It is well established that the Earth's climate system has warmed
significantly over the past several decades, and in association there have
been widespread changes in various other Earth system components. This has
been especially prevalent in the cold regions of the northern mid- to
high latitudes. Examples of these changes can be found within the western
and northern interior of Canada, a region that exemplifies the scientific
and societal issues faced in many other similar parts of the world, and
where impacts have global-scale consequences. This region has been the
geographic focus of a large amount of previous research on changing
climatic, cryospheric, and hydrological regimes in recent decades, while
current initiatives such as the Changing Cold Regions Network (CCRN)
introduced in this review seek to further develop the understanding and
diagnosis of this change and hence improve the capacity to predict future
change. This paper provides a comprehensive review of the observed changes
in various Earth system components and a concise and up-to-date regional
picture of some of the temporal trends over the interior of western Canada
since the mid- or late 20th century. The focus is on air temperature,
precipitation, seasonal snow cover, mountain glaciers, permafrost,
freshwater ice cover, and river discharge. Important long-term observational
networks and data sets are described, and qualitative linkages among the
changing components are highlighted. Increases in air temperature are the
most notable changes within the domain, rising on average 2
Recent warming of the Earth's climate system has been impacting many
biogeophysical systems and their interactions globally (IPCC, 2013). Changes
have been particularly great in the northern high latitudes, where
observations show shifts in the amount and phase of precipitation,
diminishing seasonal snow cover, retreat and loss of glaciers, warming and
thawing of permafrost, earlier breakup of seasonal freshwater ice cover,
changes in the timing and magnitude of river discharge, and altered
composition, structure, and density of terrestrial vegetation communities
(Serreze et al., 2000; ACIA, 2004; Hinzman et al., 2005; White et al., 2007;
Prowse et al., 2009b; Callaghan et al., 2011a; Derksen et al., 2012; AMAP,
2012; Bush et al., 2014). Responses to climatic and other environmental
changes may be incremental or alternately characterized by threshold-type
behaviour, often involving complex feedbacks, and there is increased
sensitivity to warming in areas with winter temperatures near freezing
associated with the phase change of water at 0
The interior of western Canada provides an immediate example of cold region environmental changes observed globally and the societal issues faced in the context of such changes. Changes, including those listed above, have been pervasive, while the costs associated with recent hydro-climatic extreme events (e.g., floods, drought, and wildfire) have been increasing (e.g., Hanesiak et al., 2011; Pomeroy et al., 2015). The principal continental-scale drainages, the Mackenzie and Saskatchewan River systems, support a major area of Canada's food and energy production, mining, forestry, critical riverine and delta ecosystems, growing cities, rural and aboriginal communities, and freshwater supply to the Arctic Ocean and Hudson Bay. The region is highly vulnerable to climate change with pressures from natural resource and hydroelectric development, irrigation demands, and population growth exacerbating the impacts (Martz et al., 2007; MRBB, 2012). Consequently, climate and environmental change here are of concern, not only at local and regional levels but also at the global scale as this impacts the global natural resource and food trades, and regional Earth system change influences the global climate system (RIFWP, 2013).
This area of Canada is the geographic focus of a major research initiative,
the Changing Cold Regions Network (CCRN; DeBeer et al., 2015;
Map of the CCRN geographic study domain in the interior of western
and northern Canada, showing the major river systems, ecoregions, and land
cover, as well as the location of the WECC observatories. Source data are
from the North American Environmental Atlas
(
There has been a substantial amount of previous work aimed at characterizing
and quantifying recent trends and variability in the climate and other Earth
system components over this region. The CCRN builds on a legacy of other
preceding research initiatives, including the Mackenzie GEWEX Study (MAGS;
Stewart et al., 1998; Woo et al., 2008;
In the following sections we describe changes and trends in surface air temperature, precipitation, seasonal snow cover, mountain glaciers and ice fields, permafrost, freshwater ice cover, and river discharge. The focus is generally on regional assessments of change based on extensive observing networks, which provide context for more detailed local observations of change at CCRN WECC observatories and elsewhere. Some principal observation networks and other important sources of regional or long-term data for the detection of change are briefly described. We consider the issue of distinguishing long-term trends from short-term variability or periodicity in records of limited length, together with the role of large-scale, low-frequency oceanic–atmospheric oscillations in driving changes over various timescales. The paper concludes with some remarks on the quality and length of observational data sets, a synopsis of the qualitative linkages among the observed changes in Earth system components, and a short discussion on the complexities of climatic and cryospheric process interactions and hydrological responses based on insights from local-scale observations and experimental studies at some of the WECC observatories. This provides the context for the diagnosis of change to be pursued as subsequent work in CCRN, including the development of improved conceptual understanding of process response and quantitative diagnostic modelling of these changes.
Surface air temperature and its diurnal and seasonal variability is a
fundamental climatic element characterizing a region and affects nearly all
other hydro-meteorological conditions, including evaporation rate and
precipitation phase and type. In cold regions, cryospheric and hydrological
regimes are strongly influenced by air temperature and its variation,
especially around 0
Map of station density and record length for
To facilitate research on climate and environmental change, the Climate
Research Division of Environment Canada has developed an Adjusted and
Homogenized Canadian Climate Dataset (AHCCD;
Recent analyses based on the AHCCD indicate that mean annual air temperature
trends at stations across Canada have been dominated by statistically
significant increases of about 1.5
To illustrate the spatial pattern and magnitude of recent trends in surface
air temperature over western Canada, we analyzed annual and seasonal CANGRD
temperature data for the period 1950–2012. Trends were derived following
Zhang et al. (2000), with a first-order autoregression process used to
adjust the temporal autocorrelations within the climate series. A two-step
approach was used to estimate the autocorrelation parameter (
The results are shown in Fig. 3. Annual mean air temperature trends (Fig. 3a)
show strong spatial coherence with slightly greater warming in the
northern areas (as found in other studies), with statistical significance at
the 95 % confidence level at all grid points. On average, temperature over
the region has increased by just over 2
Spatial patterns of trends (
Large-scale modes of oceanic–atmospheric circulation influence surface air
temperatures on various timescales over western Canada and thus factor in to
the observed trends and interannual variability. These include, for example,
El Niño–Southern Oscillation (ENSO), the Pacific Decadal Oscillation
(PDO), the North Atlantic Oscillation (NAO), and the Arctic Oscillation,
among others. Bonsal et al. (2001b) found that ENSO and PDO influence winter
temperature here, most strongly during El Niño episodes. At these times,
positive PDO phases were associated with strong positive temperature
anomalies, while negative PDO phases were associated with strong negative
anomalies. They noted an increase in the occurrence of El Niño events and
