Landfast sea ice in Hudson Bay and James Bay: Annual cycle, variability and trends, 2000–2019

Landfast sea ice (hereinafter fast ice) refers to sea ice that is attached to the coast or seafloor and remains largely immobile. Fast ice occurs in marine and freshwater environments alike (Granskog et al., 2003; Boone, 2010), and is commonly described by its extent (the distance from the coast to the fast-ice edge), duration (the amount of time fast ice is present) and thickness. Its stability and therefore extent depends on various factors, including the seasonal ice cycle, local tidal regime, bathymetry/coastal geometry, seafloor morphology and ice thickness, specifically keel depth as grounded keels contribute to the stabilization of fast ice (Reimnitz and Maurer, 1978; Jones et al., 2016). Atmospheric conditions and pack ice drift regimes are also noted to be controlling factors for fast-ice extent, stability and growth (Mahoney et al., 2014). In the Arctic, the coastal band of fast ice provides a critical travel route and hunting platform for local communities (George et al., 2004; Laidre et al., 2015; Lovvorn et al., 2018; Babb et al., 2022).

Shoals, islands and grounded icebergs can also increase fast-ice stability and persistence, and allow the fast ice to extend further offshore (Eicken et al., 2005; Kirillov et al., 2021). Mahoney et al. (2014) compiled a list of the typical water depth along the fast-ice edge in various regions and stated the following, “Among the different water depths that have been cited as the limits of landfast ice extent are 25 m along the Siberian coast (Zubov, 1945; Barry et al., 1979), 10 m in the Kara Sea (Barry et al., 1979; Divine et al., 2004), between 18 m and 30 m in the Beaufort Sea (Canadian Hydrographic Service, 1968; Kovacs and Mellor, 1974; Reimnitz and Barnes, 1974; Stringer, 1974; Shapiro, 1975; Kovacs, 1976; Weeks et al., 1977; Stringer et al., 1980; Mahoney et al., 2007), 100 m near Severnaya Zemlya (Divine et al., 2004), and 180 m off the eastern coast of Baffin Island (Jacobs et al., 1975)”. However, the latter example was due to grounded icebergs, which allowed fast ice to stabilize at depths beyond the typical depth of grounded keels. The extension of the fast-ice edge further offshore as a result of grounded icebergs is a typical feature of the Antarctic landfast ice domain, where icebergs may permit the fast-ice edge to extend to water depths of up to 400 m (Fraser et al., 2021).

Supported by the complex coastal features and offshore islands, fast ice can sometimes form a so-called ‘ice bridge’, which is a continuous sheet of ice that crosses deeper bays or channels, or joins offshore islands to the mainland (e.g., Larouche and Galbraith, 1989; Melling, 2002; Kirillov et al., 2021; Babb et al., 2022). Ice bridges serve as important platforms for hunting and transportation for the local communities, and promote the formation of latent heat polynyas along their leeward edge (Hannah et al, 2009; Itkin et al., 2015).

The presence of fast ice has a remarkable influence on the overall Arctic and sub-Arctic coastal environment. Fast ice impacts freshwater dispersion and ocean stratification in the nearshore coastal environment (Dmitrenko et al., 2005). In the Arctic, the formation of fast ice in deltas buffers river discharge from flowing into the ocean (Prinsenberg et al., 2008). In front of large estuaries, fast ice prevents wind-influenced spreading and mixing of the river outflow, and thereby retains a large volume of river freshwater beneath the ice, forming a well stratified upper layer restricted in the nearshore zone (Ingram and Larouche, 1987; Granskog et al., 2005; Itkin et al., 2014; Kasper and Weingartner, 2015), which is often called an under-ice lake (Hutchinson and Löffler, 1956). Presence of grounded ice structures can also alter the distribution of freshwater in the nearshore coastal domain (Are and Reimnitz, 2000). Due to the presence of fast ice, additional friction is introduced at the water-ice interface and hence to the water column which reduces tidal amplitudes and velocity (Wang et al., 2012), as well as hindering surface layer mixing by blocking air-water interaction (Proshutinsky et al., 2007).

