Groundwater dynamics within a watershed in the discontinuous permafrost zone near

Umiujaq (Nunavik, Canada)

Jean-Michel Lemieux1,2*, Richard Fortier1,2, Renaud Murray1,2,ǂ, Sophie Dagenais1,2,§, Marion

Cochand1,2, Hugo Delottier1, René Therrien1,2, John Molson1,2, Alexandre Pryet3, Masoumeh

Parhizkar1,2

1. Département de géologie et de génie géologique, 1065 avenue de la Médecine, Université Laval, Québec (Québec), Canada, G1V 0A6.

2. Center for Northern Studies, 2405 rue de la Terrasse, Université Laval, Québec (Québec), Canada, G1V 0A6.

3. EA 4592 Georessources & Environnement, Bordeaux INP and Univ. Bordeaux Montaigne, ENSEGID, 1 allée F. Daguin, 33607 Pessac cedex, France

*Corresponding author. +1 418-656-7679, jmlemieux@ggl.ulaval.ca

ǂ Now at Golder Associates, 1170 Boulevard Lebourgneuf, (Québec), Canada, G2K 2E3

§ Now at WSP, 1600 Boulevard René-Lévesque, (Montréal), Canada, H3H 1P9

ABSTRACT

Groundwater distribution and flow dynamics were studied in a small watershed located in the

discontinuous permafrost zone at Umiujaq in Nunavik (Québec), Canada, to assess the

seasonal variations and perform a quantitative analysis of the water cycle in a subarctic

watershed. Due to the complexity of the subsurface geology within the watershed, an

integrated investigation was instrumental to provide a detailed understanding of the

hydrogeological context as a basis for the water balance. Based on this water balance, for

the two studied hydrological years in 2015 and 2016, the average values are 828 mm for 1

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precipitation, 337 mm for evapotranspiration, 46 mm for snow sublimation, 263 mm for runoff,

183 mm for groundwater exchange (losses with other aquifers outside the watershed), and 0

mm for change in water storage. Although these figures likely have significant uncertainty

and spatial variability, this water balance is shown to be plausible. It was also found that

permafrost limits surface and groundwater interaction, even if located in low-permeability

sediments. It is expected that permafrost degradation will likely increase stream baseflow,

especially in winter.

RÉSUMÉ

La distribution et la dynamique des eaux souterraines ont été étudiées dans un petit bassin

versant situé dans la zone de pergélisol discontinu près d’Umiujaq au Nunavik (Québec,

Canada), afin d'effectuer une analyse quantitative du cycle de l'eau dans un bassin versant

subarctique. En raison de la complexité du contexte géologique dans le bassin versant, une

approche intégrée a été utilisée afin de décrire le contexte hydrogéologique nécessaire à la

réalisation du bilan hydrique. Sur la base de ce bilan hydrique, pour les deux années

hydrologiques étudiées en 2015 et 2016, les valeurs moyennes sont de 828 mm pour les

précipitations, de 337 mm pour l’évapotranspiration, de 46 mm pour la sublimation dans la

neige, de 263 mm pour le ruissellement et de 183 mm pour l’échange entre les eaux

souterraines avec d’autres aquifères situés en dehors du bassin versant et 0 mm pour les

variations d’emmagasinement. Même s'il existe une incertitude significative et une variabilité

spatiale de ces valeurs, ce bilan hydrique s'avère plausible. Il a également été constaté que

le pergélisol limite les interactions entre les eaux de surface et les eaux souterraines même

s’il se retrouve uniquement dans des sédiments peu perméables. On s'attend à ce que la

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dégradation du pergélisol ait pour conséquence l’augmentation du débit de base des cours

d'eau, en particulier l’hiver.

Keywords: cold regions, water supply, water balance, permafrost, Nunavik.

1 INTRODUCTION

Temperatures in Arctic regions are rising faster than the global average (AMAP, 2017) while

the Arctic climate is also becoming wetter and increasingly variable. Compared to current

conditions and under a high emission scenario, it is predicted that in 2050 the Arctic snow

cover will have decreased by 10-20% and that near-surface permafrost will have decreased

by as much as 35% (AMAP, 2017). These changes, which are already occurring at high

latitudes, are having significant impacts on northern communities, water resources, and

ecosystems (AMAP, 2017). One foreseen consequence of permafrost degradation is to

transform the hydrological cycle from a surface-water to a groundwater-dominated system,

which would significantly modify the terrestrial portion of the hydrologic cycle. For instance,

an increase in fresh-water discharge was observed in many Arctic and subarctic rivers during

low-flow conditions, which is attributed to a reactivation of groundwater flow systems caused

by permafrost degradation (e.g. Walvoord and Striegl 2007; Bense et al., 2009; Duan et al.,

2017; Lamontagne-Hallé et al., 2018).

The shift toward groundwater-dominated hydrological systems will likely increase the

availability of groundwater as a source of drinking water. Currently, water supplies for

northern Canadian communities come mainly from surface water such as rivers and lakes.

While these sources of water are usually abundant, their quality is variable, they are

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vulnerable to contamination, and are unreliable as they often either freeze or dry up in winter

(Lemieux et al, 2016). In contrast, groundwater is generally of better quality, is less

vulnerable to contamination, and usually requires only minimal treatment. However, being

currently stored as ground ice in permafrost, access to groundwater is limited for most

northern communities. Due to permafrost degradation, the increased availability of

groundwater in the context of climate change and its exploitation as a source of drinking

water could potentially improve water quality and security of water supply for northern

communities.

Among the fourteen Inuit communities in Nunavik, which is located in the northern portion of

the Province of Québec (Canada) (Fig 1), groundwater is used as a source of drinking water

in only two, including one located in the continuous permafrost zone (Lemieux et al., 2016).

At Salluit, for example, the water supply for the local community is provided from an aquifer in

a closed talik below the Kuuguluk River (Lemieux et al., 2016). In the discontinuous

permafrost zone in Nunavik, groundwater can also be available in unconsolidated deposits

partially affected by permafrost, and may be exploited, provided that the hydrogeological

context is delineated.

Groundwater, as baseflow, also sustains many surface water bodies such as lakes and rivers

that many communities depend on as a source of drinking water. According to Lamontagne-

Hallé et al. (2018), groundwater discharge is expected to increase in winter due to climate

warming. Groundwater also plays an important role on permafrost dynamics due to advective

heat transport (e.g. Rowland et. al., 2011; Mackenzie and Voss, 2013; Dagenais et al., this

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issue). However, there is a generally poor understanding of groundwater distribution and

dynamics in cold regions along with its interaction with permafrost and surface water. For

instance, quantitative understanding is lacking on how climate change in the Arctic and

subarctic will affect the water cycle and, more importantly, how it will impact the quality and

availability of surface water and groundwater resources. Answering these questions is

essential for planning effective multi-use of water resources (AMAP, 2017).

Numerical models are useful to investigate and better understand the non-linear interactions

between climate, water, permafrost and vegetation. However, several model applications to

date have been restricted to parametric studies based on simplified conceptual models.

Simulation predictions can be significantly improved when models are applied to realistic

hydrogeological contexts and are based on measured site-specific hydraulic properties and

long-term series of climate, groundwater levels and river discharge. Unfortunately, there is a

paucity of hydrogeological data to support these models in Arctic and subarctic regions

(Ireson et al., 2013; Walvoord and Kurylyk, 2016).

