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Impact of permafrost thaw on the turbidity regime of a

subarctic river: the Sheldrake River, Nunavik, Quebec

Journal: Arctic Science

Manuscript ID AS-2016-0006.R3

Manuscript Type: Article

Date Submitted by the Author: 26-Apr-2017

Complete List of Authors: Jolivel, Maxime; centre d'études nordiques, géographie Allard, Michel; Université Laval, Centre d'études nordiques

Keyword: permafrost, Northern Quebec, thermokarst, turbidity, subarctic river

https://mc06.manuscriptcentral.com/asopen-pubs

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Impact of permafrost thaw on the turbidity regime of a subarctic river: the 1

Sheldrake River, Nunavik, Quebec. 2

Maxime Jolivel and Michel Allard 3

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M. Jolivel and M. Allard, Centre d’études nordiques (CEN) and Département de 5

Géographie, Université Laval, Québec QC, G1V 0A6 Canada. 6

Corresponding author: [email protected] 7

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Abstract 22

In order to assess the impact of seasonal active layer thaw and thermokarst on 23

river flow and turbidity, a gauging station was installed near the mouth of the Sheldrake 24

River in the discontinuous permafrost zone of Northern Quebec. The station provided 25

five years of water level data and three years of turbidity data. The hydrological data for 26

the river showed the usual high water stage occurring at spring snow melt, with smaller 27

peaks related to rain events in summer. Larger and longer turbidity peaks also occurred in 28

summer in response to warm air temperature spells suggesting that a large part of the 29

annual suspension load was carried during mid-summer turbidity peaks. Supported by 30

geomorphological observations across the catchment area, the most plausible 31

interpretation is that the rapid thawing of the active layer during warm conditions in July 32

led to the activation of frostboils and triggered landslides throughout the river catchment, 33

thus increasing soil erosion and raising sediment delivery into the hydrological network. 34

These results indicate that maximum sediment discharge in a thermokarst-affected region 35

may be predominantly driven by the rate of summer thawing and associated activation of 36

erosion features in the catchment. 37

Keywords: permafrost, Northern Quebec, thermokarst, turbidity, subarctic river 38

Introduction 39

Rivers are natural pathways from land to sea that carry sediments and other matter 40

eroded from their catchments. Their behavior reflects different geomorphic and 41

biogeochemical processes in the landscape with cascading effects downstream to the 42

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coastal marine environment. In the context of climate change, river studies are essential 43

to quantify the pace and intensity of geosystem changes (Prowse et al. 2015). In subarctic 44

regions, the riverine hydrological regime is strongly linked to seasonal climate variations, 45

which generate a large annual range in water discharge (Déry et al. 2005). Sediment loads 46

can also be extremely variable even if they are sometimes low in comparison with rivers 47

from temperate and tropical regions (Syvitski 2002). 48

The hydrological cycle of high latitude rivers is regulated by snow storage and 49

melting and by the freezing of soil water. Permafrost is a major factor that restricts 50

infiltration and percolation at depth; a perched water table is maintained in the active 51

layer near the surface in summer (Carey and Woo 2001; Carey and Quinton 2005; 52

Quinton and Carey 2008). Base flow may cease in winter since sub-permafrost 53

groundwater may be non-existent or too deep to discharge in the catchment and because 54

taliks can be only poorly connected with springs on the river beds. Soil warming, 55

thinning and decay of permafrost, earlier breakups, decline of snow cover duration and 56

increase in shrub, forest and peatland covers are factors affecting the hydrology of high 57

latitude rivers under ongoing climate change (Magnuson et al. 2000; Sturm et al. 2001; 58

Payette et al. 2004; Brown and Romanovsky 2008; Jolivel and Allard 2013; Lesack et al. 59

2014). For example, it is broadly expected that the sediment load of high latitude rivers 60

would increase by 30% for every 2 °C of warming of the averaged catchment temperature 61

(Syvitski 2002). 62

Thawing of permafrost is known to release large volumes of sediments through 63

thermokarst processes such as thaw slumping and thermal erosion (Jolivel and Allard 64

2013; Kokelj et al. 2013). The released sediments are mobilized by soil erosion, in 65

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overland flow and in water courses; they feed sedimentation (Guo et al. 2004; Goni et al. 66

2005; Jolivel et al. 2015) and get involved in biogeochemical processes (Emmerton et al. 67

2008; Galand et al. 2008; Vonk et al. 2015) in lakes, deltas and coastal seas. 68

Evidence of general permafrost decay has been observed throughout all high-69

latitude regions (e.g. Sollid and Sorbel 1998; Luoto and Seppälä 2003; Jorgenson et al. 70

2006). In northern Québec, near the southern limit of permafrost distribution, thawing of 71

permafrost in large areas of palsas, lithalsas, peat plateaus and permafrost plateaus has 72

led to the reduction of permafrost extent by roughly 40% over the last 50 years (Payette 73

et al. 2004; Marchildon, 2007; Vallée and Payette 2007; Fortier and Aubé-Maurice 2008; 74

Jolivel and Allard, 2013). Continued warming will lead to further degradation, releasing 75

sediments and making previously frozen organic matter available for bacterial 76

decomposition and recycling into bio-available carbon and greenhouse gases (Schuur et 77

al. 2008; Deshpande et al. 2015; Vonk et al. 2015). 78

Thermokarst and associated landslides generate large sediment loads in rivers. 79

This is particularly evident in the case of retrogressive thaw slumps and large active layer 80

detachment slides (Kokelj et al. 2002; Lewis et al. 2005; Lewkowicz and Harris 2005a; 81

Jorgenson et al. 2006; Lantuit and Pollard 2008; Lantz and Kokelj 2008; Lamoureux and 82

Lafrenière 2009; Lacelle et al. 2010; Kokelj et al. 2013). These inputs can alter terrestrial 83

and aquatic ecosystems and affect food webs as well as primary and secondary 84

production (Kokelj et al. 2002, 2009; Bowden et al. 2008; Mesquita et al. 2010). 85

Ultimately, a significant fraction of the organic carbon released by thermokarst may 86

reach the marine environment (Jolivel et al. 2015; Vonk et al. 2015). 87

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More gauging of rivers and monitoring of processes are necessary to understand 88

the sedimentary regime of rivers in regions affected by thermokarst. This is particularly 89

true during periods of significant temporal change in fluvial fluxes resulting from 90

seasonal variations of thermokarst processes in response to climate forcing (Prowse et al. 91

