Changes in Snow, Ice, and Permafrost Across Canada

Based on an analysis of multiple datasets covering 1981–2015, SCF (characterized as the proportion of days within each month that snow was present on the ground) decreased by 5% to 10% across most of Canada during most seasons (Mudryk et al., 2018¹¹³; see Figure 5.2), notably, for eastern Canada in spring (April/May/June) and most of the Canadian land area in the fall (October/November/December). This loss of snow cover is consistent with previous studies using in situ datasets covering a longer time period (Brown and Braaten, 1998²⁰; Vincent et al., 2015¹⁸⁹), but the 1981–2015 period is characterized by stronger reductions in snow cover during the snow onset period for eastern Canada in response to enhanced fall warming (consistent with Brown et al., 2018²⁵). Decreasing SCF trends over high latitudes of Canada are consistent with documented reductions in annual snow cover duration (SCD; the number of days with snow cover) across circumpolar Arctic land areas of two to four days per decade (approximately 1% to 2% per decade, assuming 250 days mean snow cover) (Brown et al., 2017²⁶). Some studies (Derksen and Brown, 2012³⁶; Derksen et al., 2016³⁷; Brutel-Vuilmet et al., 2013²⁷; Hernández-Henríquez et al., 2015⁶⁸; Mudryk et al., 2017¹¹⁵) identified spring snow cover losses slightly stronger than those in Figure 5.2, because different datasets and time periods were considered. Despite these differences, all studies consistently show reductions in spring SCF.

Analysis of surface temperature from a blend of six atmospheric reanalysis datasets showed that warming trends over the 1981–2015 period are found in all Canadian land areas with SCF reductions (Mudryk et al., 2018¹¹³). Cooling trends in winter and spring are associated with the regions of increasing SCF (see Figure 5.2). Observations from climate stations in the regions where SCF trends increased over 1981–2015 also show decreased maximum snow depth and SCD over the longer 1950–2012 period (Vincent et al., 2015¹⁸⁹), so the positive trends over 1981–2015 reflect nature variability in regional surface temperatures and precipitation.

While SCF information is important for identifying changes in where snow covers the ground, from a water-resources perspective, it is important to understand how much water is stored in the form of snow. This is determined from the pre-melt SWE_max. SWE_max declined by 5% to 10% across much of Canada during the period 1981–2015, according to the multi-dataset analysis shown in Figure 5.3 (Mudryk et al., 2018¹¹³). This is consistent with snow depth trends from surface measurements (Brown and Braaten, 1998²⁰; Vincent at al., 2015¹⁸⁹) and other observational studies (for example, Mudryk et al., 2015¹¹⁴). Increases in SWE_max are evident across parts of British Columbia, Alberta, and southern Saskatchewan. The influences of temperature and precipitation changes need to be separated to understand the driving mechanisms behind trends in SWEmax (Raisanen, 2008¹⁴¹; Brown and Mote, 2009²⁴; Mankin and Diffenbaugh, 2014¹⁰⁶; Sospedra-Alfonso and Merryfield, 2017¹⁶⁶).

Projections of surface temperatures across Canada for the near-term under a high emission scenario (RCP8.5) show warming in all seasons in the multi-model average (see Chapter 4, Section 4.2.1.3), with concurrent decreases in projected SCF across all of Canada during all seasons (Figure 5.4; Mudryk et al., 2018¹¹³). During winter, projected snow cover reductions will be greatest across southern Canada, where temperature increases result in less snowfall as a proportion of the total precipitation. Temperatures will remain sufficiently cold at higher latitudes that winter (January/February/March) SCF in this region is not projected to change in response to warming. During spring, the region of snow sensitivity to temperature forcing is projected to shift north, as snow cover retreats across the boreal forest, sub-Arctic, and high Arctic. This leads to projected negative SCF trends (loss of snow) across these regions during the April through June period. Important differences in spring snow cover projections between emissions scenarios emerge by the end of the century, with stabilized snow loss under a medium emission scenario (RCP4.5) but continued loss under a high emission scenario (RCP8.5) (Brown et al., 2017²⁶).

Projected changes in SWE_max indicate that reductions will be extensive (5% to 10% per decade through 2050, or a cumulative loss of 15% to 30% over the entire 2020–2050 period) over much of southern Canada, with the greatest changes in the Maritimes and British Columbia (see Figure 5.5). The decreases across the prairies, Ontario, Quebec, and the Maritimes are attributable to increasing temperatures that will shift the proportion of total precipitation that currently falls as snow toward rain (Sospedra-Alfonso and Merryfield, 2017¹⁶⁶). (Note that the greatest near-term reductions in SWE_max, according to the climate model projections, will be just south of the Canadian border.) Projected changes in British Columbia are consistent with projected SWE_max reductions in the Western Cordillera (Fyfe et al., 2017⁵⁰). While SWE_max is projected to increase by mid-century in the Eurasian Arctic (Brown et al., 2017²⁶), minimal change is projected across high-latitude land areas of Canada because increased snowfall is expected to be offset by increasing temperatures that shorten the snow accumulation season.

The greatest snow loss across Canada during the 2020–2050 period is projected to occur in the shoulder seasons (October–November and May–June; Thackeray et al., 2016¹⁷⁹) (see Figure 5.6). During mid-winter, there is minimal percentage change in projected snow cover extent because winter temperatures across northern regions of Canada will remain cold enough to sustain snow cover and there is greater climatological snow extent in winter, which results in smaller percentage changes (see Figure 5.5). The projected trends are similar to the rate of change already observed during the historical period (see Section 5.2.1). Trends from a large ensemble of simulations from the Canadian Earth System Model (CanESM2) are slightly stronger than the CMIP5 multi-model mean because projected warming is greater in CanESM2 than in the CMIP5 multi-model mean (Thackeray et al., 2016¹⁷⁹).