In this paper, groundwater distribution and flow dynamics is studied in a small watershed

located in the discontinuous permafrost zone at Umiujaq in Nunavik (Québec), Canada. The

main objectives of this study are to provide a quantitative analysis of the water cycle in a

subarctic, cold-environment watershed and to capture its seasonal variations. This study

provides a baseline for determining the long-term impact of climate change on the water

cycle and creates a detailed dataset which can be used as input for numerical models. It also

contributes to the understanding of groundwater dynamics in cold environments to support

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long-term sustainable management of groundwater. Finally, the cryo-hydrogeological context

is described to assess the interaction between groundwater and permafrost dynamics within

this watershed (see Dagenais et al., this issue).

This study was conducted in an extensively instrumented watershed, which allowed

measuring the components of the water balance along with their seasonal variations. The

study area is first presented, followed by a detailed description of the investigation methods.

Observations of precipitation, air temperature, groundwater levels and river discharge are

then presented. The water balance in the studied watershed is computed from the flow of

water between the atmosphere, the land surface, the stream, the aquifers, and the

surrounding watersheds. Finally, a conceptual model of the watershed is proposed.

2 STUDY AREA

The study area is a small watershed located in the Tasiapik Valley near Umiujaq (Inuktitut:

ᐅᒥᐅᔭᖅ), a small Inuit community of about 440 inhabitants (Statistics Canada, 2018) located

on the eastern shore of Hudson Bay in Nunavik, Canada (56°33’ N, 76°31’ W; Fig. 1), within

the discontinuous permafrost zone. The watershed is bordered to the south-west by a cuesta

and to the north-east by the Umiujaq Hill, and is drained by a small stream that flows to the

south-east which discharges into Tasiujaq Lake (Fig. 1c and 2). The boundaries of the

watershed and its area (2.23 km2) were delineated in ArcGIS using a high-resolution digital

elevation model obtained from an airborne LIDAR survey. The watershed boundaries and

outlet, and the location of 9 piezometers (Pz1 to Pz9) and 3 thermistor strings (IMMATS1 to

IMMATS3), which are distributed over 7 sites, are shown in Fig. 2. The piezometers and

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thermistor strings are part of the Immatsiak monitoring network, designed and operated by

the Ministère du développement durable, de l’environnement et de la lutte contre les

changements climatiques du Québec (MDDELCC - Department of Sustainable Development,

Environment and Fight against Climate Change) to assess the impact of climate change on

groundwater resources in Nunavik (Fortier et al., 2013).

The climate in Umiujaq is subarctic with long winters and short summers. Over the period

from 2013 to 2017, the mean annual precipitation and temperature were 645 mm and -1.6°C,

respectively. Mean, minimum and maximum monthly air temperatures along with mean

monthly precipitation data are given in Fig. 3. Freezing temperatures occur between October

and May while above-zero temperatures occur from April to September. Snowfalls occur from

September to June, but are more intense from October to May. Most precipitation occurs

from July to January and snowfalls represent about 50% of total precipitation.

Streams in the watershed are mostly intermittent, except for the main stream, which is

perennial (Fig. 2). The watershed contains numerous small lakes, many of which are

thermokarst ponds (Beck et al., 2015). The mapping of groundwater discharge zones,

perennial and intermittent streams, as well as ponds, combined with the airborne LIDAR

survey, was used to draw the watershed drainage network shown in Fig. 2.

The watershed can be divided into three sections according to the topography. The upstream

and downstream sections are relatively flat compared to the central section, which is

characterized by steep slopes (Figs. 2a and 4a). The elevation difference between the

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upstream and downstream section of the valley is about 100 m. The cuesta on the

southwestern edge of the valley forms a 200 m high cliff that represents a significant

topographic feature in the area (Figs. 1, 2a, 2c and 4b).

The geological units in the watershed consist of unconsolidated Quaternary deposits

overlying bedrock. These sediments were deposited following the last deglaciation, which is

described by Fortier et al. (this issue). The surficial geology is shown in Fig. 2b while two

vertical cross-sections of the various geological units are found in Fig. 4. The bedrock

consists of two formations: an arenite of the Pachi Formation (Ri) and a brecciated basalt

belonging to the Persillon Formation (Rs). Frontal moraine deposits (unit GxT) overly the

bedrock, forming a 5 to 30 m thick layer. This moraine deposit is covered by a layer of

subaqueous fluvioglacial sediments (unit Gs), whose thickness varies between 5 and 20 m.

Deep marine sediments (unit Ma), composed mainly of silt, are located above the

fluvioglacial sediments. The extent of this marine silt unit is much smaller than the total extent

of the watershed (Fig. 5b) and its thickness can reach up to 30 m in the lower part of the

watershed. A sand layer consisting of littoral (unit Mb) and intertidal (unit Mi) sediments is

located above the marine silt unit. The extent of the combined Mb and Mi sand units is shown

in Figs 4a and 5c. Their thickness is variable but they are generally much thicker in the upper

part of the watershed (Fig. 5c).

Units Gs and Gxt form a lower aquifer that is unconfined in the upper part of the valley, and

confined in the lower part of the valley where the aquifer is overlain by the low-permeability

silt unit (unit Ma; Fig. 4a). Artesian conditions for this lower aquifer have been observed

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during the Fall in piezometer Pz4, suggesting upward groundwater flow across the silt unit

from the lower aquifer toward the surface. A thick unsaturated zone that can reach a

thickness of 30 m is located above the unconfined part of the lower aquifer (Fig. 4a). The

combined littoral (Mb) and intertidal (Mi) sand units form an upper unconfined aquifer above

the silt unit. The extent of the unconfined upper aquifer is shown in Figs. 4a and 5c and

corresponds to areas where the thickness of the Mb and Mi units is large enough to store

water and where it is underlain by the marine silt unit. Hydraulic conductivities obtained from

slug tests or grain size analyses using the Hazen (1892) formula are reported in Table 1 for

many of these geological units. While no values are given for the bedrock, field observations

suggest that it has a low hydraulic conductivity and can be considered as a leaky aquitard.

Along the cuesta, the rock is fractured and may be locally more permeable.

In the Tasiapik Valley, permafrost is discontinuous and only exists in the form of permafrost

mounds (Pmf) formed in the frost-susceptible marine silt unit (Ma). The permafrost extent is

shown in Figs 2b and 4. The dome shapes of the permafrost mounds are due to frost heave

and accumulation of segregation ice in marine silt sediments in contact with the cold air after

the last deglaciation (Fortier et al., this issue).

The tree line between the shrub tundra and forest tundra crosses the valley, resulting in a

very heterogeneous vegetation cover which varies from black spruce to lichens (Truchon-

Savard and Payette, 2012). Provencher-Nolet (2014) mapped the vegetation in the valley

using aerial photographs and field observations, and identified four vegetation types (Fig. 2c).

The first type is lichens, which are mainly found on topographic highs, permafrost mounds,

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and well-drained soils, and are most commonly found in the upstream portion of the

watershed. The second and most abundant vegetation type is shrubs that are found in

convex topographic areas and on permafrost mounds that show signs of advanced

degradation. A strong link exists between permafrost thaw, tundra shrubification and changes

in snow cover, as described by Pelletier et al (2008). The third vegetation type is spruce,

which is located in well-drained areas such as south-east of the watershed, near the rock

outcrop. The fourth and least abundant vegetation type is grass, which is located in small

depressions or near water bodies. There is no vegetation along the edges of the watershed,

where ground surface consists of rock outcrops or debris.