According to Taylor and McCann (1983), close to 90% of the Canadian coast experiences sea ice interactions that alter the local geomorphology and beach type. While ice remains attached to the land, its interaction with the coast is mainly protective in nature (Forbes and Taylor, 1994). Formation of bottom-fast ice protects the coastline against waves or tidal movements. However, before fast-ice stabilization and after break-up, ice remains mobile due to the action of wind and waves. Interactions with the mobile ice can be both constructive and destructive in nature (Dyke and Morris, 1990). Sediment carried by ice is deposited in the nearshore zone (Reimnitz and Barnes, 1987; Dionne, 1988; Reimnitz et al., 1990; Reimnitz et al., 1991), and during break-up season tidally driven movement of sea ice results in ice driven towards the coast, pushing sediment back up the coast as a result of collision (Taylor and Forbes, 1987). Elevated features like beach ridges are formed as a result of such interactions. Destructive effects of ice can be seen by observing formation of ice scours, nearshore grounding or wallows (Forbes and Taylor, 1994). Fast ice also offers a habitat to large marine mammals (ringed seals and polar bears) as well as ice-algal communities, which are an essential part of the under-ice ecosystem (Horner and Schrader, 1982; Mundy et al., 2005; Gradinger et al., 2009; Else et al., 2019). Another implication of fast ice on the human dimension is its impact on shipping. The presence of fast ice may impede the operations of ships, which arises as a problem for maritime activities related to community re-supply and mineral transportation (Gavrilchuk and Lesage, 2014; Babb et al., 2019).

Hudson Bay and James Bay are inland seas in Canada that represent the most southerly extension of Arctic-like oceanic and sea ice conditions (Martini, 1986). With cold continental winters but relatively warm summers compared to the High Arctic, the duration of the seasonal sea ice period in Hudson Bay and James Bay varies spatially from 5 to 10 months every year (Hochheim and Barber, 2014). Generally, sea ice freeze-up begins in November in the northwest and progresses across Hudson Bay with southeastern Hudson Bay and James Bay being the last area to become ice-covered (Andrews et al., 2018). Ice typically forms in shallow coastal areas first, and then extends offshore. In terms of sea ice break-up, it begins around late May to early June in northwestern Hudson Bay when new ice ceases to form in the Kivalliq polynya, although fast ice remains intact further into the melt season. The last remnants of sea ice are typically found in southern or sometimes eastern Hudson Bay and melt out in early August (Gagnon and Gough, 2006). In terms of ice thickness, community-based observations from level fast ice revealed that it was thicker in the northwest (189 cm) and northeast (237 cm), and thinner in southeastern Hudson Bay (140 cm) and in James Bay (93 cm; Gagnon and Gough, 2006). The offshore pack ice in Hudson Bay is characterized by a pronounced west-to-east gradient as a result of westerly winds that maintain the Kivalliq polynya in northwestern Hudson Bay (Bruneau et al., 2021), and dynamic sea ice growth due to convergence in eastern Hudson Bay (Landy et al., 2017; Kirillov et al., 2020). Additionally, large tides within Hudson Bay and James Bay drive a dynamic coastal flaw lead system along the edge of the fast-ice cover where new ice forms and is subsequently deformed on a semidiurnal cycle (Barber et al., 2021). This process has been observed to create heavily deformed, rough and thick ice floes within southern Hudson Bay, where they may be grounded and entrained within the fast ice and stabilize its offshore extension, or enter the ice pack and persist late into the melt season (Barber et al., 2021).

Owing to the significant role of fast ice in the coastal environment, the monitoring of changing fast-ice regimes and the processes linked to its stability, growth and decay is crucial. Yu et al. (2014) observed that the duration of fast ice across the Northern Hemisphere decreased at an average rate of 8 days/decade between 1976 and 2007. Some areas experience a more drastic shift in the fast-ice duration, almost 3 days/year, observed in the Laptev Sea (Selyuzhenok et al., 2015), with similar trends also observed in the Canadian Arctic Archipelago (Galley et al., 2012). Throughout the Arctic and sub-Arctic, fast ice has been observed to be decreasing in duration and extent; however, the observations vary regionally. Still, fast-ice variability across the sub-Arctic Hudson Bay coastal zone remains sparsely studied and understood, largely due to its complex nature, high regional variability, limited international interest and technological and logistical challenges associated with the harsh environment. The objectives of this study were to: i) examine the annual fast-ice cycle in Hudson Bay and James Bay; ii) investigate the inter-annual trends of fast-ice duration from 2000 to 2019; iii) quantify the variability of fast-ice extent and coverage across the study area, and iv) explore the role of air temperature and coastal topography on the growth and stability of fast ice.