predominantly positive PDO winters after 1976, which would, in part, account
for some of the observed winter warming. Vincent et al. (2015) quantified the
component of seasonal and annual temperature trends associated with the NAO
and PDO over Canada during 1948–2012. Their analysis confirmed that the PDO
signal explained some of the observed trends in winter and spring, accounting
for between a few tenths of a degree up to 2
Not only have average annual and seasonal temperatures increased across the
region, but major changes in extremes such as maximum, minimum, and other
percentiles of monthly and daily temperature have been observed. Earlier
work using the first generation AHCCD showed that over the latter half of
the 20th century most stations exhibited increasing trends in the lower
and higher percentiles of daily minimum and maximum temperature
distributions, and there was a reduction in areas experiencing abnormal and
extreme cold conditions with a concomitant increase in areas experiencing
abnormal and extreme warm conditions (Zhang et al., 2000; Bonsal et al.,
2001a). Bonsal et al. (2001a) noted this translates into fewer days with
extreme low temperature (mainly during winter, spring, and summer) and more
days with extreme high temperature (mostly winter and spring). They also
reported a greater increase in the daily minimum temperatures (i.e., greater
nighttime warming), thereby reducing intra-seasonal standard deviation of
daily temperature. Vincent and Mekis (2006) examined a number of other
indices, including frost days, cold days, cold nights, warm summer days,
warm days, warm nights, diurnal temperature range, and standard deviation of
minimum temperature (see their Table 1 for all definitions). They found
widespread reductions in the number of cold days and nights (maximum/minimum
temperatures < 10th percentile of the corresponding
distributions) from 1950 to 2003 ranging from 10 to 50 days per year fewer
over the period and, to a lesser extent, increases in the number of warm
days, nights, and warm summer days. Other indices showed mostly mixed,
non-significant trends. Mekis et al. (2015) examined trends in extreme heat
and extreme cold events (days with at least one hourly humidex value above
30 and with at least one hourly wind chill value below
Other studies have examined the variation in the intensity, duration, and
frequency of cold and warm spells (defined as 3 or more consecutive days
with minimum/maximum temperatures below/above the 20th/80th
percentile of the corresponding distributions). Shabbar and Bonsal (2003)
found that winter warm spells were increasing in frequency from 1950 to 1998
with up to three or four more events each year over the period and that these were
becoming on average 1–4 days longer. Winter cold spells showed mixed
trends in frequency, but a clear decrease in duration (1–4 days shorter)
and intensity (3–6
Like air temperature, precipitation is a fundamental climatic element characterizing a region that influences the cryospheric and hydrological regime. For instance, the amount, timing, intensity, duration, and phase of precipitation, along with other characteristics such as its seasonality and spatial and elevational distribution, control the formation and development of seasonal snowpacks (and their melt during instances of rain on snow), glacier accumulation, soil moisture and storage conditions in a watershed, and ultimately hydrological responses such as runoff. Annual and seasonal mean precipitation totals vary more strongly than temperature over western Canada, as influenced by latitude, elevation, distance from moisture sources, and other factors. Total amounts vary from up to a few hundred millimetres per year in the north and also parts of the interior plains of southwestern Canada, to many hundreds and up to thousands of millimetres per year in the Rocky Mountains (Environment Canada, 2016). Snowfall accounts for about 30 % of total annual precipitation in the southern interior plains and up to more than 60 % in the far north and in high-elevation areas of the Rocky Mountains (Woo et al., 2008).
As with surface air temperature, the Climate Research Division of
Environment Canada has developed an adjusted precipitation data set for
assessing changes and variability in Canadian precipitation
(
In general, studies using the AHCCD have noted an increasing trend in the
total annual precipitation over most parts of western Canada since about 1950
(Zhang et al., 2000, 2011; Mekis and Vincent, 2011; Vincent et al., 2015). To
provide an up-to-date regional picture of the annual and seasonal trends, we
analyzed the CANGRD precipitation data set over western Canada for the period
1950–2012 using the same methodology as described above for air temperature.
Figures 4 and 5 provide relative changes (as a percentage of the average) and
absolute changes, respectively. On average, annual precipitation has
increased by about 14 % (50 mm) over the region since 1950 (Figs. 4a
and 5a); however, there is considerable variability in the magnitude and
significance of local trends. Most of the increase has been in the north,
where precipitation has risen locally by as much as 60 % (
Spatial patterns of trends (percent per 63 years) in annual and seasonal totals of precipitation over the period 1950–2012 across western Canada, based on analysis of CANGRD precipitation data.
Spatial patterns of trends (millimetres per 63 years) in annual and seasonal totals of precipitation over the period 1950–2012 across western Canada, based on analysis of CANGRD precipitation data.
The seasonal precipitation trends also exhibit large variability across the
region (Figs. 4b–e, 5b–e). Broadly, the spatial patterns of trends in
summer and fall, and to some extent spring, are similar. In winter, there is
a clear divide between increasing trends in the north and decreasing trends
in the south. In most of the northern Mackenzie basin, winter precipitation
has increased by about 30–50 %, while in the southern Mackenzie basin
and most of the Saskatchewan basin it has decreased by about 20–30 %
(and as much as 50 % or more in southern Alberta) (Fig. 4c). Absolute
changes are mostly within about
In addition to changes in the amount of precipitation, there have also been observed shifts in its phase. Zhang et al. (2000) and Vincent et al. (2015) examined trends in the ratio of annual and seasonal snowfall to precipitation totals; from 1948 to 2012, this ratio decreased over southwestern Canada (from 0 to 15 %) and increased over much of northern Canada (by 5–20 %). The greatest changes occurred in spring over the western half of the country, with widespread reductions of up to 20 % or more, reflecting the effects of warmer temperatures. Mekis and Vincent (2011) separately examined annual and seasonal trends in both rainfall and snowfall. They found that over the past 60 years, rainfall totals have increased annually and in all seasons with the most pronounced change during spring, with increases of between 30 and 50 % over much of western Canada. Annual snowfall amounts had decreased across most of southwestern Canada (reductions of 10–30 %) but increased in much of the north (by up to 30 %), with the greatest changes occurring during winter and spring (reductions of up to 50 % in the southwest and increases up to 50 % for some northern stations).
Precipitation variability here is linked to large-scale modes of atmospheric–oceanic circulation, including ENSO, PDO, NAO, and others. El Niño and La Niña events tend to be associated with distinct negative and positive winter precipitation anomalies over southwestern Canada (Shabbar et al., 1997; Gan et al., 2007), while positive/negative phases of the PDO tend to be associated with drier/wetter than normal winters (Whitfield et al., 2010). Bonsal and Shabbar (2008) and Whitfield et al. (2010) provide useful summaries of the influence of these modes on the hydro-climatology of western Canada. While they play an important role in the interannual to inter-decadal variation of precipitation, longer-term trends appear to be mainly independent of this influence. For example, Vincent et al. (2015) assessed trends in annual and seasonal total precipitation and snowfall ratio after removal of the influence of PDO and NAO indices and found that the combination of these indices explained less than 10 % of the trends over 1948–2010 and 1900–2012.