3 METHODOLOGY

The water balance equation for the watershed can be written as:

P−ET−S−Q−G=∆S (1)

where P is precipitation, G is groundwater exchange with aquifers outside the watershed

boundaries (positive for a loss and negative for a gain), ET is evapotranspiration, S is

sublimation, and Q is runoff. The term S represents the change in water stored in the

watershed. Changes in water stored in the unsaturated zone, on the land surface, and within

the stream are assumed negligible because the water balance is considered over a

hydrological year and because they would represent a small volume compared to the

aquifers. The total change in water stored S is therefore the sum of changes in the upper

aquifer Su and in the lower aquifer Sl, and is given by:

∆ S=∆ Sl+∆Su (2)

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All terms in Eqs 1 and 2 were determined directly on-site, except for sublimation which was

back-calculated by closing the budget in Equation 1.

Further insights into water partitioning within the watershed can be gained by expanding the

water balance for the land surface, the stream, and for the upper and lower aquifers. Based

on the hydrogeological context presented in the previous section (Fig. 4a), a flow chart of the

water balance is proposed for these associated components (Fig. 6). For the land surface,

the water balance is:

P−ET−S−R−Q s=0 (3)

where Qs is surface runoff and R is groundwater recharge, which is the sum of groundwater

recharge into the lower aquifer Rl and into the upper aquifer Ru:

R=R l+Ru. (4)

Two different values are used for recharge. The recharge of the upper aquifer was measured

in the field but the recharge to the lower aquifer was not evaluated. Instead, recharge to the

lower aquifer was obtained with the following water balance equation, assuming there is no

surface runoff over the recharge area of the lower aquifer and that there is no drainage from

the upper aquifer to the lower aquifer across the marine silt unit (this assumption will be

discussed later):

Rl=P−ET−S (5)

The water balance for the stream is obtained by partitioning total runoff (Q) into surface runoff

Qs and baseflow Qb using hydrograph separation:

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Q=Q b+Qs (6)

where Qb can be split into baseflow from the upper aquifer Qbu and baseflow from the lower

aquifer Qbl associated with upward flow across the marine silt unit:

Qb=Qbu+Qbl. (7)

Baseflow is partitioned because the stream flows over both the upper aquifer and the marine

silt unit (Fig. 4a).

The water balance for the upper aquifer is given by:

Ru−Qbu=∆ Su. (8)

It is assumed that there is no water exchange from the upper aquifer toward the lower

aquifer. This hypothesis is motivated by the presence of a thick unsaturated zone that exists

below the silt layer, under the upper aquifer (Fig. 2), which presumably acts as a capillary

barrier and prevents downward vertical drainage of the upper aquifer.

Finally, assuming that the groundwater drainage basin for the lower aquifer corresponds to

the surface watershed, the water budget for the lower aquifer is:

Rl−Qbl−G=∆S l. (9)

All terms in Eqs 3 to 9 were directly evaluated except for the total recharge R, recharge of the

lower aquifer Rl, baseflow from the upper aquifer Qbu and baseflow from the lower aquifer Qbl

which were obtained by closing the equations. The methods used to evaluate each of the

components of the water balance are given below. As a summary, precipitation P is

measured with a rain/snow gauge, evapotranspiration ET is obtained using empirical

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relationships, runoff Q is measured using a flume, groundwater recharge of the upper aquifer

Ru is obtained using the water table fluctuation method, groundwater exchange G is obtained

using the Darcy equation, while surface runoff Qs and baseflow Qb are obtained using

hydrograph separation. Since there is always an unknown in each water balance equation,

no error terms could be evaluated to assess the validity of the component estimates in the

water balance.

The available data needed to evaluate the water balance vary in terms of temporal coverage.

For the Umiujaq site, all required data are available from July 1st, 2014 to June 30th, 2017,

which corresponds to three years of data. Furthermore, water balance is usually evaluated

over a hydrological year (HY), which in Quebec and North America begins October 1 st and

ends on September 30th. For the sake of consistency and comparison with published data,

the same period was chosen for this study. Therefore, in the Tasiapik Valley at Umiujaq, full

time series data are available for two HYs: from October 1st 2014 to September 30th 2015 (HY

2015), and from October 1st 2015 to September 30th 2016 (HY 2016).

3.1 Precipitation

Precipitation was measured by a Geonor T-200B all-weather precipitation gauge , being part

of the SILA network of the Centre d’études nordiques (CEN – Centre for Northern Studies) of

Université Laval (see the inset in Fig. 5a). The gauge consists of a reservoir for collecting and

storing precipitation, which is weighed using vibrating wire sensors. The gauge contains

antifreeze to transform snow into water and an oil layer is maintained at the surface of the

stored precipitation to avoid evaporation. The rain gauge is also equipped with a windshield

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to reduce air turbulence. Although the gauge only requires minimal maintenance, the raw

data must be corrected for air temperature and wind speed (Forland et al., 1996, Morin et al.,

2012), which are also measured on site by the SILA network.

In cold regions, biases in the precipitation data are mostly due to underestimation of snowfall

(Smith, 2007). Since snow has a low density and low falling speed, turbulent airflow over the

gauge can prevent snow from falling into the bucket, which results in underestimating the

true precipitation. Corrections due to gauge maintenance and evaporation losses are first

made using the method proposed by Pan et al. (2016). Using the method described by

Smith (2007), the precipitation data were then processed to account for the effect of wind

which tends to underestimate snowfall

3.2 Evapotranspiration

Measuring actual evapotranspiration ET in the field is a challenging task. Therefore, empirical

methods relying on weather data were used instead. First, potential evapotranspiration PET

was obtained using an empirical relationship developed by Bisson and Roberge (1983),

which requires only the minimum and maximum daily air temperatures, Tm [°C] and TM [°C],

respectively:

PET=0.029718 (T M−Tm )0.019[ 95 ((TM−Tm )+64 )] (12)

The temperature data used for this study were obtained from the meteorological station

IMMATS3 (Fig. 4a; Fortier, 2017). Dionne et al. (2008) demonstrated the applicability of

Equation 12 for the southern part of the province of Quebec. This empirical relationship was

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tested for 12 meteorological stations, including one at La Grande (53°37’ N, 77°42’ W), which

is reasonably close to Umijuaq (56°33’ N, 76°31’ W). However, rather than PET, the actual

evapotranspiration ET is the parameter required in the water balance. Fortunately, several

empirical relationships have also been developed to convert potential evapotranspiration to

actual evapotranspiration. Huet et al. (2016), for example, used the relation proposed by

Budyko (1974) for catchments located in the province of Québec:

ET=[P×(1−e(−PETP )×PET × tan h( PPET ))]

0.5

(13)

where P is the annual average precipitation and PET is obtained from Eq. 12. Because of its

simplicity and local usage, this relationship was selected for this study.

3.3 Runoff (discharge)

In order to monitor surface water runoff, an H flume (open channel flow) was installed in the

main stream near the outlet of the watershed, during the 2013 field season (Fig. 7). The

flume is composed of two parts: an approach section, which ensures laminar flow, and a

control section, which measures the flow as a function of the water level in the flume. The

water level is measured by water and atmospheric pressure sensors (Solinst Levelogger

Edge 2 m) installed in a stilling well located within the frame of the control section.