Hudson Bay and James Bay cover approximately 800,000 km² and represent the largest inland marine waterbody in North America (Dionne, 1980). Geographically, Hudson Bay and James Bay are situated centrally within the Canadian Shield. Based on the coastline defined in the Canadian Ice Service ice charts, there are approximately 9,900 km of continuous coastline around Hudson Bay and James Bay. To analyze the spatial variability of fast ice along this extensive coast we divided Hudson and James Bay coastline into eight divisions (Figure 1): A) northern Hudson Bay, B) northwestern Hudson Bay, C) southwestern Hudson Bay, D) southern Hudson Bay, E) western James Bay, F) eastern James Bay, G) southeastern Hudson Bay, and F) northeastern Hudson Bay. Given that islands are not part of this continuous coastline, we have chosen to exclude them from these divisions and only consider them part of the fast-ice extent if an ice bridge connects the islands to the mainland. The divisions were defined manually, based on a number of aspects such as similarity in patterns of fast-ice variability within regions, coastline orientation, nearshore topography and latitudinal differences.

Due to the direct impact of fast ice variability and dynamics on the coastal communities and the importance of this knowledge to local peoples, a closer look at the community level is needed. Hence, a more detailed analysis of fast-ice conditions was performed near 15 coastal locations/communities across Hudson Bay and James Bay, specifically Chesterfield Inlet, Rankin Inlet, Arviat, Churchill, Cape Tatnum (near York Factory), Fort Severn, Peawanuck, Attawapiskat, Moosonee, Chisasibi, Sanikiluaq, Inukjuak, Akulivik, Ivujivik and Coral Harbour (Figure 1). All of the selected study sites are adjacent to the coastal communities and contain some level of transportation infrastructure, as well as represent different coastal orientations.

The primary datasets used in this study were ice charts from the Canadian Ice Service Digital Archives (CISDA). These ice charts integrate observations from various satellite sensors, aerial reconnaissance, ship-based observations, operational models and observational prowess of experienced ice forecasters to provide information on ice concentration and stages of development (Tivy et al., 2011; Galley et al., 2012). The CISDA provides information of ice conditions in these regions from the early 1960s to the present. Technological advancements have greatly influenced the temporal availability and quality of these observations. Tivy et al. (2011) showed the data quality of the CISDA ice charts by spatiotemporal quality indices, indicating a high confidence of the data quality after the launch of Radarsat-1 in 1996, compared to an average confidence level during the years prior to 1996. A limitation of the CISDA is the temporal availability of ice charts. Ice charts have historically been produced weekly during the summer shipping season, but their frequency during winter has varied. During winter ice charts were produced monthly from 1996 to 2006, bi-weekly from 2007 to 2011 and weekly since 2012. Apart from the CISDA ice charts, this study also includes the use of optical imagery from MODIS, which was first launched in 2000. Hence our study period runs from 2000 to 2019. Satellite imagery from MODIS is provided daily which was important for our study as it helped us to pinpoint the dates.

Ice charts delineate different ice regimes with polygons that present the total concentration and partial concentration of up to three different stages of development according to the World Meteorological Organizations (WMO, 1970) egg code. Fast ice is denoted as an individual class in a defined polygon, characterized by a 10/10ths concentration and being contiguous or attached to the coastline. Ice charts were acquired in E00 format and converted to shape files for further analysis.