2015). There are few measurements of the impacts of eroding permafrost catchments on 92

fluvial sedimentary regimes (Bowden et al. 2008), and more data are required to better 93

understand geomorphological processes in these regions in transition. 94

The main objectives of this study were to (1) document the annual and seasonal 95

hydrologic fluctuations of a Subarctic river; (2) describe the dynamics of turbidity and 96

sediment fluxes during the thawing season; and (3) assess the relative impacts of 97

precipitation and thawing on discharge and sediment transport. Because the rate of soil 98

thawing influences the rate of thermokarst which releases sediments in the drainage 99

network, we raised the hypothesis that variations in air temperature can influence 100

turbidity of surface water, and so the amount and timing of sediment fluxes in the 101

collector river. 102

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Methods 104

Study area 105

The 25 km long Sheldrake River flows to the eastern coast of Hudson Bay. It 106

drains a 76 km2 watershed (Fig. 1). Its catchment is typical of the area of decaying 107

sporadic/discontinuous permafrost in the Tyrrell sea fine sediments of Eastern Hudson 108

Bay. The Sheldrake is among many rivers of the east Hudson Bay watershed, including 109

large fluvial systems such as the Nastapoka River that transport sediment resulting from 110

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erosional, thermokarst and periglacial processes (Fig. 2).Typically, in summer conditions, 111

the river varies in width from 25 to 50 m. Its depth varies between 3 m in water pools to 112

50 cm in rapids. In the upper part of the river, the river bed is generally composed of a 113

thin veneer of sand and gravel covering thick marine silty clay. From its passage in the 114

coastal hills to the shore, the bed and the banks are essentially composed of exposed 115

bedrock and boulders, which greatly limits bed load transport. 116

The river originates from Sheldrake Lake, on the Archean sector of the Canadian 117

Shield. Near the coast, the river valley runs across a range of coastal hills in Late 118

Proterozoic bedrock and the river flows into Hudson Bay at 56°37’N; 76°32’W (Fig. 1). 119

On a topographic 1:50,000 map, the low-gradient Sheldrake River is a third order stream. 120

However, the rapid and recent permafrost decay increased the hydrologic connectivity 121

between thermokarst ponds, hollows and gullies thereby increasing stream density 122

(Jolivel and Allard 2013). 123

The east-west elongated shape of the catchment (Gravelius index: 1.9; Gravelius 124

1914) is principally due to the carving activity of the Pleistocene glaciers that flowed to 125

the west. Inland, the topography is dominated by flat valley floors, scattered with lakes 126

and small hills with a general elevation range of 200 to 250 m a.s.l. The mean 127

longitudinal river slope is 0.6%. It is <0.5% inland and increases to 3% once crossing the 128

coastal hills near the Hudson Bay. 129

The entire watershed was submerged after deglaciation by the postglacial Tyrrell 130

Sea from 8000 BP to about 6000 BP. Therefore, 85% of the surficial deposits of the area 131

are marine silty clays (Fig. 3). Inland, sand and gravel deposits are associated with ice-132

contact glacio-fluvial deltas which mark the postglacial marine limit, whereas glacio-133

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marine sand and gravel anchored on the slopes of the coastal hills have been reworked 134

into beaches and terraces during the postglacial fall of the relative sea level (Fig. 3) 135

(Lavoie et al. 2012; Lajeunesse and Allard 2003). Because of permafrost, dispersed forest 136

and tundra cover, presence of impermeable clay and bedrock, water infiltration and 137

percolation rates are very low, therefore, most of the water input from rainfalls and 138

snowmelt flows as surface run-off. However, numerous wetlands, topographic 139

depressions, thermokarst ponds and lakes in the Sheldrake River catchment act as surface 140

water storage areas. 141

The regional climate is subarctic, characterized by cold winters (monthly average 142

air temperature of -24°C in January), cool summers (monthly average air temperature of 143

10°C in August) and mean annual temperatures varying between -4°C and -5°C. The area 144

is covered by snow for ~ 8 months every year. Rain accounts for 60% of total 145

precipitation and snow for 40%. The average annual precipitation is 550 mm 146

(Environment Canada 2013) while the average annual evapotranspiration is ~ 200 mm 147

yr-1 (Payette and Rochefort 2001). The eastern sector of Hudson Bay is generally ice-148

covered from early December to the end of May or beginning of June. However, during 149

the warm winter of 2010-2011, freeze-up did not occur before mid-January. The 150

Sheldrake River has a ~1 m thick ice cover from early November onwards. The breakup 151

occurs a few weeks before the melt of the Hudson Bay ice cover. Snowmelt generally 152

occurs in late May and early June but some thicker snow banks can last until mid-153

summer in the landscape. Between 2009 and 2014, the average date of the beginning of 154

the Sheldrake River ice breakup was 6 May (σ= 6 days) and the level of the river 155

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remained high during nearly two months before reaching its summer flow regime (Table 156

1). 157

The basin is located at the transition between the sporadic permafrost zone and 158

the widespread discontinuous permafrost zone (Allard and Seguin 1987). Permafrost is 159

present in 20% of the surficial deposits in the form of lithalsas, palsas, permafrost 160

plateaus and peat plateaus (Fig. 1) (Jolivel and Allard 2013). Lithalsas are mineral 161

permafrost mounds, while palsas are peaty permafrost mounds or mineral-cored 162

permafrost mounds with a peat cover. Permafrost plateaus and peat plateaus are 163

equivalent but wider landforms. These heaved landforms are generally 3 to 5 m high. 164

This is due to the development of ice segregation lenses formed by cryosuction in the 165

frost sensitive marine silt (Pissart 1985, 2002). Permafrost thickness in these frozen 166

landforms varies typically from 10 to 15 m (Lévesque et al. 1988). 167

The Sheldrake River catchment is actually an area of intensive permafrost decay. 168

Between 1957 and 2009, 21% of the area covered by permafrost disappeared. 169

Widespread thermokarst ponds, landslides, active layer failures, and expanding gullies 170

are the main features of permafrost degradation (Jolivel and Allard 2013). Thermokarst 171

pond coverage has nearly doubled between 1957 and 2009 allowing an increase of the 172

stream density and better connections between water tracks. The number of active 173

erosional landforms counted in the landscape has increased by 46 to 217%, 174

corresponding to an increase in the volume of eroded fine-grained sediments of 12 to 175

38% potentially released in the fluvial system (Jolivel and Allard 2013). This degradation 176

is more important inland to the extent that some sub-catchment areas are now devoid of 177

permafrost. 178

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The tree line runs across the basin and is oriented from south to north due to the 179

cooling influence of the Hudson Bay; the landscape close to the bay is dominated by the 180

shrub tundra while the forest tundra extends further inland (Fig. 3). In the western part of 181

the basin (shrub tundra area), the river and its tributaries flow through an area of 182

permafrost plateaus, with incising meanders and deep tributary gullies (Fig. 1). In the 183

forest tundra in the eastern part, some peat deposits over clay or sand extend over flat and 184

poorly drained valley bottoms. Most of the palsas in the catchment are located in those 185

bogs and fens. 186

Field instrumentation and laboratory analyses 187

A gauging station was installed 2 km upstream from the river mouth in late summer 2008 188