The water level in the flume h is converted into a flow rate Q using the following calibration

provided by the flume manufacturer:

Q=0.0223−0.5550h0.5+125.5276h1.5+939.5717h2.5 (14)

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The channel can measure a flow rate between 0.04 l/s and 311 l/s. Flow is considered

negligible for Q < 0.04 l/s, which essentially represents a dry stream. In contrast, if Q > 311

l/s which indicates a stream overflow, the rate cannot be assessed. The volume of water

flowing in the channel is assessed by integrating the flow rate over time. This volume is then

divided by the area of the watershed (2.23 km2) to obtain the equivalent water height.

3.4 Groundwater recharge

Groundwater recharge is assessed for the upper aquifer using the water table fluctuation

(WTF) method (Healy and Cook, 2002). This method was selected for its simplicity and

availability of groundwater level records. Groundwater recharge is obtained from the product

of the specific yield Sy and the effective water table rise ∆h* (Healy and Cook, 2002; Crosbie

et al., 2005):

R=S y×∆h¿ (15)

The WTF method is applied using the event-based approach where groundwater recharge

events are identified according to the delineation of a master recession curve (Nimmo et al.,

2015). This event-based WTF method requires the delineation of recession periods that can

be clearly identified on the observed hydrograph. Hence, the recession periods are manually

extracted and used to adjust a linear master recession curve used thereafter to automatically

identify groundwater recharge events. Corrections are applied to groundwater levels to

account for regional drainage and overshoot effects.

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The specific yield is a significant source of uncertainty in the estimation of groundwater

recharge with the WTF method (Healy and Cook, 2002; Crosbie et al., 2005). The estimation

of specific yield requires consistency with the applied WTF method for improving the

reliability of the recharge estimates since the soil may not be at residual saturation when a

precipitation event occurs (Delottier et al., 2018). As proposed by Sophocleous (1991),

instead of using specific yield values obtained from hydraulic tests or laboratory

measurements, fillable porosity values obtained from in situ volumetric water content (VWC)

probes are used. The fillable porosity is obtained as the area between water content profiles

before and after a groundwater recharge event (Sophocleous, 1991). A fillable porosity is

thus estimated and multiplied by the associated effective water table rise for each episode.

For this study, close to the meteorological station of the SILA network (see the inset in Fig

5a), four of the five piezometers (PzDL, PzDA, PzDP and PzDH) located in the upper

unconfined aquifer were collocated with a string of VWC probes (5TM, Decagon instruments)

for which hourly measurements are available for HY 2015 and 2016. However, only one of

these sites (PzDL) has VWC probes that span the entire unsaturated zone, which is needed

to estimate the fillable porosity. Therefore, groundwater recharge using the WTF method is

only evaluated for this piezometer. The VWC string is composed of 20 probes buried to

depths ranging from 10 cm to 4.0 m, which is slightly above the water table during the winter.

Since the upper unconfined aquifer does not extend over the entire watershed, the recharge

value has to be multiplied by the fraction of the watershed occupied by the unconfined

aquifer. It was determined that groundwater recharge only occurs where the thickness of the

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upper aquifer is greater than 5 m, which corresponds to 25 % of the total area of the

watershed (hatched area in Fig. 5c). Elsewhere, precipitation either (i) does not infiltrate and

instead flows over the marine silt unit as surface water runoff, (ii) infiltrates as hypodermic

flow within the thin layer of intertidal and littoral sediments (Mi and Mb) located above the

marine silt unit, or (iii) infiltrates to recharge the lower aquifer according to Eq. 5.

3.5 Groundwater storage

Groundwater storage variations within the upper aquifer ∆Su can be estimated using

observed water levels changes in piezometers PzDA, PzDH, PzDL, PzDP, and Pz2. The

storage variation in groundwater reserves is obtained by multiplying the hydraulic head

variation over a hydrological year by the fillable porosity. The fillable porosity used here is the

average value of 0.12 for all the fillable porosity values for each effective recharge event in

Table 2. Moreover, the storage variation obtained above must be multiplied by 25%, which is

the fractional extent of the unconfined upper aquifer.

The groundwater storage variation for the lower aquifer ∆Sd could be obtained using the

same methodology. However, the lower aquifer can be confined (elastic storage) or

unconfined (phreatic storage) and the storage calculation is different for each case. Since

there is great uncertainty in the extent of these two zones, this water budget component is

instead obtained upon closure of Eq. 9.

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3.6 Groundwater exchange

As mentioned earlier in this section, it is assumed that the drainage area of the lower aquifer

corresponds to the watershed surface and groundwater exchange with other sub-watersheds

can only occur at the outlet of the watershed as groundwater discharge. Groundwater

discharge out of the watershed from the lower aquifer (G) is calculated by multiplying the

Darcy flux in the lower aquifer by the cross-sectional area of the aquifer. Ideally, the Darcy

flux at the watershed outlet should be used. However, the best estimates for the Darcy flux

were obtained at piezometers Pz4, Pz6, and Pz9 by Jamin et al. (this issue) who conducted a

tracer experiment using the finite volume point dilution method (Fig. 2b). While this approach

cannot delineate groundwater flow directions, it is assumed that flow is generally parallel to

the valley axis and to cross-section A-A’ (Fig. 4). This is supported by the fact that the natural

discharge area for groundwater from the Tasiapik Valley is Tasiujaq Lake, which is located

down the valley along its main axis (Fig. 1).

The value of Darcy flux of 0.58 m/d measured in the piezometer Pz4 (Jamin et al., this issue)

is used since it is closest to the watershed outlet. The cross-sectional area of the lower

aquifer at the watershed outlet (2017 m2), perpendicular to the valley axis (parallel to cross-

section B-B’ – Fig.5), was obtained from the 3D geological model (Fortier et al., this issue).

3.7 Surface runoff and baseflow

Surface runoff Qs and baseflow Qb are evaluated using the hydrograph separation technique.

The web application "WHAT: Web-based Hydrograph Analysis Tool" (Lim et al., 2005)

proposes three methods that are used here. The first method is called "Local minimum" and

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was developed by Lim et al. (2005). This technique simply connects the hydrograph

minimums and does not account for the duration of flow; it generally overestimates the

baseflow.

The second method is called "One parameter digital filter" (Lyne and Hollick, 1979) and is

based on the following equation:

(16)

where qk is surface runoff to a time index k, qk-1 is a surface runoff to a time index k-1, yk is

the water flow at time k, yk-1 is the flow of the water at time k-1 and a is a filter parameter

which is assessed from the recession curve using the following equation (Meyboom, 1961):

(17)

where Q0 is the stream flow at the beginning of the recession and Qt is the flow rate at time t.

This method does not represent any physical phenomenon, but eliminates the subjectivity

caused by manual separation (Lim et al., 2005).

The last method is the "recursive digital filter" (Eckhardt, 2005) and is based on the following

equation:

(18)

where bk is the baseflow at a time k, bk-1 is the baseflow at a time k-1, yk is the water flow at

time step k, BFImax is the index of the baseflow and a is a filter parameter. The index BFImax is

defined as the ratio of the base rate to the flow of water. Eckhardt (2005) suggests taking a

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value of BFImax =0.80 when the stream flow is constant and when the aquifer is composed of

granular materials, as in this study. The parameter a is estimated as explained above. These

three methods were applied using daily discharge values.