In summary, analysis of multiple sources of SCF data from satellite remote sensing and land surface models over the 1981–2015 period show the portion of the year with snow cover decreased across Canada at a rate of 5% to 10% per decade. There is very high confidence in these trends based on consistency among multiple datasets and quantitative relationships with surface temperature trends in which there is also high confidence (see Chapter 4). Seasonal snow accumulation decreased by a rate of 5% to 10% per decade across most of Canada (1981–2015), with the exception of southern Saskatchewan, Alberta, and British Columbia (increases of 2% to 5% per decade), driven by both temperature and precipitation changes. Because of greater uncertainty in sources of data on snow accumulation (compared to those on SCF), we have medium confidence in these trends. It is very likely that snow cover duration will decline to mid-century across Canada as a result of increases in surface air temperature under all emission scenarios. This likelihood assessment is based on the strongly established sensitivity of snow cover to surface temperature in both observations and climate models. Scenario-based differences in projected spring snow cover emerge by the end of the century, with stabilized snow loss for low and medium emission scenarios (RCP2.6 and RCP4.5) but continued snow loss under a high emission scenario (RCP8.5). A reduction of 5% to 10% per decade in seasonal snow accumulation (through 2050) is projected across much of southern Canada; only small changes in snow accumulation are projected across northern regions of Canada because increases in winter precipitation are expected to offset a shorter snow accumulation period. There is greater uncertainty in SWE projections (compared to SCD) because of greater spread in climate model responses due to the competing effects of temperature and precipitation, so there is medium confidence in these results.

Estimates of total ice and MYI area within Canadian Arctic waters are available from the Canadian Ice Service Digital Archive (CISDA), which is an integration of a variety of datasets, including satellite measurements, surface observations, airborne and ship reports, and operational model results (see Canadian Ice Service, 2007³⁰ and Tivy et al., 2011a¹⁸² for complete details). This record has been shown to provide more accurate estimates of sea ice concentration (SIC) in Canadian waters compared to satellite passive microwave estimates (Agnew and Howell, 2003²). Analysis of seasonally averaged trends in SIC over the 1981–2015 period (selected to match the period of snow datasets described in Section 5.2.1) found reductions over Canadian waters in all seasons (see Figure 5.2). Regions with the strongest SIC declines were eastern Canadian waters in winter and spring, and the CAA and Hudson Bay in summer and fall. SIC trend patterns are closely associated with warming patterns during the seasons of ice onset and growth (from October through March). However, dynamic effects (such as wind, which redistributes sea ice) also influence the observed ice reductions in spring and summer (Mudryk et al., 2018¹¹³).

The CISDA archive also extends the record of total and MYI back to 1968, almost 10 years earlier than coverage by satellite passive microwave observations. Between 1968 and 2016, sea ice area, averaged over the summer period, has decreased significantly in almost every region of the Canadian Arctic, by up to 20% per decade in some regions (e.g., the Hudson Strait and Labrador Sea; see Figure 5.7). Compared with trends computed over the periods 1968–2008 (Tivy et al., 2011a¹⁸²) and 1968–2010 (Derksen et al., 2012³⁶), more regions are now experiencing significant decreases, and the rate of decline is stronger in all regions except Hudson Bay. The largest declines in MYI have occurred in the CAA (approximately 9% per decade) and Beaufort Sea (approximately 7% per decade).

While there are high year-to-year differences due to natural variability, time series of sea ice area (see Figure 5.8) clearly show negative trends. The Beaufort Sea experienced record-low sea ice area in 2012, becoming virtually ice free near the end of the melt season (Figure 5.8a; Babb et al., 2016⁸). This was nearly repeated in 2016. The CAA had record-low ice years in 2011 and 2012, eclipsing the previous record set in 1998 (Figure 5.8b; Howell et al., 2013⁷⁵). Baffin Bay has experienced consistently low sea ice area since 1999 (Figure 5.8c), while Hudson Bay sea ice area has declined since the mid-1990s (see Figure 5.8d; Tivy et al., 2011b¹⁸³; Hochheim and Barber, 2014⁶⁹). Modelling has demonstrated that the recent extreme lows in Arctic SIC would not have occurred without anthropogenic climate change (see Box 5.1).

The decline of sea ice across the Canadian Arctic is driven by increasing spring air temperature and resulting increases in the length of the melt season. This results in more open water, increased absorption of solar radiation (which further contributes to ice melt), increased water temperature, and delayed fall freeze-up (Howell et al., 2009a⁷³; Tivy et al., 2011a¹⁸²; Stroeve et al., 2014¹⁶⁸; Parkinson, 2014¹²⁸). Changes in sea ice cover are also driven by atmospheric circulation. The Beaufort Sea was once a region where ice would thicken and age before being transported to the Chukchi Sea and recirculated in the Arctic (Tucker et al., 2001¹⁸⁵; Rigor et al., 2002¹⁴³), but now the region has become a considerable contributor to the Arctic’s MYI loss (Kwok and Cunningham, 2010⁹⁴; Maslanik et al., 2011¹⁰⁹; Krishfield et al., 2014⁹³; Galley et al., 2016⁵²). Ice is still being sequestered from the Canada Basin (one of the two ocean basins in the Arctic Ocean) and transported through the Beaufort Sea during the summer months, but this ice is now younger and thinner and melts en route to the Chukchi Sea (Howell et al., 2016a⁷⁰). The CAA was also a region with historically heavy MYI throughout the melt season, but MYI conditions have become lighter in recent years (see Figure 5.7; Howell et al., 2015⁷²).