(see Fig. 3 for location). As the river bed is rocky and the current is strong in the lower 189

reach of the river, the instrumentation had to be installed between two rapids in a 190

convenient pool of calmer water. This automated station continually records water 191

temperature, water level and turbidity. The mooring consists of a dead weight (20 kg) 192

attached to a buoy that rests on the river bed and is tied to the shoreline with a steel cable 193

(Fig. 4). The buoy was submerged under ~ 1.2 m water at a distance of 4 m from the river 194

bank in order to avoid being swept away during ice breakup. The instrumentation 195

attached to the mooring, 20 cm from the river bottom, consists of a Levellogger (Solinst) 196

and an OBS 3+ (Optical Backscatter Sensor, Campbell Scientific, Inc.). As the depth 197

sensor is close to the river bed, possible changes in the verticality of the mooring cable at 198

higher flow speed can introduce only a very slight error in stage measurement. Given the 199

swift current (maximum measured velocity in summer: 0.5 m/s) that favors mixing in the 200

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pool, we consider that the measured turbidity at this point represents the average turbidity 201

of the whole river cross-sectional area, despite the possibility of a gradient with depth or 202

along cross-bank distance. Finally, the bedrock riverbed prevents change in the stream 203

cross section and bed scour that could locally affect turbidity values at the measurement 204

site during high stages. 205

The levelogger was calibrated for barometric compensation with a barologger 206

(Solinst). The accuracy of the water level sensor is ±0.3 cm. The turbidity sensor has a 207

range of 0–1000 nephelometric turbidity units (NTU) with an accuracy reading of 2% or 208

0.5 NTU (whichever is larger). Both the levelloger and the turbidimeter have a 1-hour 209

time step output which was then converted to daily average data to facilitate the reading 210

of the graphs. No noise effect was recorded during ice-free conditions. The mooring is 211

situated several meters upstream of a 20 m waterfall. Thus, at spring, breakup at the 212

gauging site is facilitated and backwater effects are prevented even if the river upstream 213

breaks up several days later. Data presented in this paper reflect hydrological conditions 214

at the gauging station only. 215

An attempt of conversion of water level into discharge was done but the 216

difficulties to measure current velocity during extreme high flow due to remote access 217

prevented the calculation of a robust relationship. Thus, in this study, we have settled for 218

water level data only. However, stage and discharge are closely linked in river hydrology: 219

variations in water stage give indications of variations in discharge, i.e. the lower the 220

river stage the lower is the discharge and conversely (Herschy, 1995). Thus, in this study, 221

we considered that the highest water level recorded in spring during snowmelt 222

corresponds to the annual peak flow. 223

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The OBS 3+ uses its sideways-facing optics to emit a near-infrared light beam 224

into the water. It then measures the light that bounces back from the suspended particles 225

in the water. The sensor head was cleaned at every site visit while servicing the station to 226

prevent accumulation of algae that could have invalidated the data (Schoellhamer 1993). 227

In fact, we never observed any algae on the sensor. It is assumed that sediment particle 228

colour and reflectivity are not significant factors affecting sensor accuracy (Schoellhamer 229

and Wright 2003). However, besides the choice of the turbidity range by the operator, the 230

color and reflectivity of sediment are unique for each river and consequently, no absolute 231

comparison with NTU curves in other rivers is possible. 232

When there is an ice cover, pressure and roughness of the ice and snow cover 233

make the sensors readings inaccurate. Some sections of the Sheldrake River are less than 234

1 m deep and should be frozen to the bed in winter time. Thus, in this study, the winter 235

flow is considered as negligible. Curves of water stage and turbidity start with the break 236

up and end with the first signs of ice on the river when air temperature drops under 0⁰C 237

for several consecutive days. 238

Water analyses were made by the Laboratoire de l’INRS, Centre Eau, Terre, 239

Envionnement and by the Environex laboratory in Quebec City. Total suspended solids 240

(TSS) were measured from filtration through a 0.45 µm filter and weighting after drying. 241

Air temperature and precipitation data were provided by an automated regional 242

meteorological station, operated by the Centre d’études nordiques, located in Tasiapik 243

valley, at 125 m above sea level (a.s.l.), near the village of Umiujaq 8 km south of the 244

Sheldrake River (See Fig. 1 for location) (CEN 2014). Through the area, the altitudinal 245

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gradient is small (coastal range: < 400 m a.s.l.; inland Canadian Shield: ̴ 200 to 250 m 246

a.s.l.) whereas vegetation cover types and variable snow cover conditions maintain soil 247

surface temperatures within a limited range around 0 °C (Ménard et al. 1997). Similarly, 248

convective precipitation are highly improbable within such a small catchment with low 249

relief amplitude. 250

Sheldrake Lake, where the Sheldrake River originates, has not been included in 251

the delineation of the Sheldrake River catchment area for this study. The lake is a 252

sedimentation basin which receives sediments from its surrounding catchment; which has 253

an estimated area of 91 km2. We found no visible signs of erosion on the lake catchment, 254

as it is surrounded by wetlands and bedrock outcrops (Jolivel and Allard 2013). Thus, the 255

suspended sediment released by the lake as a potential cause of turbidity of the Sheldrake 256

River is considered as negligible. This is confirmed by the clarity of the water of the lake 257

outlet over which we flew each summer. However, water input coming from the 258

Sheldrake Lake accounts for an unmeasured fraction of the Sheldrake River flow. 259

Spring flood 260

The freshet flood event is clearly visible on hydrographs (Fig. 5). We estimated its 261

starting date as the period during snowmelt time (several consecutive days) when the 262

water level curve starts to rise after being stable and low during winter. The end of 263

snowmelt time is associated with the resumption of river flow to its summer stage and 264

when air temperature remains above 0°C. For each year separately, we estimated this 265

summer stage by calculating average water stage in July and August. In 2014, water level 266

data ended on 4 October, the date of last visit on site for downloading the datalogger. 267

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Turbidity/Total Suspended Solids calibration 268

Our interpretations of turbidity events over time and with climate parameters use the 269

NTU values. Nevertheless an attempt to approximately assess the concentration of 270

suspended sediment was made by calculating a NTU-TSS relationship: 271

TSS = 0.46NTU + 2.457 (1) 272

with r2= 0.88, p< 0.001, n=22. Due to the remoteness of the site, only three samples in 273

the calibration data were obtained under conditions of relatively high TSS concentration 274

(≥14 mg.L-1) (Fig. 6). The NTU/TSS relationship was used here only to provide a crude 275

estimate of sediment concentration since no TSS samples were recovered at very high 276

NTU and no robust discharge data are available. However, the correlation NTU/TSS with 277

this instrument is reported to be linear up to 4000 NTU (Downing 2006). This is 278

particularly true for fine-grained sediments, such as glaciomarine silt and clay (Lewis 279

1996). In this study, the spectrum of calibration NTU/TSS only covers the range 3-15 280