4 RESULTS

The corrected precipitation values as a function of time are shown in Fig. 8a. The total

amount of precipitation for HY 2015 and 2016 is 850 and 805 mm, respectively (Table 3). It

can be noted that no precipitation was recorded from September 24 to October 26, 2015,

which was due to instrument failure. In order to fill this gap with realistic values of

precipitation, the average proportion of precipitation for this period over the total precipitation

was calculated for years 2013, 2014 and 2016. The measured precipitation for HY 2015 and

2016 were then corrected by adding the proportion of missing precipitation based on this

average value. The total precipitation given above includes this correction.

Using this precipitation along with the air temperature values, PET and ET were then

evaluated as explained above. For the HYs 2015 and 2016, PET is 396 and 366 mm

respectively, while ET is 349 and 324 mm, respectively (Table 3).

Stream discharge is shown in Fig. 8b, which shows that only for a few periods, the water

level was above the flume (the y-axis limit in Fig. 8b is set to the overflow limit of the flume).

However, the same rating curve was used to convert the water levels into discharge values.

Since the section over which the water flows during these periods is wider than the flume, the

discharge values are underestimated for these periods. Although no data were collected for

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some periods in Winter 2016 and 2017, this does not mean that no water was flowing in the

flume but rather that water in the stilling well, and the data logger itself, were frozen. During

these periods, a linear interpolation was applied in order to calculate baseflow volumes. The

measured air temperature at the meteorological station IMMATS3 is also shown in Fig. 8b

which was used to identify (in shaded areas) the periods where the temperature was below

the freezing point. For these periods, the stream discharge linearly decreases, while for the

periods where the temperature is above the freezing point (spring/summer/fall), large

variations in stream discharge are observed due to precipitation events. The total stream

discharges for HYs 2015 and 2016 are 268 and 258 mm, respectively (Table 3).

Hydraulic heads for the piezometers in the upper and lower aquifer are shown in Fig. 8c and

8d, respectively. Seasonal variations observed for piezometers in the lower aquifer (28 to 42

m) are much greater in magnitude than for the upper aquifer (127 to 129.8 m). A phase shift

in seasonal cycles is also perceptible for the piezometers in the two aquifers. Moreover,

changes in hydraulic heads are much smoother for the lower aquifer than for the unconfined

aquifer. However, in piezometer Pz4, the changes are not as smooth as in piezometers Pz6

and Pz9 and the heads exhibit sharp low-amplitude variations. For piezometer Pz4 in the

lower confined aquifer, the hydraulic head is constant from October 2015 to mid-January

2016 and then exhibits a slight increase before the winter recession. The relatively uniform

heads over this period are associated with artesian conditions, and correspond to the

elevation of the piezometer casing. In January, the water froze in the piezometer casing,

which allowed the hydraulic head to increase above the piezometer casing elevation since

the water was capped below the ice-plug.

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Groundwater recharge events observed in piezometer PzDL are shown in Fig. 9, along with

their associated recession curves. The greatest recharge events take place during snowmelt,

which starts around May. Other recharge events related to precipitation events occur during

the summer until November/December, when winter starts.

The VWC profiles before and after recharge events 7 and 8, along with the associated water

table in piezometer PzDL, are provided in Fig. 10. These two specific events are shown since

they exhibit two contrasting but characteristic behaviors during recharge. Recharge event 7

corresponds to a relatively small water level increase that occurs in late fall and early winter

for which the refillable porosity is small since the VWC is already high before the recharge

event starts. Recharge event 8 is an important recharge event that corresponds to the

snowmelt period after winter. At the end of the winter season, corresponding to the beginning

of the recharge event, the water content did not reach its residual values at depths of 3.8 and

4 m. For this reason, the fillable porosity is much lower than the specific yield. The water

table rise, associated fillable porosity, and recharge values are given in Table 2 for each

recharge episode. For a few events, the fillable porosity is zero since the water table rise is

due to a marginal amount of groundwater recharge. Total groundwater recharge values for

HYs 2015 and 2016 are 580 and 472 mm, respectively. However, as mentioned previously,

since the upper aquifer is not continuous over the sub-watershed, these values were

multiplied by 25 %. The equivalent recharge for the sub-watershed area is then 142 and 116

mm for HYs 2015 and 2016, respectively (Table 3).

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Groundwater storage variations in the upper aquifer for HYs 2015 and 2016 for each

piezometer are presented in Table 4. For HY 2015, the storage increases from 2 mm to 10

mm with an average of 6 mm, while for HY 2016 the storage decreases from 3 mm up to 8

mm, with an average of 6 mm. Thus, the mean groundwater storage increase in HY 2015 is

equal to the decrease in HY 2016, which means that the water level at the end of HY 2016

was the same as the water level at the beginning of HY 2015. The values used in the water

balance are the average values of 6 and -6 mm for HY 2015 and 2016, respectively (Table

3).

As shown in Table 4, the results of the three different hydrograph separation methods used

for HYs 2015 and 2016 are very similar. The average baseflows are 153 and 161 mm for

HYs 2015 and 2016 (Table 3), which represents a baseflow index of 57 % and 62.5 %,

respectively.

Groundwater flow out of the basin from the lower aquifer was assessed from a tracer

experiment conducted in July 2016 (Jamin et al., this issue; 190 mm). The same term could

also be found by closing the balance for the average of the two HYs, since the average

groundwater storage is zero (Equation 9). Applying this approach, the groundwater budget

term for flow out of the basin is 183 mm, which is very close to the value obtained with the

tracer dilution test (190 mm). The value of 183 mm is then used as the groundwater

exchange term G in order to maintain consistency in the water balance (Table 3). This value

was also applied to HY 2015 and 2016 since changes in groundwater storage were not

available for these two years. Changes in groundwater storage Sl in the lower aquifer for HY

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2015 and 2016 are assessed by closing Eq. 9, giving values of 31 and -31 mm for HY 2015

and 2016, respectively (Table 3).

The values of the components of the surface and subsurface water balance for the

watershed in the Tasiapik Valley at Umiujaq as estimated above are summarized in Table 3

for HYs 2015 and 2016. An average value for these two years is also given. As mentioned in

the Methodology section, the values for sublimation S, total groundwater recharge R,

recharge of the lower aquifer Rl, baseflow from the upper aquifer Qbu and baseflow from the

lower aquifer Qbl are obtained by closing Eqs 1 to 9.

5 DISCUSSION

The discussion is divided into three sections. First, the observations are interpreted for

explaining the water dynamics in the watershed. The water balance components are then

discussed. The discussion concludes with the presentation of a conceptual model of the

watershed in the Tasiapik Valley at Umiujaq.

5.1 Water flow dynamics

Based on the monitoring of hydraulic heads for the piezometers in the upper and lower

aquifers (Fig. 8c and 8d, respectively), the seasonal variations in the lower aquifer are much

greater than those in the upper aquifer. Since the piezometers in the lower aquifer are

located mostly in the lower part of the watershed, which acts as a funnel, this lower aquifer is

collecting water from a large area, whereas the piezometers in the upper aquifer are in a

recharge zone and only respond to vertical recharge. Because the lower aquifer at

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piezometers Pz6 and Pz4 is confined, the water level fluctuations are much greater than for

the unconfined part, for a similar amount of groundwater recharge.