Arctic sea ice thickness has declined in recent years, largely associated with a reduction and thinning of the MYI fraction (e.g., Kwok and Rothrock, 2009⁹⁶; Haas et al., 2010⁶⁶; Laxon et al., 2013¹⁰²; Richter-Menge and Farrell, 2013¹⁴²; Kwok and Cunningham, 2015⁹⁵; Tilling et al., 2015¹⁸¹). These studies indicate thickness declines are greater in the Beaufort Sea compared with the north-facing coast of the CAA, which still contains some of the thickest sea ice in the world (Haas and Howell, 2015⁶⁵). Unfortunately, the spaceborne sensors used to obtain sea ice thickness information over the Arctic Ocean are not of sufficient spatial resolution to provide thickness estimates within the narrow channels of the CAA. Although there are only four locations with consistent records and point measurements may not capture regionally representative conditions, the Canadian Ice Service record of in situ landfast ice thickness represents one of the longest datasets in the Arctic, spanning over five decades. Maximum ice thickness has declined significantly at three sites in the CAA (Cambridge Bay, Eureka, and Alert), with decreases ranging from 3.6 to 5.1 cm (± 1.7 cm) per decade from the late 1950s to 2016 (Howell et al., 2016b⁷⁴). No significant trend was found at Resolute, a result that differs from an earlier study by Brown and Cote (1992), which reported a significant increase in maximum ice thickness at Resolute over the 1950–1989 period.

Sea ice along the east coast of Canada is seasonal, with ice melting completely each spring. A robust indicator of change is winter season sea ice area, defined as the average from January through March of each year. The rate of decline between 1969 and 2016, determined from the CISDA for the entire east coast region, is 7.5% per decade (statistically significant at the 1% level; there is only a 1% possibility that the decline is due to chance; see Figure 5.10). This is consistent with the passive microwave time series, which indicates a decline of 9.5% per decade over the 1979–2015 period (Peng and Meier, 2017¹³⁰). There is regional variability within the east coast region, as the rate of decline for the Gulf of St. Lawrence (8.3% per decade) is less than that for eastern Newfoundland waters (10.6% per decade), while the decline for the southern Labrador Sea is not statistically significant at the 5% level (there is a possibility of more than 5% that the decline is due to chance; see Figure 5.10). Years with extensive ice cover are more prominent before 1995, but the region has experienced recent heavy ice years as well (in 2014 and 2015). Sea ice variability in this region is driven largely by temperature and atmospheric circulation (i.e., winds) associated with the Arctic Oscillation (also called the Northern Annular Mode [see Chapter 2, Box 2.5]; Deser and Tang, 2008³⁹; Peterson et al., 2015¹³³).

The narrow channels in Canadian Arctic waters are poorly represented by the coarse spatial resolution of climate models. While uncertainty in model projections is therefore higher for the CAA than for the pan-Arctic, evaluation of historical simulations shows the CMIP5 multi-model ensemble (see Chapter 3, Box 3.1) still provides a quantitative basis for projecting future sea ice conditions (Laliberté et al., 2016⁹⁷). Under a high emission scenario (RCP8.5), the CMIP5 multi-model projections indicate widespread reductions in SIC for the ice melt (summer) and ice formation (fall) seasons (Mudryk et al., 2018¹¹³; see Figure 5.4). For the east coast, virtually ice-free conditions during the winter months are projected by mid-century under a high emission scenario (RCP8.5) with uncertainty in these projections due to potential changes in the transport of sea ice from the Arctic to the east coast (Loder et al., 2015¹⁰⁴).

The probability and timing of future sea ice–free conditions are sensitive to the definition of ‘ice-free’ (Laliberté et al., 2016⁹⁷). When using a threshold of 5% ice area, there is a 50% probability that all Canadian regions will be sea ice–free in September by 2050 under a high emission scenario (RCP8.5; see Figure 5.11). The probability that all regions will be ice free is similar for August, but lower for October and November. Hudson Bay, which is already largely ice free in August and September, has a high probability of being ice free for four consecutive months (August through November) by 2050. Using a definition of 30% ice area, more persistent ice-free conditions are projected. Baffin Bay is projected to be ice free for August through October, and the Beaufort Sea and the CAA may be ice free in August and September by 2050.

The likelihood of summer ice-free conditions in the central Arctic is connected to the magnitude of projected global temperature increases, with a much greater probability of ice-free conditions for 2°C global warming compared to 1.5°C (Jahn, 2018⁷⁸; Sigmond et al., 2018¹⁵⁶). The area to the north of the CAA and Greenland will be the last refuge for summer sea ice (including MYI) in the Arctic during the summer (Wang and Overland, 2012¹⁹³; Laliberté et al., 2016⁹⁷), so ice will still drift into the Northwest Passage, where it will present a navigation hazard for shipping, even when the Arctic Ocean is sea ice–free during the summer (see FAQ 5.1: Where will the last ice area be in the Arctic?).