NTU; maximum turbidity peaked at 160 NTU in summer 2010. Thus, we assume that the 281

correlation NTU/TSS is linear all along the range covered by this study (3-160 NTU). 282

The cable linking the submerged instrument and the datalogger onshore was twice 283

severed by ice, and therefore complete summer coverage of turbidity was obtained only 284

in 2010, 2013 and 2014. 285

As 85% of the catchment is covered by silty clay and numerous pools in bedrock and 286

reaches on boulder beds very likely prevent or limit bed-load transport, turbidity of the 287

Sheldrake River is considered as a reliable indicator of sediment transport. 288

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Both water level and turbidity data are expressed as mean daily values. This time unit 289

also corresponds to the time unit used to calculate the soil thaw regime, i.e. degree-days. 290

Active layer depth 291

The rate of active layer thawing during a summer is key to many erosional and 292

sediment releasing processes. Downward migration of the thaw front is rapid at the 293

beginning of the season, following Stefan’s equation of heat transfer with phase change. 294

Typically, the pace of thawing is fast in June and July and it slows down towards the end 295

of summer. Cryoturbation during thaw brings loose sediments to the soil surface in the 296

center of frostboils (Mackay and Mackay 1976; Egginton and Dyke 1982), which are 297

extensive in the studied catchment on lithalsas and permafrost plateaus. Generation of 298

high pore pressures at the thawing front when the thawing rate is fast is the main trigger 299

mechanism of active layer detachment slides (Lewkowicz and Harris 2005b). Melting of 300

residual snowbanks in gullies is faster during warm summer spells, generating erosion 301

(Jolivel and Allard 2013). Therefore, we considered the air temperature curve, the 302

cumulative curve of thawing degree-days and the derived curve of thaw front progression 303

as the best integration of controlling variables of the geomorphic processes that 304

potentially release sediments in summer. 305

To evaluate the thermal regime of the active layer in 2010, 2013 and 2014, a one-306

dimensional heat transfer model, TONE was used (Goodrich 1978). This numerical 307

model is widely used for active layer and permafrost modelling with good results and 308

allows researchers to simulate the evolution of the active layer under different 309

environmental conditions (Riseborough et al. 2008; Bouchard 1990; Zhang et al. 2008; 310

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Barrette 2010; Larouche 2010). The thermal model is driven by cumulative thawing 311

degree-days gathered from the meteorological station. Clay and silt thermal parameters of 312

the model’s PRAM routine with a water content of 20% and only one stratigraphic layer 313

were applied in the simulations as in other studies in the region (Buteau et al. 2004). 314

Grain size composition is rather homogenous throughout the clay soils of the region 315

(Calmels 2005). The sediment is composed of fine silt and clay deposited in the Tyrrell 316

sea during deglaciation. N-factors measured by Ménard et al. (1998) in the region were 317

applied to account for the buffer effect of snow cover and vegetation between the 318

atmosphere and the soil surface. They measured a thawing n-factor of 1.17 and a freezing 319

n-factor of 0.64 for a typical lithalsa composed of postglacial silt with a lichen cover. 320

Those values are therefore generally applicable to most of the permafrost patches and 321

landforms across the Sheldrake River catchment. Soil temperatures profiles were 322

simulated at 10 cm intervals (model node spacing). 323

Once the soil temperatures for 2010, 2013 and 2014 were calculated with the 324

model, a piecewise cubic hermite interpolating polynomial (PCHIP) allowed estimations 325

of the depth of the thaw front according to the calculated depth of the 0⁰C isotherm over 326

the thawing period (L'Hérault 2009). The downward thaw of the active layer was 327

simulated for the summers of 2010, 2013 and 2014, i.e. from 1 May to 1 October, which 328

is considered for this study as the day of the maximum depth reached by the 0°C 329

isotherm. 330

Results 331

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Seasonal variations of water stage 332

The breakup generally occurs in late April or early May. The freshet high stages, 333

i.e. between break up and drawdown to summer stage, generally last about 8 weeks, 334

except in 2014 (5 weeks) which had also the earliest hydrological summer condition 335

recovery date of the six years of the study period (Table 1). The highest water stage, 336

corresponding to the annual discharge peak, generally occurs between end of May and 337

mid-June, except in 2013 where the river flow was maximum on 3 May (Table 1, Fig. 5). 338

The Sheldrake River reaches its summer flow regime at the end of June or at the 339

beginning of July. Starting in late August, successive rainfall events generate an increase 340

in water flow until the river freezes. 341

Turbidity regime over the thawing season 342

Turbidity variations during the snowmelt period 343

In 2010, the snowmelt period lasted from 25 April to 17 June (Table 1, Fig. 7). 344

During this period, daily NTU values varied between 5 and 15 and peaked at 22 on 6 345

May (Fig. 7). This value was reached after five days with average daily air temperatures 346

fluctuating around 0°C. No significant rain event occurred and the active layer began to 347

thaw only after this date, i.e. from 8 May onwards (Fig. 8). During the spring flood, 348

turbidity values did not increase significantly. However, in early June, five days after the 349

maximum water stage of the freshet flood, turbidity increased and stayed at a level of ~7 350

NTU, compared with a mean 4 NTU in May. From 25 April to 17 June, only two days of 351

light rain were recorded (19 May: 4 mm ; 29 May: 2 mm) (Fig. 7). 352

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In 2013, the period of high water levels associated with the snowmelt period 353

lasted 59 days and ended at the end of June (1 May to 28 June) (Table 1, Fig. 9). Over 354

this period, only 6 days of light rain totaled less than 40 mm. Daily NTU values generally 355

varied between 3 and 20 NTU. Contrary to 2010, the highest water stages were associated 356

with a higher turbidity, ranging from 14 to 20 NTU from 1 May to 5 May. During the 357

first half of May 2013, two major peaks of turbidity occurred. The first one (average 18 358

NTU from 1 to 5 May) was correlated with the peak of the freshet flood and followed the 359

first rain of the year on 28-29 May (25 mm of rain). The second peak of turbidity was of 360

higher amplitude (average 80 NTU from 11 to 14 May, with a peak of 135 NTU reached 361

on 12 May) and was independent of rain and river stage. This peak coincides in time with 362

the fast initial thawing of soils to a depth of about 15 cm (Fig. 10). 363

From 8 to 12 June 2013, turbidity averaged at 59 NTU, with a peak at 80 NTU on 364

10 June. Before and during this period, no significant rain was recorded (4 mm on 10 365

June). However, it occurred simultaneously with a significant increase in air temperature 366

(+15 °C) and a rapid deepening of the active layer (Figs. 9 and 10). It was also correlated 367

with the onset of the period when air temperature maintains itself continuously above 368