A phase shift in seasonal cycles is also apparent for the piezometers in the upper and lower

aquifers. For the piezometers in the upper aquifer, groundwater recharge starts immediately

when the air temperature is above the freezing point (Fig. 8c) since the water comes from the

melting snow immediately above the aquifer. For the lower aquifer, the hydraulic head only

starts to increase several weeks after snowmelt has started (Fig. 8d). The lower aquifer is

either confined by the marine silt unit or overlain by a thick unsaturated zone, which is also

overlain by the low permeability marine unit. Groundwater recharge for this aquifer therefore

occurs in areas without the marine silt unit above (see Fig. 4). Since the piezometers in the

lower aquifer are located much further down-gradient from these recharge zones, there is a

delay between the time when snow-melt recharges the upgradient lower aquifer and the time

when hydraulic heads begin to increase in the downgradient piezometers.

Moreover, the variations in hydraulic heads are much smoother for the lower aquifer than for

the upper aquifer (Figs 8c et 8d). Since the piezometers in the lower aquifer are far from the

recharge zones, the high-frequency component of the recharge signal is lost along the flow

path. However, hydraulic heads in piezometer Pz4 within the lower aquifer show high

frequency and low amplitude variations. This may indicate that the lower aquifer in the sector

of piezometer Pz4 is hydraulically connected to a recharge zone at the base of the nearby

cuesta since the frontal moraine deposits (unit GxT) nearly outcrops here and is in contact

with the littoral and pre-littoral sand unit (unit Mb) (Fig. 8b). In contrast, piezometer Pz6,

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which is located in a similar hydrogeological context, does not show these high frequency

variations. A discontinuous outcrop area for the lower aquifer along the bottom of the cuesta

could explain this observation. The hydraulic link between the lower aquifer and the cuesta

base is unclear.

The stream discharge during summer low flows that occurr between precipitation events is

lower than stream discharge during winter (Fig. 8b). However, hydraulic heads in the upper

aquifer are at a maximum level during the summer and decline during the winter period. If the

stream is fed by baseflow from the upper aquifer, the opposite would be expected: higher

stream discharge during summer low-flow and lower winter stream discharge. In contrast, the

hydraulic heads in piezometers Pz4 and Pz6, located in the lower confined aquifer, peak

around December/January (3 to 5 months later than for the upper aquifer), and slowly

increase during the summer period, declining afterward. This could explain the unusual

seasonal behavior of the stream baseflow. During the summer, the stream is mostly fed by

surface drainage and hypodermic flow following precipitation events along with drainage from

the upper aquifer between precipitation events. Baseflow from the lower confined aquifer is

low since the heads in that aquifer are also low. During the summer, the heads in the lower

confined aquifer increase steadily, reaching a maximum around December/January.

Significant upward flow then sustains the stream baseflow, which becomes higher than in the

summer season. This hypothesis is further supported by the fact that the piezometer Pz4 was

under artesian conditions (free flowing), from October 2015 to February 2016.

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In order to better understand the link between the lower aquifer and the stream, a cross-

section was drawn along the stream onto which the hydraulic heads measured in the

piezometers were projected for two different periods on June 1 rst and December 1rst 2016

(Fig. 11). While the hydraulic heads directly below the stream may be different at these

locations than where they were measured, the vertical gradients nevertheless suggest that

the lower aquifer is under artesian conditions during the winter period and groundwater flows

upward toward the stream throughout most of the lower part of the valley (Fig. 11). In the

summer, only the portion downgradient from piezometer Pz4 would be under artesian

conditions. These observations are compatible with those of Cochand et al. (this issue), who

showed, from total dissolved solids (TDS) concentrations along the stream and other

hydrogeochemical signatures, that groundwater from above the confined part of the lower

aquifer provides a significant contribution to the stream flow.

5.2 Water balance

Because one component of the global water balance of the watershed and other sub-water

balance components could not be independently evaluated, and were instead assessed by

closing the water balance equations, there is no means to assess their validity. Moreover,

significant uncertainties exist in all components, for example due to gaps in the field data,

measurement reliability, and discrepancies between the underlying hypothesis of the

methods used and the field conditions. Discussing all sources of uncertainty of each method

used here would be almost endless and is beyond the scope of this paper. Instead, in the

following, the focus will be on the plausibility of the obtained values.

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Precipitation is the most important part of the water balance since it determines the amount

of water that will partition into the watershed. To put the Umiujaq precipitation values in

perspective, they are compared here to the closest available precipitation measurements,

which are in Kuujjuarapik, Inukjuak and Kuujjuaq (Fig. 1). Environment Canada reports

average precipitation values of 661 and 542 mm at Kuujjuarapik and Kuujjuaq, respectively,

for the reference period 1981-2010, and 460 mm at Inukjuak for the reference period 1971-

2000. The average precipitation at Umiujaq for 2013-2017 is 645 mm, which is very close to

values in Kuujjuarapik. Lower values were expected since there is a north-west trend of

decreasing precipitation in Nunavik (Québec), Canada (Statistics Canada, 2008). This could

be explained by the summer climate near Umiujaq which has a maritime character with

heavy rain, frequent mists and cyclonic conditions (Klock et al., 2000). Moreover, there is a

microclimate in the enclosure of Tasiujaq Lake, where the Tasiapik Valley is located,

because the high cuestas form a natural protective barrier against exposure to cold winds

from Hudson Bay (KRG, 2007).

For HYs 2015 and 2016, the precipitation values are 677 mm and 673 mm, with an average

of 675 mm. However, the corrected precipitation for wind bias is 828 mm for the same period,

which is 153 mm above the raw precipitation data. The correction therefore enhances the

precipitation data by 23 %, which is slightly higher than the reported range (5 – 20 %)

reported by Yang et al. (2005) for 4802 stations located in cold environments.

The values for evapotranspiration reported herein were obtained with empirical relationships.

In order to assess their credibility and put the results into perspective, they are compared

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with values from other basins located in high-latitude watersheds. To this end, Kane and

Yang (2004) compiled meteorological data from 39 watersheds located worldwide between

latitudes from 44°N to 80°N and ranging in surface area between 0.1 km2 and 432 km2 (Woo,

2012). The reported evapotranspiration values range from 35 to 537 mm with a general trend

of decreasing evapotranspiration with increasing latitude (Fig. 12a). The average value of

evapotranspiration in the Tasiapik Valley at Umiujaq is 337 mm, which fits well with the

general trend for these watersheds. This value also plots at the boundary between

watersheds with and without permafrost, which seems plausible since the valley contains

discontinuous permafrost. The general agreement between these values suggests that the

calculated evapotranspiration reported here is realistic.

Kane and Yang (2004) also compiled the runoff ratios (surface runoff over total precipitation)

for these watersheds. This runoff ratio shows an increasing trend with latitude (Fig. 12b),

which is partly due to a decrease in precipitation with latitude (Kane and Yang, 2004). Again,

the average value of 0.13 in the Tasiapik Valley at Umiujaq fits well within this general trend,

which also suggests that it is plausible. Although the runoff ratio plots on the lower end of the

expected values, this can be explained by the complex hydrogeological settings where most

of the precipitation that recharges the lower aquifer flows out of the basin as groundwater

exchange below the outlet of the stream, instead of returning to the stream as baseflow

(Table 3).

Sublimation is also an important component of the water cycle in cold environments, but it is

one of the most poorly known (Liston and Sturm, 2004). The average value obtained here is

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46 mm for the two HYs, which represents about 6% of the annual precipitation and about

11% of the winter precipitation. According to Liston and Sturm (2004), reported values for

sublimation usually vary between 10 and 50% of winter precipitation. The obtained value falls

within the lower part of this range. However, only an independent measurement of this

component would resolve the uncertainty on this value. To this end, an Eddy-Covariance flux

tower was recently installed in the Tasiapik Valley, which will help to resolve the

evapotranspiration and sublimation values for the watershed.