In summary, the Arctic sea ice environment has changed profoundly over recent decades (Barber et al., 2017¹⁰). Perennial sea ice that survives the summer melt is being replaced by thinner seasonal sea ice that melts in the summer. Summer sea ice area (particularly MYI) declined across the Canadian Arctic by 5% to 20% per decade (1968–2016, depending on region); winter sea ice area decreased in eastern Canada (by 7.6% per decade, 1969–2016). There is very high confidence in these trends, which are derived by trained analysts from Canadian Ice Service ice charts that provide region-specific sea ice information. It is very likely that continued reductions in summer and fall sea ice across the Canadian Arctic, and winter sea ice in eastern Canadian waters, will result from increased temperatures under all emission scenarios. Most Canadian Arctic marine regions could be sea ice–free for at least one month in the summer by 2050 (high confidence) based on simulations from CMIP5 models. There is very high confidence that the region to the north of the CAA and Greenland will be the last area where thick MYI will be present in the Arctic during the summer. Current understanding of sea ice dynamics, based on satellite observations, indicates this MYI will continue to drift into Canadian waters.

Climate warming, combined with periods of reduced precipitation in Western Canada, has contributed to total thinning of glaciers in the southern Cordillera by 30 to 50 m since the early 1980s (see Figure 5.13) (Zemp et al., 2015¹⁹⁷). By the mid-1980s, glaciers in Garibaldi Provincial Park, southern British Columbia, had contracted in area by 208 km² since the Little Ice Age maximum extent of 505 km², with accelerated shrinkage by another 52 km² (or 7% of the Little Ice Age maximum) by 2005 (Koch et al., 2009⁸⁴). Glacier extent at several sites in the central and southern Canadian Rocky Mountains decreased by approximately 40% from 1919 to 2006 (Tennant et al., 2012¹⁷⁸). Glaciers of the Columbia Icefield, in the Canadian Rocky Mountains, also experienced dramatic changes from 1919 to 2009, losing 22.5% of total area while retreating more than 1.1 km on average (Tennant and Menounos, 2013¹⁷⁷). Aerial photography shows that all glaciers in British Columbia’s Cariboo Mountains receded over the 1952–2005 period, with a loss of approximately 11% in surface area (Beedle et al., 2015¹⁵). In eastern Canada, small alpine glaciers in the Torngat Mountains, Labrador (see Figure 5.13 for location) shrunk by 27% between 1950 and 2005, with current thinning rates as high as 6 m per year across the 22 km² of glaciers that remain in this area (Barrand et al., 2017¹²).

Glaciers and ice fields covering approximately 10,000 km² of the Yukon have decreased in area by approximately 22% between 1957 and 2007, and thinned by 0.78 m water equivalent (90% uncertainty range 0.44 to 1.12 m per year) contributing 1.12 mm (90% uncertainty range 0.63 to 1.61 mm) to global sea level over this period (Barrand and Sharp, 2010¹¹). Mass balance of glaciers, measured at three monitoring sites in Alaska (all within 300 km of the Kluane ice field, Yukon) indicate a rapid change from positive to negative glacier mass balance in this region beginning in the late 1980s (Wolken et al., 2017¹⁹⁵).

Long-term in situ glacier monitoring indicates a trend of significant loss of mass for glaciers and ice caps in the CAA beginning in the early 1990s (see Figure 5.13). Acceleration of glacier thinning in this region in the mid-2000s coincided with increases in summer warming driven by the advection of warm air masses to the Arctic from more southerly latitudes (Sharp et al., 2011¹⁵⁴; Mortimer et al., 2016¹¹²). Based on satellite measurements and surface mass-budget models, total mass loss from glaciers and ice caps in the CAA has increased more than two-fold, from 22 gigatonnes (Gt) per year between 1995 and 2000 (Abdalati et al., 2004¹), to 60 Gt per year (90% uncertainty range 52 to 66 Gt per year) over the 2004–2009 period (Gardner et al., 2013⁵³), and 67 Gt per year (90% uncertainty range 61 to 73 Gt per year) over the 2003–2010 period (Jacob et al., 2012⁷⁷), with mass losses continuing to accelerate to 2015 (Harig and Simons, 2016⁶⁷). According to the most recent assessments of regional glacier change, glacier melting in the CAA has contributed 0.16 mm per year to global sea-level rise since 1995, 23% of the contribution of the Greenland ice sheet and 75% of the Antarctic Ice Sheet (Gardner et al., 2013⁵³; Shepard et al., 2012¹⁵⁵; Sharp et al., 2016¹⁵³).

The Barnes Ice Cap on Baffin Island, the last remnant of the Laurentide Ice Sheet that covered most of Canada during the last glaciation, lost 17% of its mass from 1900 to 2010 (Gilbert et al., 2016⁵⁵). Approximately 10% of the total area of ice in the CAA is composed of small, stagnant ice caps (the oldest are less than 3000 years old), located almost entirely under the regional equilibrium line altitude, meaning they do not have an accumulation zone and experience net thinning across their entire area in most years. These ice caps are shrinking rapidly (Serreze et al., 2017¹⁵²) and fragmenting (Burgess, 2017²⁸), with many expected to completely disappear within the next few decades. Of similar age to the small ice caps are the ice shelves of northern Ellesmere Island, which are composed of floating glacier ice and/or very thick old sea ice. These ice shelves have decreased in area by about 90% since 1900 (with more than 50% of that loss since 2003) and are expected to survive for only the next decade or two (Mueller et al., 2017¹¹⁶).