Aspects of the seasonality, frequency, duration, and intensity of
precipitation events have changed over the region in recent decades. For the
period 1950–2003, Vincent and Mekis (2006) reported more days with
precipitation and especially with rainfall (10–30 days per year over the
period) and a decrease in average daily intensity (i.e., total annual
precipitation divided by number of days with precipitation;
0–3 mm day
There have been a number of severe prairie droughts documented over the instrumental record, with multi-year droughts occurring in the 1890s, 1930s, late 1950s and early 1960s, 1980s, and 2000s (Bonsal and Regier, 2007; Bonsal et al., 2011). In the first half of 2015, much of western Canada was experiencing abnormal to record dry conditions – in many areas immediately following a several-year period of historical record wetness and flooding. This led to widespread forest fires, low water levels in streams, lakes, and reservoirs and low soil moisture levels, and it was unusual in that it stretched over a vast area from Mexico to Alaska. Since the beginning of the 20th century there has been decadal-scale variability in drought occurrence as indicated by various precipitation and soil moisture indices, but there has been no consistent long-term trend in drought frequency or magnitude (Millett et al., 2009; Qian et al., 2010; Bonsal et al., 2013). This variability has tended to coincide primarily with precipitation variations modulated by large-scale modes of oceanic–atmospheric circulation (Bonsal et al., 2011; Shabbar and Skinner, 2004; Bonsal and Shabbar, 2008). Through the use of proxy information it appears that extended drought conditions during the 20th century have been relatively mild in comparison to the pre-settlement era on the prairies, and there is evidence of climate-driven non-stationarities in hydrological variables over the past several centuries or millennia (Bonsal et al., 2013; Razavi et al., 2015). Bonsal et al. (2011) provide a useful review of drought research in Canada. The recently completed DRI network was established to conduct a comprehensive study of the severe 1999–2005 Canadian prairie drought (Hanesiak et al., 2011; Stewart et al., 2011).
The seasonal snow cover that forms over the landscape during the winter months is an important feature influencing the climate, hydrology, and ecology of cold regions and mountainous areas (Male, 1980; Gray and Male, 1981). For instance, it has a strong control on the surface energy exchange with the atmosphere, represents an important storage of water that is released during the spring melt period, and affects ground thermal regimes and surface vegetation through its insulating properties. Changes and variability in snow cover therefore have a large impact on many other Earth system processes in these regions, and feedbacks with the climate system can intensify local and regional patterns of change and have global implications (Callaghan et al., 2011b). Snowpack evolution is closely linked with climatic elements such as air temperature and precipitation, and interannual variations and trends in snow cover conditions coincide strongly with those of such climate variables.
The Meteorological Service of Canada (MSC) has produced the Canadian daily snow depth database and the SWE database, based on in situ observations of surface snow cover at climatological stations across Canada and at snow course locations (MSC, 2000). Brown and Braaten (1998) describe the snow depth database, including quality control procedures for internal consistency, the effects of station shifts and urban warming, and the reconstruction of missing values. The data represent about 400 stations with varying record lengths, few of which began before the mid-1940s, while the spatial density and record length tend to decrease considerably in the north. The measurements also tend to be biased to low elevations and open areas (Brown and Braaten, 1998). The SWE database contains weekly, biweekly, or monthly measurements by a number of agencies, but since the late 1990s it has not been actively maintained; in many cases, updates to snow course data can readily be obtained from the various provincial/territorial agencies involved (MSC, 2000). The data primarily cover the period from about 1950 to the mid-1990s, with a pronounced decline after 1985.
Remotely sensed snow cover data sets provide a useful source of information, supplementing in situ observations and extending coverage over broad regions (Hall et al., 2006). Brown et al. (2010), Frei et al. (2012), and Kelly (2012) describe some of the more widely used satellite and model-derived snow cover products. The longest time series is the National Oceanic and Atmospheric Administration (NOAA) weekly snow cover product, which includes near-consistent snow cover mapping since 1966 (Robinson et al., 1993). NASA's Moderate Resolution Imaging Spectroradiometer (MODIS) also provides a range of valuable snow cover products, although data availability is only from 1999 (Hall et al., 2002; Hall and Riggs, 2007). Coarse resolution and classification thresholds lead to uncertainty in these products, particularly in mountain regions (e.g., Brasnett, 1999; Brown et al., 2010). Passive microwave sensors such as the Scanning Multichannel Microwave Radiometer (SMMR), Special Sensor Microwave/Imager (SSM/I), and Advanced Microwave Scanning Radiometer, EOS (AMSR-E), provide information on snow depth, SWE, and melt onset but are limited by variations in snowpack physical properties and wetness that affect the signal, low spatial resolution that does not capture local-scale accumulation, and restrictions associated with both shallow or intermittent snow and deep snow (> 120–150 mm SWE) (Frei et al., 2012; Kelly, 2012).
Over most of Canada there has been a pattern of deceasing snow depths and
snow cover duration and extent since the mid-1970s, with the largest
declines in western Canada and proportionally greater changes later in
winter and spring (Brown and Braaten, 1998; Dyer and Mote, 2006; Derksen et
al., 2008; Derksen and Brown, 2012). Brown and Braaten (1998) analyzed the
Canadian daily snow depth database for the period 1946–1995 and found
widespread and spatially coherent decreases in depth that increased in
magnitude and spatial extent from January through March. Maximum changes
were found over western Canada, where reductions of between 1.0 and 1.5 cm year
To illustrate the pattern over western Canada, Fig. 6 shows recent trends in
snow cover duration from the Canadian daily snow depth data set (1950–2012;
Fig. 6a) and the NOAA weekly snow cover product (1972–2013; Fig. 6b). Trends
were computed using the method of Zhang et al. (2000) described above, with
significance assessed at the 95 % confidence level. The magnitude of
trends in annual snow cover duration have ranged from about 1 to 15 fewer
days decade
At regional to continental scales, snow cover trends are mostly associated with pervasive climate warming patterns independent of low-frequency atmospheric circulation patterns and teleconnections (Derksen and Brown, 2012). Ge and Gong (2009) noted that at these scales, the regional domain of climate mode teleconnections is exceeded. More locally, certain modes of variability (e.g., ENSO, PDO, NAO) influence the interannual variability of snow cover (Derksen et al., 2008; Ge and Gong, 2009; Bao et al., 2011) but are not the main driving factor behind the decreasing trends in western Canada. For instance, Vincent et al. (2015) noted that trends in various snow cover indices from the Canadian daily snow depth data set were almost identical after removal of the influence of PDO and NAO.