0°C, as shown by the increase of the slope of the degree-days curve (Fig. 10). No change 369

in precipitation and thus, in water stage was noticed. 370

In 2014, the snowmelt period lasted 42 days, from 4 May to 14 June, which is 371

considerably shorter than in the other years (Table 1, Fig. 11). During this period, 372

turbidity averaged 8 NTU and again, no significant rain occurred. However, two peaks of 373

turbidity were recorded on 20 May (17 NTU) and on 24, 25 and 26 May (respectively 32, 374

29 and 21 NTU). As during the 2013 spring high water levels, pulses of high turbidity 375

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corresponded with relative high river stage and coincided with rapid and short increases 376

of the active layer depth (Figs. 11 and 12). Meanwhile, air temperatures increased from -377

1°C on 19 May to 11°C on 22 and 23 May, accelerating snowmelt, and the thaw front 378

quickly reached a depth of ̴ 30 cm. 379

Turbidity variations during summer 380

In 2010, the Sheldrake River water level lowered to reach its mean summer level 381

on 17 June (Table 1, Fig. 7). Before that date, as seen above, snowmelt driven high levels 382

did not strongly influence the turbidity. Whereas the period of higher river stage and thus 383

maximum discharge occurred during the spring freshet, the turbidity in the river reached 384

its summer maximum in July. Indeed, between 30 June and 23 July 2010 (24 days), NTU 385

values averaged at 83 with a peak at 160 on 10 July, while the rest of the 2010 thawing 386

period registered an average turbidity of only 4.2 NTU (Fig. 7, Table 2). The 2010 387

turbidity maximum was reached on 10 July following several days of light rain (Fig. 7). 388

On the other hand, the NTU curve rises three days following the increase of the average 389

daily air temperatures from 0-5 °C up to above 10 °C for the rest of the turbidity period 390

(Fig. 7). The turbidity peak also occurred at a time when the thaw rate was the fastest in 391

the soil (Fig. 8). 392

In 2013, summer river stage was reached on 28 June (Table 1, Fig. 9). An 393

increase in turbidity similar to 2010 occurred in midsummer but with a lesser magnitude. 394

The general period of high turbidity lasted 19 days, i.e. 15 July to 2 August, with a daily 395

average NTU of 33, while the rest of the thawing period registered an average turbidity of 396

only 8.2 NTU (Table 2). During this period, 10 days with light to moderate rain were 397

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registered (4 to 13 mm/day). The two main peaks of turbidity, on 17 July and 29 July, 398

were associated with two (10 mm) and three (25 mm) rainy days, respectively. As a 399

consequence the higher water levels, and thus discharge, were also correlated with 400

turbidity and generated small peak events (Fig. 9). 401

In 2014, several pulses of turbidity were recorded in June, e.g. 5 June with 26 402

NTU, 12 June with 18 NTU and from 23 to 27 June, with an average NTU of 31 and a 403

peak of 64 NTU reached on 26 June. On 12 June, the turbidity pulse was associated with 404

an increase in air temperature from 9 to 21°C and very light rain (2 mm). The two other 405

episodes were not associated with rain or significant changes in air temperatures. 406

However two days with air temperatures above 15°C preceded the turbidity pulse which 407

occurred from 23 to 27 June (Fig. 11). 408

In early July 2014, a turbidity increase, with a maximum of 19 NTU on 4 July, 409

was recorded during a rainy and warm week (74 mm of rain between 28 June and 5 July) 410

(Figs 11 and 12). Then, as in 2010 and 2013, a period of high turbidity, distinctive from 411

the rest of the thawing season, occurred in midsummer, from 12 July to 28 July 2014 412

(Fig. 11). During these 17 days, average turbidity was 72 NTU, with a peak of 141 NTU 413

on 20 July, discharge was low and stable and air temperatures fluctuated between 7 and 414

21°C (Table 2). The onset of this event was associated with an increase in air temperature 415

from 8 to 13°C and light rain (5 mm) two days before the increase in NTU. However, 416

from 20 to 24 July, 55 mm of rain were recorded and coincided with the maximum of 417

turbidity on 20 July (Fig. 9). 418

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In 2010, 2013 and 2014, small peaks of turbidity that occurred in August were 419

directly correlated with precipitation and river levels (Figs. 7, 9 and 11). In autumn, air 420

temperatures decreased until freeze-up and no turbidity event like those recorded in 421

spring and midsummer occurred despite an elevation of water stage and thus an increase 422

of discharge. several peaks of turbidity unrelated to high flows are noticeable on the 423

hydrographs. 424

Fig. 13 shows relationship between turbidity and water level and between 425

turbidity and air temperature. We selected NTU values <10 in order to remove small 426

peaks caused by light rains. Thus, there exists no correlation between turbidity and water 427

stage. In 2013 and 2014, there appears to be a positive relation between turbidity and air 428

temperature whereas 2010 showed no trend (Fig. 13). 429

The major turbidity period occurred when the thaw front reached 40% of its 430

maximum end of summer depth in 2010, 60% in 2013 and 64% in 2014 (Figs. 8, 10 and 431

12). In July 2010, the high turbidity period coincided with the steepest part of the 432

cumulative thawing degree-days curve, i.e. when the thaw front was still rapidly 433

penetrating in the ground (Fig. 8). In 2013, except for two short but intense pulses of 434

suspended sediment in May and June, the major period of high turbidity occurred in mid-435

July-early August as the thaw front was progressing at a sustained high rate (Fig. 10). In 436

2014, during the second part of July, turbidity increased significantly and dominated the 437

thawing season, also coinciding with an increase of thaw rate and a period of faster rise in 438

the cumulative degree-days (Fig. 12). Except for some small peaks of turbidity induced 439

by rain events, NTU values generally stayed low in September and October as 440

temperatures fell and the thaw front tended to stabilize before freeze back. 441

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Discussion 442

Hydrological regime of a typical subarctic river 443

The flow regime of the Sheldrake River is typical of those of the subarctic 444

environment and is closely linked to precipitation and air temperature (Déry et al. 2005). 445

Water stages and relative water flow are high at the end of spring (snowmelt); they rise 446

again in autumn as heavy rains precede the coming of winter, contrasting with low flow 447

in summer (July and August) and practically no flow in winter. Snowmelt runoff is 448

clearly the dominant hydrologic event, surpassing peaks generated by rain event, as 449

observed in other subarctic catchments (Déry et al. 2005; Jing and Chen 2011). In details, 450

the magnitude and timing of water level fluctuations varied significantly each year and 451

from one year to another. For example, the highest water stage reached 366 cm in 2012 452

but only 313 cm in 2010 and 2011. Peak flow recession at the recording station does not 453

show any irregularity that could have arisen due to some disturbing event such as a 454

temporary ice-jam up-river. Moreover, during our early summer visits, we never 455

observed any signs of ice jams (remaining ice blocks and boulder barricades, wood debris 456

on banks) along the river. 457

The Great Whale River (GWR), located 150 km south of the Sheldrake River, is 458

the closest river where a similar study was done previously, allowing some broad 459

qualitative comparisons on the spring high water stage (Hudon et al. 1996). GWR is the 460

only other river with a functioning gauging station in the region (operated by the Centre 461

d'Expertise Hydrique du Québec). Comparisons are only indicative since the hydro-462

geomorphological regime of the GWR is not studied and its catchment is considerable. 463