Like many glaciers in the world, Canada’s glaciers are out of equilibrium with current climatic conditions and will continue to lose mass for the foreseeable future. Summer warming in the Arctic has driven extreme melting of ice caps and glaciers over the past two decades, resulting in this region becoming the most significant cryosphere contributor to global sea-level rise after the Greenland and Antarctic ice sheets.

Climate model projections indicate that western Canada and the western United States together (grouped together in many studies because of their similar mountainous domain) could lose approximately 85% (90% uncertainty range 74% to 96%) of the 2006 volume of glaciers by the end of the century under a medium emission scenario (RCP4.5). Under a high emission scenario (RCP8.5), this loss could exceed 95% (Radic et al., 2014¹⁴⁰). Glaciers in the coastal ranges of western Canada are predicted to lose 75% (90% uncertainty range 65% to 85%) of their 2005 ice area and 70% (90% uncertainty range 60% to 80%) of their volume by 2100 based on the mean of four emission scenarios (RCP2.6, 4.5, 6.0, 8.5) (Clarke et al., 2015³²). Glaciers in the western Canadian interior are projected to lose more than 90% of the 2005 volume under all scenarios except a low emission scenario (RCP2.6) (Clarke et al., 2015³²). These changes, in combination with the projected loss of alpine snow cover, will impact regional water resources (Fyfe et al., 2017⁵⁰; see Chapter 6, Section 6.2). Glacier-fed rivers may experience periods of increased discharge due to greater meltwater contributions in a warmer climate, but this response is finite, and glacier mass loss associated with warming is projected to result in reduced summer streamflow by mid-century (Clarke et al., 2015³²). The rate and timing of this transition will have important consequences for stream and river water quality and temperature, and for the availability of water for human uses such as hydro-electricity generation and agriculture.

Regional land ice models project that glaciers and ice caps in the Canadian Arctic will lose 18% of their total mass by 2100 (Radic et al., 2014¹⁴⁰; relative to a baseline mass reference estimated by Radic and Hock, 2011) under a medium emission scenario (RCP4.5), equivalent to 35 mm of global sea-level rise (Lenaerts et al., 2013¹⁰³; Marzeion et al., 2012¹⁰⁸). This loss of land ice volume in Arctic Canada by 2100 will contribute 41 mm of sea-level equivalent (90% uncertainty range 26 to 56 mm) under RCP4.5, and 57 mm of sea-level equivalent (90% uncertainty range 39 to 75 mm) under a high emission scenario (RCP8.5) (Radic et al., 2014¹⁴⁰). Densification of high-elevation firn (partially compacted granular snow that is the intermediate stage between snow and glacial ice) has reduced or eliminated the internal storage capacity of the larger (more than 2000 km²) ice caps in this region, thus increasing their sensitivity to future warming (Noël et al., 2018¹²⁴). Based on the trajectories of observed loss over recent decades, many of the remaining small ice caps (less than 2000 km²) and ice shelves in the Canadian Arctic are expected to disappear by 2100.

In summary, Canada’s Arctic and alpine glaciers have thinned over the past three to five decades, due to increasing surface temperatures (very high confidence). While spatial sampling is sparse, these long-term trends in glacier thickness change have been measured annually following standardized protocols and agree with independent remote sensing and model-based approaches. Multiple assessments using satellite data and models show that mass loss from glaciers and ice caps in the Canadian Arctic represents the largest cryosphere contributor to global sea-level rise after the Greenland and Antarctic ice sheets (very high confidence). Based on a regional mass balance model forced by future climate scenarios, glaciers across the Western Cordillera are projected to lose up to 85% of their volume by the end of the century (high confidence). This will lead to a decline in glacial meltwater supply to rivers and streams, although there is only medium confidence in the absolute impacts on freshwater availability because of multiple other contributors to projected streamflow changes (see Chapter 6, Section 6.2). Based on output from various independent models, glaciers and ice caps in the Canadian Arctic will lose 18% of their total mass by the end of the century, and so will remain important contributors to global sea-level rise beyond 2100 (high confidence). Small ice caps and ice shelves in the Canadian Arctic are shrinking rapidly. Based on observed changes in recent decades, and ice cap and ice shelf sensitivity to projected temperature increases, most are expected to disappear well before 2100 (very high confidence).

Surface observations show that ice breakup is occurring earlier, and freeze onset later, across small lakes in southern Quebec, Ontario, Manitoba, and Saskatchewan (Brown and Duguay, 2010²²). A significant declining trend in annual maximum ice cover was observed for the Laurentian Great Lakes over the 1973–2010 period (71% decline for all of the Laurentian Great Lakes), with the largest declines occurring in Lake Ontario (88%), Lake Superior (79%), and Lake Michigan (77%) (Wang et al., 2012¹⁹¹). Heavy ice years in 2014, 2015, and 2018, however, result in no trend over the full 1973–2018 period (Figure 5.14). The large year-to-year variation is associated with the Arctic Oscillation/North Atlantic Oscillation (AO/NAO) and El Niño–Southern Oscillation (ENSO) (see Chapter 2, Box 2.5). For example, the record-breaking low in maximum ice cover in the winter of 2011/2012 occurred during a strong positive-phase AO/NAO and the cold phase of ENSO (La Niña event) (Bai et al., 2015⁹). Whether variable ice cover contributes to observed increases in water temperature in the Laurentian Great Lakes is a topic under debate. Recent findings suggest that changes in winter lake ice cover play only a minor role in the observed warming trend (Zhong et al., 2016²⁰²), whereas ice cover duration was linked to summer surface water temperature (particularly in nearshore areas) when the lakes were examined at a finer spatial scale (Mason et al., 2016¹¹⁰).