Detecting regional changes in SWE is more challenging due to difficulties in estimating variable snow density and problems retrieving SWE information from microwave sensors, and robust information on large-scale, long-term SWE trends from satellite data is not available (Kelly, 2012). Brown (2000) used a combination of in situ snow course and synoptic station snow depths, along with NOAA weekly snow cover data to reconstruct monthly SWE for the period 1915–1997 over North America. He found overall statistically significant increasing trends in SWE for December, January, and February and a significant decrease in April, but he did not provide details on regional patterns. Some work has examined SWE variations over western Canada using SMMR and SSM/I data (Walker and Silis, 2002; Derksen et al., 2002, 2003; Tong et al., 2010), but the analysis periods were relatively short and clear trends could not be detected. Derksen et al. (2004) combined SMMR and SSM/I with in situ SWE observations over the provinces of Saskatchewan and Manitoba and several states in north-central USA to examine SWE variability from 1915 to 2002. Derksen et al. (2008) used in situ observations, NOAA weekly snow cover, and SMMR and SSM/I SWE retrievals to examine trends over the Mackenzie River basin from 1950 to 1999. In both studies, monthly SWE in spring did not show clear long-term trends but rather indicated decadal-scale variability that varied regionally. Squires (2014) focused on a more limited part of the Mackenzie basin around the Great Slave Lake region using snow course data collected by Aboriginal Affairs and Northern Development Canada to assess trends in snow depth and SWE from the mid-1960s up to 2010. They reported progressively earlier onset of snowmelt across most of this area, along with regionally varying changes in maximum SWE.
Mountain glaciers represent a major geomorphic agent shaping the landscape, directly affect local surrounding climatic conditions, and provide a source of runoff well beyond the spring snowmelt freshet period (Moore et al., 2009). Glacier mass balance responds directly to variations in snow accumulation and snow and ice ablation, as influenced by air temperature, precipitation, radiation, and other meteorological variables, while extended perturbations in the mass balance tend to produce a delayed response in glacier geometry, surface elevation, and terminus position (Oerlemans, 1989). Thus, patterns of glacier behaviour over time provide a proxy indicator of the local to regional prevailing climatic conditions. Globally, mountain glaciers have predominantly been losing mass and retreating, and there is concern that as this continues their ability to provide a reliable source of meltwater runoff during the summer and fall period will be reduced (Radić and Hock, 2014). In western Canada, glacier runoff has been shown to provide a buffering mechanism during extremely dry years (Hopkinson and Young, 1998; Comeau et al., 2009), but as glacier cover diminishes they may no longer be able to effectively augment low flows (Demuth and Pietroniro, 2003; Marshall et al., 2011).
Map of western Canada showing the distribution of glaciers and ice
fields within the major interior mountain ranges; the locations CCRN glacier
observatories are indicated. Glacier extents are from the GLIMS database
(
Glaciers and ice fields occur extensively throughout the mountain regions of
western Canada, covering an area of roughly 50 000 km
Detailed in situ observations and mass balance records are available only for a very limited number of glaciers over the region. Ommanney (2002a) and Østrem (2006) provide accounts of the history of glaciological investigations in western Canada and the selection of reference mass balance glaciers. The most well-documented and studied glacier in the region is Peyto Glacier (Fig. 7), which has been documented since the late 1800s and has near continuous annual and seasonal mass balance records going back to 1965, beginning as part of the IHD activities (Østrem, 2006; Demuth and Keller, 2006). In contrast, monitoring activities at most other glaciers have been short-lived and are largely discontinued. A concern for ongoing initiatives is the retreat and disintegration of glacier termini, making access difficult and restricting continuity of measurement programmes.
Net annual mass balance series for Peyto Glacier (1966–2012) and cumulative mass balance over the 46-year period. Note that values in 1991/92 are reconstructed from proxy information and values in 2010–12 are preliminary (data provided by M. Demuth, Geological Survey of Canada).
Glaciers across western North America have predominantly retreated over the
past century and the observational records have shown mostly negative net
mass balances (Moore et al., 2009). As shown in Fig. 8, the net annual mass
balance of Peyto Glacier has exhibited strong interannual variations but
has been mostly negative since 1965 (average
The rates of retreat have varied, with initially rapid rates in the first half of the 20th century, followed by a short-lived period of stabilization or advancement until 1980 and continuing retreat since then (Moore et al., 2009). The records at the Athabasca and Saskatchewan glaciers, together with observations elsewhere in the Rockies, show a sharp decline in glacier recession in the 1960s and 1970s, attributed to a short-lived period of positive net mass balance (Luckman, 1998). McCarthy and Smith (1994) documented historical glacier length variations within the southern Rockies and showed that individual glaciers displayed a wide range of patterns, with some stabilizing after the 1950s, others slowing in their rate of retreat, and still others exhibiting an accelerated recession beginning in the last few decades of the 20th century.
Glacier area changes have been investigated in many recent studies and, at the regional scale, have been reported to have declined from as little as 11 % over 53 years to as much as 29 % over 13 years, with most studies reporting about a 25 % decline in glacier cover since the mid- to late 20th century (Hopkinson and Young, 1998; DeBeer and Sharp, 2007; Schiefer et al. 2007; Demuth et al., 2008, 2014; Jiskoot et al., 2009; Bolch et al., 2010; Tennant et al., 2012; Tennant and Menounos, 2013; Beedle et al., 2015). Individually, glaciers have exhibited a wide range of local changes from small net advances to complete disappearance, reflecting the strong control of local factors on glacier mass balance, dynamics, and response. A number of features have been commonly reported in western Canada: (1) glaciers have generally been continuing to lose mass and retreat, and rates of loss have accelerated in the last few decades; (2) smaller glaciers have tended to exhibit greater variability in their relative area changes, while larger glaciers have exhibited more consistent and moderate relative changes; and (3) collectively most of the glacier area loss has been due to the retreat of larger glaciers that comprise a greater proportion of the regional ice cover.
Volumetric changes have been measured through comparison of repeat digital
elevation models, which has indicated thinning rates from about 0.5
to 0.9 m year
In western Canada, several recent studies have compared glacier losses with observed discharge and used various modelling approaches to estimate the proportional contribution by glaciers (Hopkinson and Young, 1998; Comeau et al., 2009; Marshall et al., 2011; Jost et al., 2012; Naz et al., 2014; Bash and Marshall, 2014). Glacier wastage, referring to ice loss as a result of any negative net mass balance, has been found to account on average for about 1–5 % of the annual discharge of the larger rivers exiting the eastern slopes of the Rockies and up to 10 % or slightly more of the summer (July–September) flow (Comeau et al., 2009; Marshall et al., 2011; Bash and Marshall, 2014). In an extremely low flow year, Hopkinson and Young (1998) found that glacier wastage supplied 13 % of the annual flow of the Bow River at Banff, with a maximum monthly contribution in August of 56 %. Jost et al. (2012) found that glacier ice melt (excluding seasonal snow cover) supplied an average of 6 % of the annual flow of the upper Columbia River between 1970 and 2007, which increased to as much as 25 and 35 % of August and September flow. Naz et al. (2014) simulated glacier contributions to the upper Bow River basin over 1981–2007. Their results indicated that while summer and annual glacier melt showed increasing tendencies, summer and annual discharge showed decreasing tendencies. They speculated that the decline in discharge is associated with a combined effect of decreases in glacier cover and precipitation. Demuth and Pietroniro (2003) observed that the discharge from glacierized catchments in the North Saskatchewan River basin during the transition-to-baseflow (TBF) period has decreased since the mid-20th century, despite increasing precipitation during the TBF period and greater ice melt. This response is commensurate with reductions in glacier area. They also observed an increase in the variability of discharge during the TBF period, pointing to a reduction in the buffering capacity of glaciers in this system.