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GWR basin is 43,200 km2 and its mean annual discharge is 600 m3 s-1 (Déry et al. 2005; 464

Hülse and Bentley 2012). Compared to the GWR, the highest flow in Sheldrake River 465

occurred generally one to two weeks later (Table 3), average annual air temperatures 466

being 1°C colder in the Sheldrake River area than in the GWR area, which delays the 467

freshet (Environment Canada 2013). 468

An exception occurred in 2013 when the highest water stage of the Sheldrake 469

River preceded the one of the GWR by 18 days. Five days before the peak flow of the 470

Sheldrake River, an early spring storm (25 mm of rain and daily air temperatures of 5°C) 471

occurred in the Umiujaq area. The meteorological station situated at the mouth of the 472

GWR, in Kuujjuarapik, registered a coincident warm spell in late April 2013 but with 473

lighter rain than in the Sheldrake River area (14 mm of rain spread over 3 days) 474

(Environment Canada 2013). This comparison between the Sheldrake River and the 475

GWR should be interpreted with caution however. 476

In periglacial regions, runoff rates and discharge response time tend to wane over 477

the summer period as snow depletes, seasonally frozen soils thaw and the active layer 478

increases in thickness over permafrost areas, allowing water infiltration and soil water 479

storage (McNamara et al. 1998; Kane et al. 2000; Carey and Woo 2001; Hinzman et al. 480

2003; Woo et al. 2008). For the Sheldrake River, the snowmelt and the thawing depth in 481

soils (permafrost or not) control the seasonal trend of the river flow. In spring, 482

precipitation has only a limited and temporary influence on river stage and discharge. 483

Then, after the freshet high water stages, our hydrographs and field observations suggest 484

that summer river flow depends on both rain storm frequency and intensity and water 485

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flow to the main stream after a short delay by retention in soils, ponds, lakes and 486

wetlands, as observed in southern temperate Québec (James and Roulet 2009). 487

Rain generally stops at the end of October and is replaced by snow when air 488

temperature drops under 0⁰C. However, some variations of discharge still occur in the 489

first half of November. The presence of the Hudson Bay cools the coastal zone, whereas 490

temperature remains warmer inland (Payette and Rochefort 2001; Bhiry et al. 2011; 491

Jolivel and Allard 2013). Autumn rains inland likely raise the level of the Sheldrake Lake 492

therefore sustaining a relatively high flow in the river until freeze-up. But turbidity does 493

not increase during that autumn period of high flow, suggesting that no erosion is 494

generated in soils and on slopes in the catchment. 495

Turbidity and sediment transport 496

Turbidity and sediment transport during spring high water levels 497

In rivers flowing in permafrost-free watersheds or in Arctic catchments with 498

permafrost, transport of sediment is generally a function of discharge and dependent on 499

rainfall and on the amount of snow accumulation during the previous winter (Kriet et al. 500

1992; Hudon et al. 1996; Braun et al. 2000; Forbes and Lamoureux 2005). However, in 501

the Sheldrake River, during the spring, several peaks of turbidity unrelated to high flows 502

are noticeable on the hydrographs. The absence of correlation between turbidity and 503

water stage (Fig. 13) clearly demonstrates that the hydrological regime is not the 504

dominant factor of suspended sediment transport in the catchment. However the rate of 505

soil thawing driven by cumulative warming (here expressed as degree-days) stands out as 506

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a very probable driver of turbidity. Otherwise, there is considerable differences of water 507

level from one year to another, reflecting an important interannual variability. For 508

example, in 2010 and 2014, the spectrum of turbidity was broader than in 2013 which 509

showed larger variation of water level. Even if we can not totally exclude a slight 510

displacement of the mooring during precedent breakup, this further suggests that water 511

level of the Sheldrake River is linked to winter snow accumulation in the catchment 512

and/or the level of the Sheldrake Lake in Spring. 513

In winter, the frozen soils, the snow cover and the absence of runoff impede 514

overland flow and soil erosion as well as river flow and bank erosion (Hudon et al. 1996; 515

Burn et al. 2004). In spring, the rise in solar radiation and above freezing air temperatures 516

rapidly melt the snow cover and result in an increase in water flow and transport (Hudon 517

et al. 1996). However, during the freshet flood, a relatively low turbidity indicates that 518

transport of TSS likely remains low, because some channels stay armored with ice and 519

snow and soils are still frozen, as observed in other catchments in the Arctic and 520

Subarctic (McDonald and Lamoureux 2009). 521

Inter-annual fluctuations of turbidity are significant and can be partly explained 522

by the length of the snowmelt high flow period, which was ̴ 2 weeks shorter in 2014 than 523

in the five other years. 524

In May and June 2013 and 2014, i.e. during high water levels, 3-4 days pulses of 525

sediment were recorded independently from temporary increase in water stage. These 526

short events are likely due to episodic landslide activity, as observed in Arctic catchments 527

(Lewis et al. 2005; Lamoureux and Lafrenière 2009). The increase in air temperature 528

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causes rapid deepening of the active layer and trigger shallow active layer landslides 529

along the river banks. Such mass movements are widespread in the Sheldrake catchment 530

(Jolivel and Allard 2013). Such an event occurred in spring 2010 (Fig. 14) before 15 June 531

(date of a visit in the field). This landslide was likely responsible of the turbidity peak 532

recorded on 6 May 2010, even if no direct correlation was made in the field, the precise 533

date of the landslide occurrence being unknown. 534

The mid-summer turbidity period 535

Mid-summer is the time of the year when frostboils on clay and silt rich soils are 536

reactivated by excess pore water pressure caused by soil thawing, bringing to the surface 537

new sediments that can be washed away on the slopes of the lithalsas by surface runoff 538

(Shilts 1978; French 2007) (Fig. 15A). In the Sheldrake River catchment, where 1100 539

lithalsas covering a cumulated area of 3 km2 were inventoried in 2009 (Jolivel and Allard 540

2013), this runoff is shown by the milky colour of thermokarst ponds (Fig. 15B). During 541

rain events, thermokarst ponds overflow into channels and feed the fluvial system (Jolivel 542

and Allard 2013). Moreover, rapid thawing of the active layer causes small active layer 543

detachment failures on the side slopes of palsas and lithalsas releasing silt and clay into 544

tributary streams (Jolivel and Allard 2013). Larger slides occur along gullies and 545

immediately along river banks (Fig. 15C and 15D); activity and contribution of gullies to 546

the fluvial load are confirmed by the presence of small deltas at the confluence with 547

streams or with the Sheldrake River during low stages (Fig. 15E). Simple observations of 548

rivers during helicopter flights in the silt-rich permafrost-dominated region east of 549