Satellite measurements show that lakes in Arctic Canada have also been experiencing an earlier ice minimum (the last date of floating ice cover on the lake surface) and an earlier date when the water is clear of ice (see Figure 5.15; see also Duguay et al., 2006⁴⁵; Prowse, 2012¹³⁵; Cooley and Pavelsky, 2016³⁴). These changes are consistent with a recent circumpolar assessment, which showed that approximately 80% of Arctic lakes experienced declines in ice cover duration from 2002 to 2015, due to both a later freeze-up and an earlier breakup (Du et al., 2017⁴¹). Results from northern Alaska (which are likely similar to those in northwestern Canada) show that lake ice has begun to thin in recent decades (Alexeev et al., 2016³). From 1992 to 2011, approximately one-third of shallow lakes in which the entire water volume historically froze to the bed by the end of winter had changed to floating ice (Arp et al., 2012⁷; Surdu et al., 2014¹⁷³). Canada’s northernmost lake, Ward Hunt Lake (located on Ward Hunt Island), had maintained stable, continuous year-round ice cover for many decades until very warm summers of 2011 and 2012, when the ice cover fully melted (Paquette et al., 2015¹²⁷). This loss of inland perennial freshwater ice cover occurred nearly simultaneously with the collapse of the nearby Ward Hunt ice shelf (Mueller et al., 2009¹¹⁷; Veillette et al., 2010¹⁸⁸). Analysis of a 15-year time series (1997–2011) of radar and optical satellite imagery provides further evidence that some lakes in the central and eastern Canadian high Arctic are transitioning from continuous (year-round) to seasonal ice cover (Surdu et al., 2016¹⁷⁴).

It is difficult to provide an assessment of river ice changes across Canada because of sparse observations and a lack of recent assessments of the available data. There is evidence of earlier river ice breakup, consistent with increases in surface temperature (Prowse, 2012¹³⁵). However, the impact that climate-driven changes in ice phenology and thickness, combined with changing seasonal flow regimes (see Chapter 6, Section 6.2) and the influence of hydraulic processes (i.e., changing ice strength), will have on ice jams and flood events is not fully understood (Beltaos and Prowse, 2009¹⁶).

Changes in lake ice can be projected only indirectly, because lake models are not embedded within global climate models and individual lakes are not spatially resolved. When forced by a future climate under a medium emission scenario (RCP4.5), lake ice models project that spring breakup will occur between 10 and 25 days earlier by mid-century (compared with 1961–1990), and freeze-up will be five to 15 days later across Canada (Brown and Duguay, 2011²³; Dibike et al., 2012⁴⁰) (Figure 5.16). This results in a reduction of ice cover duration of 15 to 40 days for much of the country. More extreme reductions of up to 60 days are projected in coastal regions. The range in projected changes is due to regional variability in temperature and snowfall changes, and to lake-specific variables such as size and depth. The Laurentian Great Lakes can be resolved by lake models if the projected climate forcing data are downscaled. This approach has identified consistent results, with reduced ice cover duration of between 25 to 50 days across the Laurentian Great Lakes by mid-century, due to both later freeze-up and earlier melt (Gula and Peltier, 2012⁶³). Mean seasonal maximum ice thickness is projected to decrease by 10 to 50 cm by mid-century, with a more pronounced decrease in the eastern Canadian high Arctic (Brown and Duguay, 2011²³).

Warming is projected to drive an earlier river ice breakup in spring, which is due to decreased mechanical ice strength and earlier onset of peak discharge (Cooley and Pavelsky, 2016³⁴). More frequent mid-winter breakup and associated ice jam events are anticipated (Beltaos and Prowse, 2009¹⁶), although projected changes in river ice properties may reduce ice obstructions during the passage of the spring freshet (the increased flow resulting from snow and ice melt in the spring) (Prowse et al., 2010¹³⁷). A shorter ice cover season and reduced ice thickness may affect food security for northern communities by reducing the reliability of traditional ice-based hunting routes and the safety of ice-based travel. The reliability and predictability of ice roads as supply lines to northern communities and development sites is not fully dependent on climate, because these ice roads are partially engineered each season (i.e., snow is removed to accelerate ice growth). However, there have been instances of severely curtailed ice-road shipping seasons due to unusually warm conditions in the early winter (Sturm et al., 2016¹⁷¹). The seasonal operational duration for such ice roads is expected to decrease as a result of winter warming (Perrin et al., 2015¹³²; Mullan et al., 2017¹¹⁸).

In summary, the duration of seasonal lake ice cover is declining across Canada due to later ice formation in fall and earlier spring breakup, with implications for freshwater ecosystem services, tourism and recreation, and transportation. Although the surface monitoring network is sparse, there is high confidence in this trend because of consistency between satellite observations and historical lake ice model simulations. There is a weak negative trend in seasonal maximum lake ice cover for the Laurentian Great Lakes (1971–2017); large year-to-year variation is the primary feature of the time series (very high confidence). Changes in lake ice are difficult to project because lake models are not embedded within global climate models and individual lakes are not spatially resolved. Instead, estimates of changing lake ice phenology are derived from lake ice models forced by projected future climates. These simulations indicate spring lake ice breakup will be 10 to 25 days earlier by mid-century, with a five to 15-day delay in fall freeze-up, depending on the emission scenario. While the impact of warming temperatures on ice phenology is clear, there is only medium confidence in these projections because of numerous sources of uncertainty, including the quality of snowfall projections, the limitations of lake ice models, and the role of lake-specific characteristics such as depth and morphology.