The groundwater recharge value corresponding to 41% of the precipitation (Table 3) is high

for a cold-region watershed with long winters. These high values can be explained with the

following reasoning. First, there is no drainage network, and therefore no surface runoff, for

about half the watershed in the areas not underlain by the marine silt layer (Fig. 2b).

Secondly, more than half the watershed is covered with permeable sediments that promote

infiltration (Fig. 2b). Finally, the greatest recharge event occurs in the spring during snowmelt.

5.3 Conceptual model

The above observations, along with quantification of the water balance components, have

been integrated to develop a conceptual flow model for the watershed in the Tasiapik Valley

at Umiujaq (Fig. 13). In this conceptual model, water flow within the valley is controlled by the

complex geological setting that was delineated in detail using subsurface investigation

techniques (Fortier et al., this issue).

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An apparent correlation between the extent of the silt layer and that of the drainage network

(Fig. 5b) suggests that runoff occurs only above the silt layer and that elsewhere, infiltration

of precipitation feeds the lower aquifer. Water flowing out of the basin at the stream outlet

measured at the gauging station mostly comes from a sub-watershed corresponding to the

extent of the silt layer, and which contains the upper unconfined aquifer, instead of the entire

watershed delineated from the digital elevation model. Above the marine silt unit, there is a

thin upper unconfined aquifer composed of sandy material while a lower aquifer made of

sand and gravel is located below the marine silt unit. The lower aquifer is unconfined in the

upper part of the valley and becomes confined, with seasonal artesian conditions, in the

lower part. Groundwater within the confined aquifer becomes focused downgradient and

flows laterally outside the watershed, toward Tasiujaq Lake. Groundwater also flows upward

from the lower aquifer toward the stream, especially during the winter season when the water

levels in the lower aquifer are highest. Active groundwater flow in the confined aquifer has an

important role on permafrost dynamics within the marine silt unit (Dagenais et al., this issue).

The upper aquifer is recharged from precipitation over its entire area, while the lower aquifer

is recharged mostly outside the extent of the marine silt unit – in the upper part of the valley

or at the base of the cuesta ridge. However, recharge is not well understood in this area

since the outcrop zone of the lower aquifer is not well delineated and may be discontinuous.

A thick unsaturated zone exists below the silt layer in the upper part of the valley (Fig. 4a),

which presumably acts as a capillary barrier preventing vertical drainage of the upper aquifer.

Because the upper aquifer has a limited extent and pinches out in the steep central part of

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the valley, groundwater discharges in ravines in the form of springs. These springs are

mostly ephemeral during the spring season and possibly during important precipitation

events.

While the Tasiapik Valley hosts discontinuous permafrost, its distribution is limited to the

marine silt unit; some discontinuous permafrost also likely exists within the bedrock, outside

of the valley. This discontinuous nature of permafrost could suggest that it may not have a

strong impact on groundwater dynamics and the water cycle. However, significant interaction

between the upper and lower aquifers due to vertical flow across the marine silt unit has been

observed where the permafrost is located. This interaction is especially active in winter when

river baseflow comes mostly from the lower aquifer. Usually, the hydraulic conductivity of

frozen ground is less than its unfrozen counterpart (Kane and Stein, 1983), therefore

permafrost within the marine silt unit should act as a barrier for groundwater flow. Since the

permafrost mounds are degrading due to climate warming, it is likely that groundwater flow

across the aquifers will increase in the future. This will in turn increase winter stream flow as

suggested by Bense et al. (2009) and Lamontagne-Hallé et al. (2018).

6 CONCLUSION

This study aimed at improving the understanding of the water cycle, and especially

groundwater flow dynamics, in subarctic environments. A conceptual model for groundwater

flow is proposed here for a small watershed within the Tasiapik Valley near the Inuit

community of Umiujaq in Nunavik (Québec), Canada. The watershed hosts two aquifers that

are separated by a discontinuous leaky aquitard of marine silt. This marine silt unit has a

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determinant role on the hydrology of the watershed since the drainage network develops only

above this low hydraulic conductivity layer. Groundwater recharge toward the lower aquifer

occurs exclusively outside the extent of this marine silt unit. However, lower in the valley,

groundwater flows upward from the lower aquifer toward the surface, sustaining stream

baseflow during the winter. Delayed recharge in the lower aquifer compared to the upper

aquifer was observed due to the presence of this silt layer.

Being relatively less permeable and located within a leaky aquitard, permafrost likely has an

impact on groundwater flow dynamics and on the water balance in the watershed. Its main

role is to limit the amount of upward groundwater flow from the lower aquifer toward the

stream, especially in winter when the upward flux is highest. Since these permafrost mounds

are predicted to thaw over the next few decades (Dagenais et al., this issue), an increase in

winter baseflow and streamflow can be expected.

Due to the complexity of the subsurface geology, a detailed understanding of the

hydrogeological context was instrumental to delineate the watershed water balance. The

water balance reveals that the average values for the two studied hydrological years in 2015

and 2016 are 828 mm for precipitation, 337 mm for evapotranspiration, 46 mm for

sublimation, 263 mm for runoff, 183 mm for groundwater exchange as losses with other

aquifers outside the watershed, and 0 mm for changes in water storage. While there is

significant uncertainty in these values, they are plausible. A more accurate water balance

could be obtained by independently quantifying the terms that were estimated by closing the

water balance equations. To this end, the Eddy covariance tower that was recently installed

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in the Tasiapik Valley at Umiujaq should help resolve the evapotranspiration and sublimation

components. Numerical simulations using an integrated surface-subsurface model should

also help to resolve the water balance.

ACKNOWLEDGEMENTS

This work was funded by the Ministère du développement durable, de l’environnement et de

la lutte contre les changements climatiques du Québec (MDDELCC - Department of

Sustainable Development, Environment and Fight against Climate Change), the Fonds de

recherche Nature et technologies du Québec (Quebec Research Fund – Nature and

Technology - Establishment of New Researchers Grant), the Northern Scientific Training

Program (NSTP) administered by Polar Knowledge Canada and the Natural Sciences and

Engineering Research Council of Canada (Strategic Project Grant and Discovery Grants).

The authors thank the Centre d’études nordiques (CEN - Centre for Northern Studies) at

Université Laval for their logistical support at the field site, Marie-Catherine Talbot-Poulin as a

research assistant for her help in the field and at Université Laval, Georg Lackner and Daniel

Nadeau for their advice with the precipitation data, and Barret Kuryluk who reviewed the

manuscript.

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TABLES

Table 1. Hydraulic conductivity of selected key geological units in the watershed in the

Tasiapik Valley at Umiujaq. Numbers in parenthesis are the number of samples for these

analyses.

Hydraulic conductivity (m/s)Geological unit Slug test Grain size

(Hazen)Littoral and pre-littoral sediments (sand) - Mb 6.4 × 10-4 (1) 5.9 × 10-4 (30)

Marine sediments (silt) - Ma - 3.7 × 10-5 (4)*

Subaqueous fluvioglacial sediments (sand and gravel) - Gs - 2.4 × 10-4 (1)

Frontal moraine deposits (gravel, pebble and stone) - GxT 4.4 × 10-5 (3) 3.4 × 10-4 (2)*

Bedrock (basalt) - Ri 2.2 × 10-7 (1) -*Likely not valid since the Hazen formula was designed for sand materials

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Table 2. Details of the recharge events in the watershed in the Tasiapik Valley at Umiujaq

obtained with the water table fluctuation method.