Hudson Bay also reveals the summer high turbidity of small rivers flowing into Hudson 550

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Bay or as tributaries of major rivers such as the Nastapoka River, suggesting that the 551

processes may be similarly over the wider region. 552

Precipitation recorded during the three mid-summer turbidity periods generate 553

surface runoff and facilitate the washout of the thaw-released sediments as shown by the 554

time correspondence of small turbidity peaks with rain events. Indeed, hot summer 555

temperature with moderate rains can trigger earth flows, active layer detachments slides 556

and gully erosion, delivering considerable volume of sediment into Subarctic fluvial 557

systems, while overland flow can also simply wash out sediments released by soil 558

thawing and thermokarst (Lamoureux 2000; Lewkowicz and Harris 2005a; Lamoureux 559

and Lafrenière 2009; Lewis et al. 2012). 560

After the mid-summer maximum, turbidity rapidly comes back to lower values. 561

Sediment-releasing and periglacial erosion processes slow down as the thaw front nearly 562

stabilizes near the base of the active layer over permafrost areas and when soils are 563

finally thawed in the permafrost free areas of the catchment. Finally, in October, when air 564

temperatures get closer to 0°C, turbidity comes back to very low values despite higher 565

discharge than in midsummer. Furthermore, the well-defined turbidity-concentration 566

relationship suggests that particle size does not change significantly with variable flow 567

rates. This confirms that the eroded soils are well sorted and fine-grained, i.e., typically 568

belonging to the ice-rich marine silty clays, characterized by their homogeneity through 569

the catchment (Calmels 2005) and currently affected by thermokarst. 570

Implications for sediment transport 571

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Daily turbidity peaks recorded during the three summers corresponded to 572

sediment concentrations of 76 g m-3 in 2010, 29 g m-3 in 2013 and 67 g m-3 in 2014. SSC 573

reported in the Sheldrake River is far less than SSC measurements made in small Arctic 574

catchments immediately downstream from large mass movements. For example, hourly 575

SSC reached 60,800 g m-3 in a small creek on Ellesmere Island (Lewis et al. 2005). 576

Short-lived peaks up to 800 g m-3 were recorded in a small creek on Melville Island 577

(Lamoureux and Lafrenière 2009) and in rivers on the Peel Plateau (Kokelj et al. 2013) 578

downstream of large active layer slides and retrogressive thaw slumps. Pulses of TTS 579

generally occur within two or three days of landslides (Lewis et al. 2005; Lamoureux and 580

Lafrenière 2009). However, those conspicuous processes also occur in summer and are 581

triggered by soil thawing under warming temperature. In the low relief and discontinuous 582

permafrost of the Sheldrake River catchment, the thawing processes are smaller in scale 583

and more diffusively scattered in the landscape. A dilution effect associated with the 584

contribution of low-SSC water from the outlet of Sheldrake Lake should also be 585

mentioned. 586

The 24 days period in 2010, the 19 days period in 2013 and the 17 day periods of 587

sustained high turbidity recorded in 2014 in the Sheldrake River catchment are likely 588

related to frostboil activity, many small soil instabilities, landslides on permafrost slopes 589

and along riverbanks, and to erosion of gullies in the catchment. In this periglacial 590

environment, all those mass wasting processes are driven by the rate of summer warming 591

and thaw of the active layer especially in the ice-rich permafrost in marine silty clays. 592

Conclusion 593

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Our results indicate that the sedimentary regime of a river that drains geosystems and 594

ecosystems on fine-grained permafrost, such as the Sheldrake, can in large part be 595

affected by summer thaw rates in the active layer and associated thermokarst processes. 596

The annual peak flow of the Subarctic Sheldrake River shown by the annual highest 597

water stage typically occurs during the snowmelt period. However large pulses of 598

suspended sediments of a few weeks of duration occur in mid-summer, during the 599

steepest part of the cumulative thawing degree-days curve. Light rains in summer also 600

drive the flow regime that assumes transport of the suspended sediment load during the 601

high turbidity period. Indeed, in 2010, 2013 and 2014 the period of high turbidity lasted 602

for ̴ 3 weeks and occurred in July when warm summer air temperatures transferred heat 603

into the active layer and provoked the rapid penetration of the thaw front. This is likely a 604

time of the year when frost boils on permafrost and active layer detachment failures on 605

the slopes of palsas and lithalsas are the most active, releasing fine-grained sediments that 606

are rain-washed at the soil surface and transported to the ponds, gullies and stream 607

system. Thaw front penetration in steep river banks likely also favors slumping and 608

sediment release. We suggest that these mid-summer episodes of sediment transport that 609

correspond with warm days and light rains carry a large part of the annual suspended 610

sediment load although further more investigations are needed to quantify these summer 611

exports. This is important because it would imply that maximum sediment transport in an 612

active thermokarst region dominated by fine-grained soils may be primarily thermally-613

driven rather than hydrologically-driven. More intensive observations within the 614

catchment will help better appraise the scale and extent of the thaw-related sediment 615

delivery processes that are diffusely distributed over the territory. 616

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617

Acknowledgements 618

Climate data to support this article are from the Centre d'études nordiques. This 619

work received financial support from grants to M. Allard from ArcticNet and the Natural 620

Science and Engineering Research Council of Canada. The Centre d’études nordiques of 621

Université Laval provided important logistical support. We thank Denis Sarrazin for the 622

installation and maintenance of the gauging station, Marc-André Ducharme for field 623

assistance, and Sarah Aubé-Michaud for help with the Tone model. The comments of 624

Mickael Lemay, Daniel Fortier, Patrick Lajeunesse and Guillaume St-Onge greatly 625

improved the manuscript. We are also grateful to the Inuit community of Umiujaq for its 626

generous hospitality. We finally thank the two anonymous reviewers. 627

628

629

630

631

632

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Table legends

Table 1 Features of the period of spring high water stage of the Sheldrake River from 2009 to 2014. Table 2 Comparison between the summer periods of turbidity in 2010, 2013 and in 2014. Table 3. Comparison of the date of the peak flow in the Great Whale River and in the Sheldrake River, from 2009 to 2014. For the Sheldrake River, the date of the peak flow corresponds to the day where the annual highest water stage was recorded.

Figure legends

Figure 1 Location of the study area and photograph of the Sheldrake River flowing through an area of permafrost plateaus (foreground), and finally flowing to Hudson Bay across rock outcrops. VDT indicates the location of the meteorological station, in the Vallée des Trois (also called Tasiapik)

Figure 2 Two tributaries of the Nastapoka River. Note the high turbidity of the streams reflecting high thermokarst activity in catchments. Figure 3 The Sheldrake River catchment with the distribution of surficial deposits, the position of the tree line and the organization of the drainage network. Sheldrake Lake does not appear on this map (see Fig.1 for location). Figure 4 Schematic representation of the positioning of the gauging station and photograph of its location on a bank of the Sheldrake River.