Permafrost, defined as ground that remains at or below a temperature of
0
Distribution of permafrost over western Canada (from Brown et al., 2001) and mean annual ground temperature (MAGT) recorded during the IPY (2007–2009) (from IPA, 2010).
Two of the largest monitoring networks are maintained by the Geological Survey of Canada (GSC) and the Centre d'études nordiques, Université Laval. The GSC, in collaboration with other partners, has been developing and maintaining a network of sites that contribute to the Canadian Permafrost Monitoring Network as well as the Circumpolar Active Layer Monitoring (CALM) programme and the Global Terrestrial Network for Permafrost (GTN-P), providing long-term measurements of permafrost thermal state and active layer conditions. About 75 sites operated by government and university scientists contribute to these networks, with more than 20 boreholes available since the mid-1980s of up to 20 m depth in the Mackenzie River valley and delta in the Northwest Territories (Fig. 9; Smith et al., 2005). A number of additional monitoring sites were established during the IPY period 2007–2009, bringing the total number of sites in North America up to 350 (Smith et al., 2010). The distribution of these sites tends to be along roads and pipelines, while the boreholes themselves vary in depth, measurement technique, and recording frequency; the network, however, provided a useful snapshot of permafrost in northern North America for the IPY period (Smith et al., 2010). It is important to note that for all ground temperature and active layer thickness data, local site characteristics and conditions can have a considerable influence on the variability of permafrost conditions (e.g., Burgess and Smith, 2003; Smith et al., 2009). Useful reviews of permafrost conditions and their recent changes over northern Canada are provided by Smith (2011), Burn (2012), Derksen et al. (2012), and Bush et al. (2014).
Observations from monitoring sites across northern Canada have shown that
ground temperatures to depths of 10–15 m or more have been increasing in
response to recent climate warming (Smith et al., 2005, 2010). The magnitude
of this warming has varied with latitude, generally exhibiting greater
trends in the northern continuous permafrost zone and more limited
temperature change in the southern discontinuous and sporadic zones, where
permafrost thaw has been widespread. Smith et al. (2005) and Smith (2011)
presented temporal variations in permafrost temperature from the Canadian
Permafrost Monitoring Network and the GTN-P. They showed that in the
southern Mackenzie Valley, where permafrost is warmer than
As ground temperatures have risen, there has also been an increase in the
thaw depth of the seasonal active layer across northern Canada and local to
widespread permafrost degradation in the discontinuous permafrost zone
(Smith, 2011). In many areas, the melting of the near-surface ground ice has
resulted in decreased stability and strength of the substrate, leading to
surface subsidence and waterlogging of soils (both of which further amplify
permafrost thaw), thermokarst development, and collapse of forested peat
plateaus (Jorgenson et al., 2008; Smith et al., 2008; Baltzer et al., 2014).
Burn and Kokelj (2009) reported interannual variability of maximum seasonal
thaw depths that ranged between 0.3 and 0.55 m, with a trend toward
increasing thaw depths of 0.08 m over the period 1983–2008 in the outer
Mackenzie Delta area. Smith et al. (2009) examined summer thaw depths over
Studies have examined changes in the spatial extent and connectivity of permafrost in the Taiga and Boreal Plains within the discontinuous zone and generally found that it is shrinking in area and becoming increasingly fragmented. Beilman and Robinson (2003) used historic aerial photos and recent satellite imagery to compare permafrost extent between about 1950 and 2000 and found that at peatland sites in the Northwest Territories, as much as 50 % of permafrost had degraded and thawed, while in more southern locations in Alberta, Saskatchewan, and Manitoba, between 30 and 65 % of localized permafrost has thawed. Similar analysis at the Scotty Creek Research Basin near Ft. Simpson (Quinton et al., 2011) showed that the extent of forest-covered permafrost plateaus had decreased by over 38 % in localized areas between 1947 and 2008, with associated collapse of black spruce forest on the plateaus and an increasingly connected surface drainage network. This has the potential to substantially alter basin runoff production in the region (Connon et al., 2014). Further, it has been shown that the rates of permafrost and forested plateau loss have been accelerating, with the average rate of loss in seven different local areas of interest being over 3 times greater during the period 2000–2010 compared to 1977–2000 (Baltzer et al., 2014).
Virtually all lake and river systems in the interior of western Canada are seasonally ice-covered, and this influences the timing and severity of low flows and floods associated with freeze-up and break-up periods (Beltaos and Prowse, 2009; Prowse, 2012) and has an important role in the ecology of these systems (Prowse, 2001; Prowse and Culp, 2003). Spatially, ice thickness ranges from skims in more temperate southern parts of this region to several metres in colder northern parts, while ice cover duration ranges from being a transient feature to existing for over 6 months of the year (Prowse, 2012). The formation and melt of the ice cover are sensitive to air temperature, precipitation, and other meteorological variables, which influence the composition, thickness, and stability of the ice (Beltaos and Prowse, 2009). In addition to thermal properties, river ice formation and break-up also depend on hydrodynamic processes that affect the mechanical properties of the ice (Beltaos, 2003). Ice cover thickness, timing (i.e., freeze-up and break-up), and duration therefore respond to changes in climatic conditions and other controls such as changes in terrestrial hydrologic regimes (Prowse et al., 2011a, b). Some useful reviews on the climatic controls, historical trends, and future projections of river ice formation and break-up are provided by Prowse and Beltaos (2002), Beltaos and Burrell (2003), Prowse and Bonsal (2004), Prowse et al. (2007), Beltaos and Prowse (2009), and Prowse (2012).
There are limited in situ records of ice cover variables over the interior of
western Canada, which poses a challenge for systematic regional monitoring
and assessment of ice cover regimes and their changes over time. Various
observations on freshwater and coastal sea ice conditions in Canada have
been gathered into a common database known as the Canadian Ice Database
(CID) in order to aid climate monitoring efforts and improve numerical
prediction models and remote sensing methods (Lenormand et al., 2002). The
database contains records related to ice thickness and freeze-up and break-up
dates for 757 sites across Canada (including 312 lakes and 288 rivers).