Permafrost conditions are challenging to monitor because they cannot be directly determined using satellite measurements. They are therefore determined largely from in situ monitoring, which results in gaps in the spatial distribution of measurement sites because of the relative inaccessibility of large portions of northern Canada and historical emphasis in monitoring regions with infrastructure development potential (such as the Mackenzie Valley; Smith et al., 2010¹⁶³). Changes in permafrost conditions over the last few decades can be assessed by tracking changes in two key indicators: permafrost temperature and thickness of the active layer. Ground temperature, measured below the depth where it varies from one season to the next, is a good indicator of decadal to century changes in climate, while the active layer responds to shorter-term climate fluctuations (Romanovsky et al., 2010¹⁴⁵).

Ground temperature is measured in boreholes, generally up to 20 m deep, across northern Canada. Some of these monitoring sites have been operating for more than two decades, while many others were installed during the International Polar Year (IPY, 2007–2009) to establish baseline measurements of the temperature of permafrost (Smith et al., 2010¹⁶³; Derksen et al., 2012³⁸). A comparison of data collected for about five years after the establishment of the IPY baseline indicates that permafrost has warmed at many sites from the boreal forest to the tundra (Smith et al., 2015a¹⁶⁰), with greater changes in the colder permafrost of the eastern and high Arctic, where temperatures increased by more than 0.5°C at some sites over this short time period. Continued data collection has extended the time series beyond 30 years for some sites, allowing researchers to place the changes since IPY in the context of a longer record.

The temperature of warm permafrost (above −2°C) in the central and southern Mackenzie Valley (i.e., Norman Wells, Wrigley) has increased since the mid-1980s, but the rate of temperature increase has generally been lower since 2000 — less than about 0.2°C per decade (see Figure 5.17 and Table 5.1). The low rate of increase is observed because permafrost temperatures are already close to 0°C in this region, so energy is directed toward the latent heat required to melt ground ice rather than raising the temperature further. In the Yukon, comparison of recent ground temperature measurements with those made in the late 1970s and early 1980s suggests similar warming of approximately 0.2°C per decade (Duguay, 2013⁴³; Smith et al., 2015b¹⁶¹). In contrast, in the northern Mackenzie Valley (sites designated Norris Ck and KC-07 in Figure 5.17 and Table 5.1), recent increases in permafrost temperature have been up to 0.9°C per decade, likely associated with the greater increases in surface air temperature in this region over the last decade when compared with the southern Mackenzie Valley (Wrigley, Norman Wells in Figure 5.17; Smith et al., 2017¹⁵⁹).

Since 2000, high Arctic permafrost temperatures have increased at higher rates than those observed in the sub-Arctic, ranging between 0.7°C and 0.9°C at 24 m depth and more than 1.0°C per decade at 15 m depth (see Table 5.1), consistent with greater increases in air temperature since 2000 (Smith et al., 2015a¹⁶⁰). Short records from sites in the Baffin region indicate warming at 10–15 m depth since 2000 (see Figure 5.17 and Table 5.1), but there has been a decline in permafrost temperatures since 2012 (Ednie and Smith, 2015⁴⁶) that likely reflects lower air temperatures in this region since 2010. In northern Quebec, where measurements at some sites began in the early 1990s, permafrost continues to warm at rates between 0.5°C to 1.0°C per decade (Smith et al., 2010¹⁶³; Allard et al., 2016⁴). Permafrost can exist at high elevations in more southerly locations. Canada’s most southerly occurrence of permafrost, at Mont Jacques-Cartier on the Gaspé Peninsula, shows an overall warming trend at 14 m depth of 0.2°C per decade since 1977 (Gray et al., 2017⁵⁸).

A network of thaw tubes throughout the Mackenzie Valley has provided information on trends in the active layer thickness (ALT) between 1991 and 2016 (see Figure 5.18; Smith et al., 2009¹⁶⁴). ALT exhibits greater variability among years than does deeper ground temperature, with higher values of ALT in extremely warm years such as 1998 (Duchesne et al., 2015⁴²). ALT generally increased between 1991 and 1998 but decreased over the following decade in response to lower annual air temperatures in the region. Since 2008, there has been a general increase in ALT in Mackenzie Valley, with peak values in 2012 (Duchesne et al., 2015⁴²; Smith et al., 2017¹⁵⁹). At sites where the permafrost is ice-rich, increases in summer thawing have been accompanied by significant settlement (subsidence) of the ground surface (Duchesne et al., 2015⁴²).

