the manuscript.

Edited by: M. Hrachowitz

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