During the late 1980s and 1990s, however, there was a drastic decline in the
observation network and more recent observations only represent a very small
percentage of those available in the mid-1980s. Lenormand et al. (2002)
provide a complete description of the CID and its historical evolution. A
volunteer monitoring program, IceWatch (
Observations have shown a general reduction in ice cover duration on lakes
and rivers across much of Canada since the mid-20th century due
primarily to earlier spring break-up (Prowse, 2012). Zhang et al. (2001a)
and Burn and Hag Elnur (2002) analyzed trends in a number of hydro-climatic
variables from the Canadian Reference Hydrometric Basin Network (RHBN)
database (see next section) and found that the break-up of river ice and
the spring freshet showed significant trends toward earlier occurrence at
many sites in Canada over the latter half of the 20th century. The most
pronounced changes were observed in western and southwestern Canada, with a
greater degree of change over the last few decades of the century. Lacroix
et al. (2005) used observations from the CID to examine spatial trends in
river freeze-up and break-up over Canada and related these to the timing of
the fall and spring 0
The magnitude of trends in ice cover formation, break-up, and overall
duration has varied among studies, depending on data sets and temporal
intervals of analysis. Lacroix et al. (2005) reported trends in river ice
break-up over Canada ranging from 1.0 to 2.2 days earlier decade
Other characteristics of the seasonal ice cover are important, such as trends in the severity of river ice break-up events and ice thickness, but there have been few broad-scale analyses in Canada. In the Mackenzie Delta, a widespread tendency toward earlier break-up initiation and peak water level has been reported, associated with a tendency towards the occurrence of more ice-driven break-up events as opposed to discharge-driven events (Goulding et al., 2009). Ice-driven events are characterized by higher upstream backwater levels from ice jams and earlier break-up initiation of a more competent ice cover. It has been suggested that mid- to late-spring air temperatures during the pre-break-up melt and runoff period more strongly influence the timing of break-up than overall winter temperatures and maximum ice thickness (Prowse and Bonsal, 2004). Prowse (2012) noted that the CID observations of Lenormand et al. (2002) over Canada do not show any obvious trends in ice thickness over the latter part of the 20th century.
The flow of water in rivers and streams provides a vital water resource to society and serves an important ecological role; however, hydrological extremes represent serious potential risks and costs to society, such as during floods or droughts. River discharge represents the spatially integrated response to inputs from rainfall, snowmelt, and glacier melt, and its character is influenced by internal watershed storage capacity, vegetation, soils, topography, and other landscape features affecting the delivery of water to the stream channel network. In cold regions, the response to input events is highly sensitive to antecedent conditions in the basin, including preconditioning of soils, snowpacks, surface and sub-surface water storage, and surface drainage network connectivity (Spence and Woo, 2008; Carey et al., 2010; Bring et al., 2016), while the nature of response is also dependent on the character of precipitation and the storm type (e.g., Shook and Pomeroy, 2012). The timing, magnitude, duration, and other characteristics of discharge are therefore closely linked with climate and other Earth system components, and they are sensitive to their changes within the watershed over multiple timescales.
The WSC maintains a network of over 2000 hydrometric monitoring stations
across Canada, and daily and monthly mean streamflow and stage levels are
available through their hydrometric database (HYDAT;
Many studies have examined variability and trends in the magnitude, timing, and other characteristics of river discharge over Canada (e.g., see Koshida et al., 2015; Mortsch et al., 2015). Direct comparison of the results is often difficult due to differences in temporal analysis periods, statistical methodology, region of focus, and data sets (i.e., RHBN and non-RHBN). Annual mean flow has been observed to vary regionally, with studies documenting both increasing and decreasing trends since the 1960s (Prowse et al., 2009b). Major river systems such as the Mackenzie and Nelson (which includes the Saskatchewan River system) have shown no detectable long-term trends at their mouths over this time (Woo and Thorne, 2003; Déry and Wood, 2005; McClelland et al., 2006; Déry et al., 2011), while statistically significant declines in annual flow have been observed for some smaller systems within these, such as the Athabasca River and its tributaries, and other rivers draining from the eastern slopes of the southern Rocky Mountains (Burn et al., 2004b; Rood et al., 2005; St. Jacques et al., 2010; Peters et al., 2013). Seasonally, a consistent pattern of increasing flow in the winter has been reported – especially for the Mackenzie River and many of its northern tributaries – with significant trends in the annual minimum discharge and lower flow percentiles (Burn et al., 2004b; Rood et al., 2008; St. Jacques and Sauchyn, 2009). Discharge rates during other times of the year have varied, but studies across the region have found decreasing trends in the annual maximum discharge rate and spring freshet flow volume (Burn et al., 2004b, 2008, 2010). Burn et al. (2008, 2010) have also reported increasing trends in the number of rainfall–runoff events later in summer and their peak magnitude, but little trend in runoff volume of these events.
Spatial patterns of trends in
To synthesize recent changes in flow across western Canada, we performed an
analysis of monthly and annual discharge from RHBN records. Trends were
computed for three different periods beginning in 1960, 1970, and 1980, and
all ending in 2010. The analysis was restricted to currently active stations
and excluded most higher-order rivers as these are mostly affected by
upstream regulation and other human influences (and thus are not part of the
RHBN). Missing data thresholds were set at no more than 3 days month
Annual and January trend results are shown in Fig. 10. Trends since 1960 are not shown as most were rejected due to missing data. Annual flows exhibited a mixture of increasing and decreasing trends (Fig. 10a), while their magnitude, significance, and, in some cases, direction depended on the length of the analysis period. There is no well-defined regional pattern of consistent, long-term annual discharge change, except perhaps in the southern mountain regions of British Columbia and Alberta, where flows at a number of stations have decreased slightly over time. During the winter months there was a clear pattern of increasing flows across the north, with many stations exhibiting significant positive trends in all three periods (January flows in Fig. 10b provide a good example). Due to the low flow rates at this time of year, however, this has little influence on the annual mean flow of most rivers. In the southern half of the domain, trends in winter and spring discharge showed no consistent regional pattern, except later during summer and fall (not shown), when many streams and rivers in the North Saskatchewan and Peace–Athabasca River basins exhibited significant declining flows.
The most consistently reported characteristic of discharge trends across
western Canada has been an earlier onset of the spring freshet since at least
the mid-1960s, as indicated by features such as the initial hydrograph rise
and the timing of peak spring flow (Woo and Thorne, 2003; Burn et al., 2004a,
b, 2008, 2010; Burn, 2008; Abdul Aziz and Burn, 2006; Rood et al., 2008;
Cunderlik and Ouarda, 2009). Woo and Thorne (2003) reported that between 1973
and 1999 the date of spring hydrograph rise for the Mackenzie River and
several of its major tributaries (i.e., the Peace, Liard, and Slave Rivers)
advanced by about 3 days decade
The timing of low flows has also received attention (Ehsanzadeh and Adamowski, 2007, 2010; Khaliq et al., 2008; Burn et al., 2010). Annual low flow events of various durations have generally been observed to show a decreasing (i.e., arriving later) trend in southwestern Canada and an increasing trend in the northwest (Khaliq et al., 2008). In the prairie region, Burn et al. (2008) found a decreasing trend in the timing of summer rainfall-driven peak flow events, while over a broader part of western Canada, Cunderlik and Ouarda (2009) found no significant trends in the timing of fall rainfall-dominated high flow events.