A number of recent studies provide other evidence of changing permafrost conditions. Observations of landscape change over time, often based on air photo or satellite imagery interpretation, have identified areas undergoing thermokarst processes, such as lake formation and collapse of peat plateaus and palsas (e.g., Olefeldt et al., 2016¹²⁶; Kokelj and Jorgenson, 2013⁸⁶). Over the last 50 years in northern Quebec, there has been a loss of permafrost mounds, collapse of lithalsas, and increases in the size of thermokarst ponds (Bouchard et al., 2014¹⁹; Beck et al., 2015¹⁴; Jolivel and Allard, 2017⁸⁰), while palsa decay has been observed in the Mackenzie mountains of the Northwest Territories (Mamet et al., 2017¹⁰⁵). A recent repeat of a 1964 survey of permafrost conditions along the Alaska Highway corridor between Whitehorse and Fort St. John indicated that permafrost continues to persist in organic-rich soils, but is no longer found at other sites (James et al., 2013⁷⁹). Changes in lake area in Old Crow Flats since 1951 have also been linked to thermokarst processes (Lantz and Turner, 2015¹⁰⁰). A recent intensification of thaw slumping may also be tied to changes in climate, including increases in precipitation (Kokelj et al., 2015⁹⁰, 2017a⁸⁸; Segal et al., 2016¹⁵⁰; Rudy et al., 2017¹⁴⁸). In the southern Northwest Territories, forest die-off has been attributed to permafrost thawing and ground subsidence (Sniderhan and Baltzer, 2016¹⁶⁵). Erosion of Arctic coasts in the form of retrogressive thaw slumps can result from a combination of mechanical (wave action) and thermal (warming permafrost) processes, potentially exacerbated by sea-level rise (see Chapter 7, Section 7.5; Ford et al., 2016⁴⁸; Lamoureux et al., 2015⁹⁸; Lantuit and Pollard, 2008⁹⁹).

Climate models project large increases in mean surface temperature (approximately 8°C) across present-day permafrost areas by the end of the 21st century under a high emission scenario (RCP8.5) (Koven et al., 2013⁹¹) (see Chapter 3, Section 3.3.3). While this dramatic warming will no doubt affect permafrost temperatures and conditions (e.g., Slater and Lawrence, 2013¹⁵⁷; Guo and Wang, 2016⁶⁴; Chadburn et al., 2017³¹), it is challenging to project associated reductions in permafrost extent from climate model simulations because of inadequate representation of soil properties (including ice content) and uncertainties in understanding the response of deep permafrost (which can exceed hundreds of metres below the surface). Simulations from a model considering deeper permafrost and driven by low and medium emission scenarios project that the area underlain by permafrost in Canada will decline by approximately 16%–20% by 2090, relative to a 1990 baseline (Zhang et al., 2008a¹⁹⁹). These declines are smaller than projections from other modelling studies that only examined near-surface ground temperature (Koven et al., 2013⁹¹; Slater and Lawrence, 2013¹⁵⁷). These simulations also show that permafrost thaw would continue through the late 21st century, even if air temperatures stabilize by mid-century (Zhang et al., 2008b²⁰⁰).

Other climate-related effects also influence the future response of permafrost to warming and complicate modelling of future conditions (e.g., Kokelj et al., 2017b⁸⁹; Romanovsky et al., 2017a¹⁴⁴). For example, intensification of rainfall appears to be strongly linked to thaw slumping (Kokelj et al., 2015⁹⁰). New shrub growth in the tundra can promote snow accumulation and lead to warmer winter ground conditions (Lantz et al., 2013¹⁰¹). Thaw and collapse of peat plateaus and palsas into adjacent ponds increase overall permafrost degradation, and gullies that form because of degrading ice wedges can result in thermal erosion and further permafrost degradation (Mamet et al., 2017¹⁰⁵; Beck et al., 2015¹⁴; Quinton and Baltzer, 2013¹³⁸; Godin et al., 2016⁵⁷; Perreault et al., 2017¹³¹). Damage to vegetation and the organic layer due to wildfires (which are projected to occur more frequently under a warming climate) can lead to warming of the ground, increases in ALT, and degradation of permafrost (Smith et al., 2015c¹⁶²; Zhang et al., 2015²⁰¹; Fisher et al., 2016⁴⁷). Similarly, vegetation clearing and surface disturbance due to human activity and infrastructure construction can also lead to ground warming and thawing and enhance the effects of changing air temperatures on permafrost environments (Smith and Riseborough, 2010¹⁵⁸; Wolfe et al., 2015¹⁹⁴).

In summary, a large proportion of the northern Canadian landscape has undergone, or will soon undergo, changes brought about by permafrost thaw. The temperature of permafrost is increasing across sub-Arctic and Arctic Canada, and the ALT has increased over the past decade in the Mackenzie Valley (high confidence). The rate of increase differs within and among regions due to variability in surface temperature changes, soil properties, and preceding temperature conditions. There is high confidence in these trends: they are derived from high-quality borehole measurements (permafrost temperature) and thaw tube networks (ALT), although the measurements are spatially sparse. The observed permafrost temperature and ALT changes are consistent with regional surface air temperature trends, but additional factors such as snow cover, vegetation changes, and disturbance can also modulate the response of permafrost to a changing climate. Ongoing landscape change across northern Canada associated with the expansion of thermokarst landforms was identified from surface observations and remote sensing. There is medium confidence in the assessment of thermokarst changes: they are associated with well-understood processes associated with the thaw of ice-rich permafrost, but region-specific rates of change are difficult to determine. Projected increases in mean air temperature over land underlain with permafrost under all emission scenarios will result in permafrost warming and thawing across large areas of Canada by the middle of the 21st century (high confidence), with impacts on northern infrastructure and the carbon cycle. This is the expected permafrost response to the high probability of increased surface temperature across Arctic land areas, with additional factors such as changes to snow cover, surface wetness, vegetation, and disturbance also influencing permafrost conditions. Confidence in projected permafrost changes from climate model simulations is affected by inadequate representation of soil properties (including ice content) and uncertainties in understanding the response of deep permafrost.