This section involves a high-resolution analysis of fast-ice time-related events, i.e. freeze-up/break-up days and ice persistence. Instead of broad divisions, this part of the study was carried out at 15 selected locations across Hudson Bay and James Bay. Determination of the fast ice freeze-up and break-up dates have been explained in a variety of ways by previous researchers. Most of these definitions and methods vary in their descriptions, suited well to the aims of their research, geographical settings and availability of datasets. Most of their methods were based on the definition of fast ice that focuses on no detectable motion over a specified time, e.g. 1–4 days of no detectable motion (Stringer et al., 1978; 1980; Barry, 1979) or 20 days of no detectable motion (Mahoney et al., 2007). For this study, we relied on the definition of the fast-ice polygon in the ice charts, following the approach taken by Galley et al. (2012), where the authors considered polygons representing a 10/10ths concentration of ice as fast ice, instead of using only such polygons designated as fast ice by the CISDA. For maintaining accuracy, the selected polygons were further checked to ensure the polygons were continuous and attached to the shore.

Given that the ice charts were produced weekly, bi-weekly or monthly, the exact date of fast ice freeze-up cannot be determined, hence the week when fast ice is first present is treated as the week of freeze-up, while the week when fast ice was first absent is treated as the week of fast-ice break-up. In some instances, multiple fast-ice freeze-up and break-up events were observed within the 15 locations. Suited to our research goal, we considered the week with the first freeze-up event as a start period and the week with the last break-up event as the end period. As most of the ice charts are produced at an interval of 7 days, estimating the exact day of freeze-up or break-up is often inaccurate. Instead it tells us more precisely about the week these events took place. Because in some cases the ice charts were more than a week apart (see Section 2.2), here we used the charts to identify the week when these events took place by identifying two consecutive charts depicting a fast-ice-present or fast-ice-absent scenario.

Once the week of freeze-up or break-up was determined, daily MODIS and VIIRS imagery from the NASA Worldview observation platform (see Section 2.5) was used as a tool to further identify the exact day of these events. These observations were cross-checked with Landsat satellite imagery available during the desired timeline, wherever possible. For this analysis, the stretch of 50 km of the coastline adjacent to the selected 15 study locations was considered. The first day of the ice season with complete presence/absence of fast ice along the entire 50 km stretch was considered as the freeze-up/break-up day. The use of optical remote sensing was often limited by the presence of clouds. When the study locations were obscured by persistent cloud coverage, we considered the first cloud-free day as the freeze-up/break-up day. While in most cases daily observations were required for only seven consecutive days (being weekly apart), some ice charts (before 2006) placed almost a month apart would require daily observations of thirty consecutive days. In cases where all seven consecutive days within the weekly images were not useful for determining the dates, we used the date of the weekly ice charts showcasing the first freeze-up/break-up events for the study. In some cases, detecting changes in some fast-ice properties through thinner clouds was possible.

After determining the days of freeze-up and break-up, we estimated the decadal trends in the timing of break-up and freeze-up. Fast-ice duration at each location was also determined by the number of days between freeze-up and break-up for the subsequent year. Due to the intense variations in the inter-annual records and no assumed relationship between freeze-up/break-up days and years, we opted for a non parametric test (Mann–Kendall’s Test) to examine the significance of the trend in the observed datasets (Mann, 1945; Kendall, 1948), where the magnitude of the trend (positive or negative) was estimated using Sen’s Slope (Sen, 1968).

Fast-ice polygons from the ice charts were used to determine the seasonal development of the fast-ice area over the 2000–2019 time period. The mean fast-ice area was calculated for each month and across each ice season, and the maximum fast-ice area determined for each ice season. Additionally, the cumulated annual fast-ice cover (CAFIC) was calculated for each ice season by merging all fast-ice polygons and determining the total area covered by fast ice at some point during each ice season. The CAFIC differs from the maximum fast-ice area as it does not represent the maximum fast-ice area within a given ice chart, but rather the total area covered by fast ice during each ice season even if the fast ice was only present for a short period.