Given the limited record length for flow analyses, the potential for computed trends to be influenced by shorter-term variability and large-scale modes of oceanic–atmospheric circulation variability is high. Woo et al. (2006) discuss this issue and illustrate the difficulty in attributing trends detected in short records to long-term climate change. Some studies have accounted for modes of large-scale variability to examine the trends in the absence of their influence. For example, St. Jacques et al. (2010) removed the signal due to the PDO in flows at a number of stations across the southern Canadian Rockies and southern Alberta, and they found that observed and naturalized flows still exhibited predominantly decreasing trends since as far back as the early 20th century. Burn's (2008) analysis of flow timing indices in the Mackenzie River basin examined the relationship with meteorological variables and six large-scale circulation indices. They concluded that although several of the timing measures are related to large-scale circulation patterns, there is a relationship with increasing spring air temperatures, and the large-scale periodic influences are not a dominant contributor to the observed trends. Bonsal and Shabbar (2008) review the past Canadian research on the impacts of large-scale circulation patterns on low flows and report a higher frequency of low flow events in western Canada associated with the warmer and drier conditions during El Niño events and positive phases of the PDO and the Pacific/North American (PNA) pattern. They also note that the spatial and temporal aspects of the relationships between low flows and these climatic patterns are strongly influenced by local hydro-climatic complexities, particularly in the mountainous watersheds.
The observational data sets described throughout this paper provide the best available long-term, quality-controlled products for examining Earth system change and variability in western Canada. Considerable effort has gone into ensuring their reliability for trend detection and their intercomparability for regional assessments, including corrections for known sources of inhomogeneity, station relocation, measurement error, and data gaps. The data are, however, subject to several notable limitations. First, the hydro-climatic observational network is sparse in northern Canada and regional assessments require extrapolation and interpolation over vast areas. This is likely less problematic for air temperature than for precipitation, for example, which can exhibit greater local variability and is more sensitive to the density of the observing network and methods for adjusting systematic measurement errors (Rapaić et al., 2015). Rapaić et al. (2015) noted that considerable care is needed when using gridded climate data sets in local or regional-scale applications in data sparse regions such as the Canadian Arctic, where, for example, they found that the adjustments used in the CANGRD product led to a doubling of precipitation trend magnitude compared to a multi-data set average. Sparseness of the observing network is also an issue in mountainous areas, where there are relatively few long-term sites at high elevations – a problem that exists globally (Adam et al., 2006). Monitoring in cold, harsh conditions and the associated measurement error also present a challenge; in particular, solid precipitation measurements are highly susceptible to error as a result of wind and turbulence around the gauge and require careful correction (Yang et al., 2005).
Another issue involves the record length of these observational data sets. The longest climate station records go back to the early 20th century or late 19th century for a few select sites, while across much of the northern parts of western Canada, there are very few climate records going back much further than about 1950. This is also a problem for other variables, such as snow and freshwater ice cover, and particularly river discharge. As noted previously, it is difficult to discern long-term trends from periodic effects in short records, and the results may be highly sensitive to analysis period. Large-scale oceanic–atmospheric circulation patterns, such as ENSO and PDO, are known to influence the variability of cryospheric and hydrological systems on timescales of years to decades (Whitfield et al., 2010) and can thus significantly complicate the detection and attribution of trends (Woo et al., 2006). In many instances, the closure of stations or the automation of measurement procedures has caused gaps or inconsistencies in the data, also limiting the length of observational data sets. Since about the mid-1990s there has been a considerable reduction in active climate, snow, and discharge monitoring stations across this region, which restricts the analyses of long-term change.
Despite the data limitations noted, observations have shown clear and
systematic patterns of change in climatic regime and cryospheric response
over western Canada. The various lines of evidence are consistent and
mutually supportive. Warming has been pervasive, especially during winter
and spring and at higher latitudes (warming rates in western Canada are
among the highest globally), while changes in precipitation have been more
varied, both regionally and seasonally. Widespread decreases in winter
precipitation in much of southwestern Canada have been observed, and a
decline in the fraction of precipitation falling as snow is very likely
associated with rising winter/spring temperatures and a shift in the timing
of the 0
In general, river discharge magnitude has not shown long-term and regionally coherent trends, other than increasing winter flows in the north. This may partly reflect aspects of periodicity and the influence of large-scale modes of climate variability over interannual to inter-decadal scales (e.g., ENSO, PDO). However, it has been found that these modes are not the sole driver of observed trends in other variables (Vincent et al., 2015). The mixed responses in discharge magnitude are probably the result of interactions between multiple processes and drivers – which may be confounding and multi-directional – across various temporal and spatial scales and geographic locations.
While in most cases, the length of record at the CCRN WECC observatories (Fig. 1) precludes robust time series analysis as conducted on the national climatological and hydrometric data sets, the observatories provide critical platforms to understand hydrological responses and process interactions and to diagnose how changes are manifested at scales where field-based scientific investigations are undertaken.
One of the most notable changes in the cryosphere is the widespread thaw of
permafrost across the circumpolar region (Hinzman et al., 2005; Schuur et
al., 2015). In the Mackenzie and Yukon River basins of North America, recent
changes in winter low flows have been ascribed to permafrost thaw and
adjusting pathways (Walvoord and Striegl, 2007; St. Jacques and Sauchyn,
2009). In contrast, at the headwater and mesoscales
(< 1000 km
Several WECC observatories have been operational since the 1990s in permafrost-underlain catchments and have undergone change associated with warming. The most dramatic change has been observed at Scotty Creek in the Taiga Plains near Fort Simpson, NT, where degradation in permafrost has resulted in decreasing and fragmented forest cover (Baltzer et al., 2014) and increasing wetland coverage and connectivity and drainage efficiency (Quinton et al., 2011; Connon et al., 2014). From 1947 to 2008, permafrost coverage declined at the site from 70 to 43 % (Chasmer et al., 2010) and wetland permafrost-free areas expanded and coalesced. Connon et al. (2014) showed that increases in flows observed in the region were not a result of reactivation of subsurface flow pathways from permafrost thaw but from an increased connectivity of surface pathways and, to a lesser extent, from increases in overall precipitation. Bogs that were formerly isolated on the landscape became incorporated into the basin drainage network, increasing runoff efficiency through time.