Figure 5 Hydrographs for 2009, 2010, 2011, 2012, 2013 and 2014. Water level record starts with the first signs of the break up (see in text) and end at mid-November (except for 2013 and 2014, due to lack of data). Figure 6 Relationship between turbidity (NTU) and total suspended sediment (TSS) (g m3). Figure 7 Water level and turbidity regime of the Sheldrake River with air temperature and rainfall events in 2010: rain (mm d-1), air temperature (°C), water level (cm) and turbidity (NTU). Data are presented on a daily basis. Figure 8 Turbidity (NTU), thaw front depth (cm), rain (mm d-1) and cumulative degree-days ≥ 0⁰C from 1 May to 1 October 2010.

Figure 9 Water level and turbidity regime of the Sheldrake River with air temperature and rainfall events in 2013: rain (mm d-1), air temperature (°C), water level (cm) and turbidity (NTU). Data are presented on a daily basis.

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Figure 10 Turbidity (NTU), thaw front depth (cm), rain (mm d-1) and cumulative degree-days ≥ 0⁰C from 1 May to 1 October 2013. Figure 11 Water level and turbidity regime of the Sheldrake River with air temperature and rainfall events in 2014: rain (mm d-1), air temperature (°C), water level (cm) and turbidity (NTU). Data are presented on a daily basis. Figure 12 Turbidity (NTU), thaw front depth (cm), rain (mm d-1) and cumulative degree-days ≥ 0⁰C from 1 May to 1 October 2014. Figure 13 (A) Relationship between turbidity and air temperature; (B) Relationship between turbidity and water stage. One point represents one day during the thawing season of 2010, 2013 and 2014. Days when turbidity was less than 10 NTU were excluded to limit the effects of light rain. The regression line and the coefficient of correlation represent only turbidity values ≥10 NTU.

Figure 14 Typical landslide on the bench of the Sheldrake River. The landslide probably triggered on 6 May 2010 and caused a peak of turbidity.

Figure 15 Thermokarst as source of suspended sediment in the Sheldrake River catchment (A): active and flowing frostboil on the slope of a permafrost plateau (1 meter length) (photo: Denis Sarrazin); (B): Typical thermokarst ponds, the strong turbidity is caused by frostboils activity and surface runoff; (C): Network of gullies between permafrost plateaus (Jolivel and Allard, 2013); (D): typical landslide on a permafrost bank along the Sheldrake River; (E): small delta at the confluence between a thermokarst gully and the Sheldrake River.

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Table 1. Features of the period of spring high water stage of the Sheldrake River from 2009 to

2014.

Start/end

snowmelt high

water stage (date)

Length snowmelt

(number of days)

Peak of the freshet

flood (date)

Highest water

stage (cm)

2009 16 May/10 July 56 14 June 333

2010 25 April/17 June 54 27 May 313

2011 12 May/2 July 52 5 June 313

2012 4 May/1 July 59 7 June 366

2013 1 May/28 June 59 3 May 331

2014 4 May/14 June 42 28 May 338

Average 6 May/27 June 53 30 May 332

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Table 2. Features of the mid-summer turbidity periods

Context fitting with the onset of the period of turbidity

Period of high

turbidity

Number of

days

Daily turbidity

min./max./av.

(NTU)

Average turbidity for the

rest of the thawing season

(NTU)

Rain Water stage

Air temperature

30 June - 23 July 2010 24 24/160/43 4.2 No preceding rain event Decrease Increase from 0-5

to 13⁰C

15 July - 2 August

2013 19 15/57/33 8.2

2 days of light rain (8 and 6 mm) preceding the onset

of the turbidity period

Stable Generally rising

above 10⁰C

12 July - 28 July 2014 17 11/141/72 8.2 Light rain (5 mm) 2 days before the onset of the

turbidity period

Stable Increase from 8-9

to 13 ⁰C

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Table 3. Comparison of the date of the peak flow in the Great Whale River and in the Sheldrake

River, from 2009 to 2014. For the Sheldrake River, the date of the peak flow corresponds to the

day where the annual highest water stage was recorded.

Great Whale River Sheldrake River

Date Date

2009 8 June 14 June

2010 14 May 27 May

2011 20 May 5 June

2012 26 May 7 June

2013 21 May 3 May

2014 23 May 28 May

Average 24 May 30 May

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Figures

Fig. 1. Location of the study area and photograph of the Sheldrake River flowing through

an area of permafrost plateaus (foreground), and finally flowing to Hudson Bay across

rock outcrops. VDT indicates the location of the meteorological station, in the Vallée des

Trois (also called Tasiapik).

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Fig. 2. Two tributaries of the Nastapoka River. Note the high turbidity of the streams

reflecting high thermokarst activity in catchments.

.

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Fig. 3. The Sheldrake River catchment with the distribution of surficial deposits, the

position of the tree line and the main drainage network. Sheldrake Lake does not appear

on this map (see Fig.1 for location).

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Fig. 4. Schematic representation of the positioning of the gauging station and photograph

of its location on a bank of the Sheldrake River.

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Fig. 5. Hydrographs for 2009, 2010, 2011, 2012, 2013 and 2014. Water level record

starts with the first signs of the break up (see in text) and end at mid-November (except

for 2013 and 2014, due to lack of data).

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Fig. 6. Relationship between turbidity (NTU) and total suspended sediment (TSS) (g m3).

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Fig. 7. Water level and turbidity regime of the Sheldrake River with air temperature and

rainfall events in 2010: rain (mm d-1

), air temperature (°C), water level (cm) and turbidity

(NTU). Data are presented on a daily basis.

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Fig. 8. Turbidity (NTU), thaw front depth (cm), rain (mm d-1

) and cumulative degree-

days ≥ 0⁰C from 1 May to 1 October 2010.

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Fig. 13. (A) Relationship between turbidity and air temperature; (B) Relationship

between turbidity and water stage. One point represents one day during the thawing

season of 2010, 2013 and 2014. Days when turbidity was less than 10 NTU were

excluded to limit the effects of light rain. The regression line and the coefficient of

correlation represent only turbidity values ≥10 NTU.

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Fig. 14. Typical landslide on the bench of the Sheldrake River. The landslide probably

occurred on 6 May 2010 and caused a peak of turbidity.

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Fig. 15. Thermokarst as source of suspended sediment in the Sheldrake River catchment

(A): active and flowing frostboil on the slope of a permafrost plateau (1 meter length)

(photo: Denis Sarrazin); (B): Typical thermokarst ponds, the strong turbidity is caused by

frostboils activity and surface runoff; (C): Network of gullies between permafrost

plateaus (Jolivel and Allard, 2013); (D): typical landslide on a permafrost bank along the

Sheldrake River; (E): small delta at the confluence between a thermokarst gully and the

Sheldrake River.

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