Episode Start End H (mm)

Fillableporosity (-)

Recharge (mm)

1 2014-10-01 2014-10-18 272 0.036 102 2015-04-24 2015-05-27 861 0.140 1203 2015-05-29 2015-08-11 1707 0.175 2994 2015-08-16 2015-09-08 477 0.167 795 2015-09-14 2015-10-07 727 0.136 996 2015-11-02 2015-11-10 77 0.000* 07 2015-11-16 2015-11-24 290 0.091 268 2016-05-11 2016-08-20 2174 0.187 4069 2016-08-25 2016-09-05 90 0.000* 010 2016-09-19 2016-09-30 299 0.044 13

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Table 3. Water balance for the hydrogeological years (HY) 2015 and 2016 in the watershed

in the Tasiapik Valley at Umiujaq. The average values for the two years are also given.

Water budget components (mm) HY 2015 HY 2016 AveragePrecipitation (P) 850 805 828Evapotranspiration (ET) 349 324 337Sublimation (S) 13 78 46Runoff (Q) 268 258 263 Surface runoff (Qs) 115 97 106 Base flow (Qb) 153 161 157 Upper aquifer (Qbu) 136 122 129 Lower aquifer (Qbd) 17 39 28Groundwater exchange (G) 183 183 183Change in water storage (S) 38 -37 0 Change in GW storage in upper aquifer (Su) 6 -6 0 Change in GW storage in lower aquifer (Sd) 31 -31 0Recharge (R) 373 306 339 Upper aquifer recharge (Ru) 142 116 129 Lower aquifer recharge (Rd) 231 190 211

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Table 4. Change in groundwater storage in the unconfined upper aquifer in the watershed in

the Tasiapik Valley at Umiujaq. See Fig. 2 and the inset in Fig. 5a for the location of

piezometers.

PzDA PzDH PzDL PzDP Pz2

 HY

2015HY

2016  HY

2015HY

2016  HY

2015HY

2016  HY

2015HY

2016  HY

2015HY

2016

H1 (Oct. 1rst) (m) 127.77 128.29 127.92 128.49 129.14 129.48 126.51 126.68 127.61 128.34

H2 (Sept. 30) (m) 128.28 127.80 128.48 128.00 129.47 128.98 126.68 126.47 128.34 127.74

H (m) 0.51 -0.49 0.56 -0.49 0.33 -0.50 0.16 -0.21 0.72 -0,60

Su (mm) 7 -7   8 -7   5 -7   2 -3   10 -8

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Table 5. Base flow, direct runoff and base flow index in the watershed in the Tasiapik Valley

at Umiujaq for the three hydrograph separation methods used in this study (local minimum,

one parameter digital filter and recursive digital filter). Total flow and average values are also

given.

HY 2015 HY 2016Total flow mm 268 258Local minimum Base Flow mm 152 164 Direct runoff mm 116 93 Base flow index % 57 64One parameter digital filter Base Flow mm 154 161 Direct runoff mm 114 97 Base flow index % 57 62Recursive digital filter Base Flow mm 153 158 Direct runoff mm 116 100 Base flow index % 57 61Average Base Flow mm 153 161 Direct runoff mm 115 97 Base flow index % 57 62.5

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FIGURES

Figure 1. Location map. a) Location of the Nunavik region in Canada. b) Permafrost zonation

in Nunavik (after Allard and Lemay 2012) and location of the Umiujaq village. c) Location of

the studied watershed near Umiujaq.

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Figure 2. Environmental settings of the studied watershed in the Tasiapik Valley at Umiujaq.

a) Digital elevation model obtained from an airborne LIDAR survey and surface drainage

network. The main stream is orientated south-east towards the inland Tasiujaq Lake (Fig. 1).

b) Surficial deposits (A - alluvial sediments, B - glacial boulder fields, Ce - colluvial talus

scree deposits, GxT - frontal moraine deposits, Ma - fine deep water marine sediments, Mb -

littoral and prelittoral sediments, Md - deltaic and pro-deltaic sediments, O - organic deposits,

R - bedrock) and location of the Immatsiak groundwater monitoring network. Cross-sections

A-A’ and B-B’ are shown in Fig. 4 while cross-section C-C’ is shown in Fig. 11. c) Types of

land cover (Provencher-Nolet et al., 2014) overlying an Ikonos satellite image.

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Figure 3. Air temperature and precipitation at Umiujaq in Nunavik (Québec), Canada. Mean

monthly snow and rain precipitation values for the period from October 2012 to October 2016

are identified by vertical bars while the curves show the mean, minimum and maximum

monthly air temperatures for the same period.

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Figure 4. Cross-sectional cryo-hydrogeological settings in the Tasiapik Valley at Umiujaq; a)

A-A’ along the valley and b) B-B’ transverse to the valley of the. See Fig. 2 for the location of

cross-sections. Note that the vertical exaggeration is different in each figure (1:5 in Fig. 4a

and 2:3 in Fig. 4b). Aqf.: aquifer, Aqt.: aquitard.

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Figure 5. a) Map of the watershed in the Tasiapik Valley at Umiujaq showing the drainage

network, location of piezometers, weather station and gauging station. See the text for the

description of the inset. Thickness of the b) marine sediments (Ma), c) littoral, pre-littoral (Mb)

and intertidal (Mi) sediments, and d) subaqueous fluvioglacial sediments (Gs) and moraine

deposits (GxT) (Fortier et al., this issue). The hatched area in Fig. 5c is the portion of the sub-

watershed used to compute groundwater recharge.

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Figure 6. Flow chart of the water balance showing the water components of the global, land

surface, stream, upper and lower aquifer. See the text for the definition of components.

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Figure 7. Photograph of the H-flume gauging station.

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Figure 8. Meteorological, hydrological and hydrogeological data measured in the Tasiapik

Valley at Umiujaq: a) daily and cumulative precipitation, b) stream discharge and air

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temperature, c) hydraulic heads for the upper aquifer, and d) hydraulic heads for the lower

aquifer. See Figs 2b, 4, and 5a for the location of piezometers. The shaded areas correspond

to periods where the temperature is below 0°C.

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Figure 9. Variations of hydraulic heads monitored in piezometer PzDL for HYs 2015 and

2016. Recharge events numbered from 1 to 10 are identified with thick black lines while

recession curves are shown in red. See the inset in Fig. 5a for the piezometer location.

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Figure 10. Volumetric water content profiles and water table elevation before and after

recharge events 7 and 8. The hatched area corresponds to the fillable porosity associated

with the recharge event.

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Figure 11. Cross-section along the stream showing the projected piezometers and water

levels in the watershed in the Tasiapik Valley at Umiujaq. The values shown in parenthesis

are the transverse distances of the piezometer from the cross-section.

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Figure 12. (a) Mean annual evapotranspiration and (b) mean annual runoff for 39

watersheds located between latitudes 44°N and 80°N and ranging between 0.1 km 2 and 432

km2. Plotted data are from Kane and Yang (2004) and Woo (2012).

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Figure 13. Conceptual model of the watershed in the Tasiapik Valley at Umiujaq.

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