Further north at the Baker Creek, NT, WECC observatory, changing precipitation regimes from warmer temperatures are more strongly influencing the hydrological response of this watershed. Here, the Precambrian Shield geological setting, abundant lakes, and thin glacial–fluvial soils are not particularly sensitive to changes in thermal regime, but changes in lake storage can rapidly change the runoff contributing area available for river discharge formation. The WECC observatory has a distinct threshold–response runoff regime (Spence, 2006) where increasingly large areas of the watershed connect and can contribute to runoff as lake storage increases (Phillips et al., 2011). Spence et al. (2011) have documented sudden changes in streamflow regimes as the climate has changed. Previously, peak flows were observed during spring freshet in response to the relatively rapid snowmelt, whereas in more recent years, large summer and fall precipitation events that are rare in the historical record have resulted in enhanced late-season flows. In addition, Spence et al. (2015) revealed the emergence of new streamflow regimes as a result of late-season rains, with enhanced winter storage and streamflow being characteristic of a transition from a nival toward a pluvial regime, and they showed how this influences the biogeochemical system.
Other WECC observatories underlain with permafrost (notably Wolf Creek, YT, and Trail Valley Creek, NT) have undergone less dramatic changes since the 1990s despite warming air temperatures, although important insights to potential response from change are emerging. In Trail Valley Creek where permafrost is continuous, active layers have increased and there have been gradual expansions of taller shrub vegetation (Lantz et al., 2012). Shi et al. (2015) examined several timing metrics of the runoff hydrograph and revealed complex responses for 27 years of data. Unlike what has been reported in southern latitudes (Berghuijs et al., 2014), a decline in total winter snow accumulation had only minor influence on the timing of springtime streamflow. Offsetting effects of springtime temperature fluctuations and winter warming resulted in small changes to the freshet hydrograph, whereas delayed springtime rain had a more notable impact on the late-summer flow regimes. Overall, despite major changes in climate, vegetation, and active layer thickness, changes to streamflow were limited. At Wolf Creek, an alpine watershed underlain with discontinuous permafrost, recent changes in shrub abundance at higher elevation has resulted in some changes in accumulation of snow (Pomeroy et al., 2006), yet responses in the streamflow hydrograph to change have yet to be realized with 18 years of data. Using a hydrological model to simulate future climates, Rasouli et al. (2014) showed that while the hydrological impacts of climate warming are unequivocal (i.e., reduced snow contributions to streamflow, shorter snow-covered period, greater evapotranspiration), the magnitude and direction of the impact of warming on streamflow depends more on changes to future precipitation than on future warming.
In more southerly locations, the influence of land use and climate change combined to dramatically alter hydrological regimes in Smith Creek, MB, an agricultural basin characteristic of the southeastern prairies. Warming temperatures, increased rainfall fraction, earlier melt, and more multiple-day rainfall events have, in concert with extensive wetland drainage and increase in basin hydraulic connectivity, resulted in an overall dramatic increase in flows generated from snowmelt, rain-on-snowmelt, and rainfall runoff processes, with the greatest increases for rainfall runoff and a relative decline in the proportion of streamflow derived from snowmelt from over 85 % in the 20th century to less than 50 % in the last 5 years (Dumanski et al., 2015). Conversely, in the eastern slopes of the Rocky Mountains, Marmot Creek, AB, has shown considerable resilience in its flow regime. Despite winter warming, declining low-elevation snowpacks, forest cover reduction, and, more recently, extreme weather, streamflow volume, timing, and magnitude of peak have not changed since 1962 (Harder et al., 2015). An increase in the fraction of precipitation occurring in spring and the frequency of multiple-day rainfall events may have counteracted warmer conditions and lower snowpacks at low elevations. Substantial sub-surface storage in the basin may further buffer the impacts of climate variability. The devastating flood of 2013 that affected Calgary and much of the eastern slopes was in fact buffered by processes operating within headwater catchments as flow responses were less than anticipated based on precipitation statistics. Of particular interest is that 2 years later in 2015, despite relatively normal total volumes of winter precipitation, smaller snowpacks (due to more mid-winter rainfall and melt events) have resulted in dramatically lower streamflows and record low groundwater levels, exacerbating the ecological and societal impact of drought conditions that arrived in the spring.
Ongoing research at the WECC observatories will continue to address the interplay among changes in climate, cryosphere, ecology, and the hydrological responses observed, and place larger-scale patterns in a local context. There are a number of outstanding research questions that remain to be addressed and will be a focus of ongoing research: (1) the role of changing vegetation and ecological assemblages (including land-use change) on hydrology at the WECC observatories; (2) the importance of fall and over-winter hydrological processes, which are often neglected in field-based research yet are increasingly being identified as important when evaluating change; (3) the influence of thawing permafrost and a reduction in seasonally frozen ground on hydrological response across a gradient of watersheds in different thermal and physiographic environments; and (4) how the decline in snow (in terms of volume, duration, and as a fraction of total precipitation) will influence all aspects of the hydrological cycle. An important activity of CCRN will be to develop conceptual and numerical models of change for each WECC observatory using future climate scenarios and a common experimental framework linking all biophysical aspects of these systems. While large national observation networks are critical for baseline assessment of change, linking this to research at small scales is still required to interpret patterns observed in larger basins and will be a focus of the network.
To conclude we note that rapid warming is occurring throughout the interior of western Canada, as in other cold region environments globally, with profound implications in terms of landscape and hydrological and change. However, hydrological responses are complex and spatially variable, reflecting multiple process interactions at sub-basin scale that vary across the region, and depend strongly on changes in precipitation form and timing. There has been much speculation on the future trajectory of hydrological change in cold regions in relation to climate projections, but this has not often been grounded on appropriate conceptual understanding and process-based representation of the underlying response mechanisms. Given the complexities involved, this is required to develop plausible scenarios of change for the 21st century, especially considering that while certain processes and their interaction are inherently accounted for in some physically based models, other factors are likely not (e.g., land cover changes such as deglaciation, permafrost degradation, changing vegetation and ecological assemblages, and their influence on hydrology). Simple projections of change are therefore likely to be misleading. There is clearly a need to diagnose and better understand the causes and mechanisms underlying the observed pattern in flow regime and its variability over the region if projections of future change (and associated societal impacts) are to be made with any confidence.
All data sets used here are publicly available and can be accessed through the links and references we have provided.
We gratefully acknowledge the Natural Sciences and Engineering Research Council of Canada (NSERC) for funding the CCRN through their Climate Change and Atmospheric Research (CCAR) Initiative. We also thank Kevin Shook and Ross Brown for their contributions in trend analysis of the hydrometric and snow cover data, Mike Demuth for providing updated Peyto Glacier mass balance data, and John Pomeroy for help with editing the manuscript. Two anonymous reviewers provided comments that were greatly beneficial and constructive, and we thank them for this. Edited by: B. Schaefli