Furthermore, fast-ice polygons were converted to a polyline feature, and the seaward edge of these lines was used to represent the fast-ice edge. Due to the vast coastline of the Hudson Bay and James Bay combined (roughly 9900 km), we included a control point every 100 km of the coastline starting from Ivujivik and running along the coastline in a clockwise direction to the southern end of Wager Bay in the Roes Welcome Sound. The coastline provided by the Canadian Ice Service in their ice charts was used for this section. The control points were generated using a spatial analysis tool in ArcGIS to generate a point vector on the shapefile of the coastline every 100 km. The points appeared denser along complex coastlines as compared to straight open coastlines. Apart from this effect, these control points cover all coastline types, as seen later in the text (Section 4.4). This method for generating control points does result in an uneven distribution of points with respect to zones, but from a bay-wide perspective this approach helps by selecting unbiased control points and maintaining parity throughout the entire Hudson Bay and James Bay coastline. Vectors originating at these control points were drawn perpendicular to the coastline, with the point of intersection of vector with seaward ice edge noted as the fast-ice extent in kilometers (Figure S1). This process was applied to the mean yearly extent of fast ice and later compared inter-annually. Working with the fast-ice polygons from the ice charts also involved estimating the percent occurrence of fast ice in the Hudson Bay and James Bay over the 2000–2019-time period.

Other datasets include daily observations from NASA Worldview, an interactive viewing platform for daily optical satellite observations. MODIS and VIIRS corrected reflectance images were used for supporting observations of fast-ice cover during cloud-free sky conditions. These images are available in true-colour composites with 250-m spatial resolution. Satellite observations from different Landsat missions were also used in this study as a support dataset. Landsat images have a spatial resolution of 30 m and a temporal resolution of 16 days. A similar red-green-blue band combination was used from the Landsat images to aid this study. Given the long revisit interval of the Landsat images, they were used whenever the determined freeze-up/break-up days coincided with the available imaging days of Landsat to cross-validate the observations from MODIS and VIIRS.

Several multispectral images from Landsat were collected over the Hudson Bay and James Bay. Images were selected based on maximum visualization (cloud-free conditions). Multiple images were mosaicked to obtain a complete view of the selected study sites, and to delineate coastline and nearshore coastal features. The General Bathymetric Chart of the Oceans (GEBCO) bathymetry products were used to generate a bathymetry map over the entire study area (GEBCO, 2003). GEBCO provides a continuous, reviewed, processed and gridded database of bathymetry records collected by scientists and institutions worldwide, which is compiled and made accessible at a 15-arcsecond resolution. Further integrating these datasets with the fast-ice extent and duration, we mapped the depth where the fast-ice edge appears more frequently and reaches a stable phase. Similarly, we also noted the duration of ice in different coastal regions to observe the influence of coastline orientation and topography on stabilizing or sheltering fast ice, relative to other coastal types.

Daily mean air temperatures at 2 m above the surface were acquired from the MERRA-2 reanalysis for the 2000–2019 study period at a spatial resolution of 0.5° x 0.625° (GMAO, 2015). Luo et al. (2020) conducted an extensive study to examine the accuracy of MERRA-2 reanalysis products of sea surface skin temperature, atmospheric temperature and humidity. Based on the results the authors supported the use of MERRA-2 reanalysis products across various research fields. From the 2000–2019 temperature record at each location, we identify the fall transition day (FTD) as the last day of each fall season with the daily mean temperature above 0°C, and the spring transition day (STD) as the first day of each year with the daily mean temperature above 0°C. Occasionally anomalies may arise during the winter months with a day with above-freezing temperatures. To avoid classifying such days as FTD, we excluded such days if they were placed 7 days apart from the last recorded 0°C temperature. Similar exclusions were performed for spring months as well. In our analysis these anomalies were rare (11 times out of 570 observations), however, and did not occur consecutively for every year. Furthermore, these data were combined with the observations of freeze-up and break-up days to investigate the influence of air temperature on ice growth and decay.

Records of fast-ice freeze-up from the 15 study locations around Hudson Bay and James Bay indicate a northwest-to-east pattern of freeze-up across the study area (Figure 2a). Fast ice generally forms first in Chesterfield Inlet in early to mid-November, before advancing southward along the western shore through November and into early December. During December, fast ice formed at the remaining study sites and formed last in Ivujivik in northeastern Hudson Bay. In general, this timing is consistent with the bay-wide ice growth pattern in offshore waters (Stewart and Barber, 2010). The difference between the first and last records of fast-ice freeze-up observed across Hudson Bay and James Bay was found to be 47 days over the 2000–2019 study period (Figure 2a).

The progression of fast-ice break-up events across Hudson Bay and James Bay (Figure 2b) follows a south-to-north trend with fast ice breaking up as early as mid-May in southern James Bay and gradually progressing northwards, with the final break-up events taking place in the northwestern part of Hudson Bay around mid- to late July. An approximately 50-day difference between the earliest (southern James Bay) and latest (northern Hudson Bay) break-up events was observed across the Hudson Bay and James Bay (Figure 2b).

The annual average of fast-ice duration across Hudson Bay and James Bay was observed to be 185 ± 10 days (Figure 2c), and varied from a minimum of 150 days (approximately 5 months) in Moosonee to a maximum of 234 days (approximately 7 months) in Rankin Inlet (Figure 2c). The fast-ice cover remained immobile throughout April; decay started by the beginning of May. By the middle of June, southern and eastern Hudson Bay and the entire James Bay reaches an ice-free condition, with some parts of the northern and western Hudson Bay still retaining fast ice. We observed a minimum fast-ice duration of 143 days (4.7 months) at Moosonee in the southeast, and a maximum duration of approximately 233 days (7.7 months) near Rankin Inlet in the west. Rankin Inlet is the location within the Hudson Bay and James Bay system where fast ice generally appears first and its presence is most persistent.

The interannual trends in fast ice duration over the 2000–2019 period (Figure 2d) show that the fast-ice duration along the west coast of Hudson Bay has been decreasing at a rate of 1–6 days per decade. All other locations across Hudson Bay and James Bay, with the exception of Moosonee, have experienced increasing fast-ice duration. This increase has been particularly notable in Sanikiluaq and Chisasibi, where fast-ice duration increased at 9 and 6 days per decade, respectively (Figure 2d).

The seasonal change in fast-ice area (in square kilometers) and its inter-annual variability from 2000–2001 to 2018–2019 is shown in Figure 3, while Figure 4 summarizes the CAFIC in each of the eight divisions for each year. The 2014–2015 ice season exhibited the highest fast-ice area from February to May (Figure 3). From initial formation in November, fast-ice area steadily rises to a peak in March. The mean peak in fast-ice area is 88,517 km², with a range of 20,688 km². After March fast-ice area begins to decline, with a particularly noteworthy decline from May to June before there is no fast ice remaining in July. The CAFIC of 162,125 km² across the full Hudson Bay and James Bay area was also the largest for the study period (Figure 4). This large area is explained by the size of the ice bridges that formed between the mainland and Ottawa Islands, Belcher Island and Charlton Island. By contrast, the lowest yearly mean fast-ice cover over the study period occurred in 2009–2010 (84,274 km²), when none of the ice bridges formed.

This observation also provides an understanding of the fast-ice growth and decay pattern in an ice season. This analysis does not account for anomalous rare events like delayed ice freeze-up, earlier break-up or “false” ice freeze-up and break-up events described by Rolph et al. (2018) as temporary freeze-up or break-up of ice prior to a determined event of freeze-up or break-up. Instead it provides sufficient knowledge of only the mean area covered by fast ice on a monthly time scale.

In Figure 5, CAFIC is expressed within the range 0–100%, with 0% being the seaward edge of the maximum limit of fast ice recorded in the 2000–2019 period and 100% being the area that became fast-ice covered in all 19 years. The 100% occurrence line is used to represent the zone of continuous reoccurrence where the fast ice reoccurred in every ice season. Note that this 100% limit does not cover the ice bridges as they do not form every year; however, evidence suggests that they may have occurred much more frequently in the past (Flaherty, 1918). Figure S2 provides a visual representation of the inter-annual variability of the maximum extent of the fast ice cover for each year from 2000 to 2019. The variations in the extents associated with ice bridge formation between the mainland and islands are seen Figure S2a for Roes Welcome Sound, Figure S2e for southeastern James Bay, and Figure S2f for Belcher Islands and Ottawa Islands. These variations are discussed further in Section 4.4.

The annual cycle of fast ice in Hudson Bay and James Bay can be divided into four stages that classify the growth, stability and decay of ice (Figure S3). The first is the ‘growth phase’, when the ice starts to freeze in the nearshore zone as a thin sheet of ice, undergoes thermodynamic freezing and eventually attains stability. This phase usually starts from November and the gradual extension of the fast ice continues until March (Figure 3). During this time the fast ice reaches its maximum extent and remains stable, which is the second phase representing ‘maximum fast-ice period’, when the fast-ice area remains relatively stable until early May. Past this period there is a gradual decline in the fast-ice area as the ice starts breaking up, and by late June major parts of the Hudson Bay and James Bay coastlines are free of fast ice (Figure 3). By mid-July, typically no fast ice remains in Hudson Bay and James Bay, though mobile ice floes may still be present in the nearshore zone. The remnant floes either melt locally or drift offshore, where they subsequently melt. This period between fast-ice break-up and complete ice loss is described as the ‘post break-up phase’. Finally, by July–August the remaining mobile ice floes melt away initiating the final phase, ‘open-water season’.

Within Hudson Bay and James Bay there is substantial spatial variability in the pattern of fast-ice freeze-up and break-up, and therefore the pattern of fast-ice duration. The variability in freeze-up and break-up and the maximum extent of fast ice arise due to the different forcing mechanisms that dictate how fast ice forms, extends and subsequently breaks up. Hence, within this discussion section, we examine the contribution of air temperature, surface winds, bathymetry and coastal orientation to the fast-ice cover of Hudson Bay and James Bay.

The duration of fast ice over the 19 years ranged from 150–160 days in James Bay to 170–230 days in Hudson Bay (Figure 2c). Figure 2d shows the trend of changing fast-ice duration across the Hudson Bay and James Bay. Areas like Chesterfield Inlet, Rankin Inlet and Arviat were noted to have a negative trend signifying a shortening of the fast-ice period, with the most negative trend observed around Arviat (6 days per decade). Similarly, a positive trend signifying an extended presence of fast ice was observed along eastern Hudson Bay with Belcher Islands encountering a longer fast-ice period of roughly 8 days per decade. The trends of fast-ice duration across Hudson Bay and James Bay (Figure 6) indicate a decreasing period of fast-ice presence in Chesterfield, Rankin Inlet, Arviat, Churchill, Moosonee, while fast ice was staying for a longer duration in Fort Severn, Chisasibi, Inukjuak, Akulivik and Ivujivik.

Figure 6 shows the trend of annual mean air temperature at the 15 study locations along with the trend in fast-ice duration. Cooley et al. (2020) suggested air temperature as the predominant meteorological forcing controlling year-to-year variability of fast-ice duration. However, even though the patterns of air temperature change and ice duration align in Hudson Bay and James Bay, the correlation remains moderate (r² = 0.558, p < 0.05). A gradual shortening of fast-ice persistence is noted in some areas where the air temperature is observed to be increasing. Although some literature has suggested that air temperature plays a pivotal role in controlling ice persistence, in our study the correlation remains weak. Yet, air temperature can safely be assumed to be one of the many physical factors (wind, tides, coastal geomorphology, ocean heat flux, etc.) that influence fast-ice persistence.

The duration of an ice season is an important aspect of ice climatology. Variations in ice duration are indicators of inter-annual changes happening in the local environment and affecting the regional climate, e.g. temperature. For a more detailed understanding of the environmental changes taking place seasonally, fast-ice duration may not provide a direct answer. Specific events associated with changes in fast-ice duration, attributed to early or later fast-ice freeze-up and break-up, can be better indicators of changes happening over seasons, like enhanced warming in the summers or colder winters. In Figure 7a and b, the trend of fast-ice freeze-up and break-up events over the Hudson Bay and James Bay from 2000–2019 is evident. The values of this trend analysis show the rate of change (time of occurrence, in days per year) over 2000–2019. The positive values in the freeze-up trends suggest a delayed onset of fast-ice freezing, while the negative values are indicating an earlier onset of freezing. Similar inference can be made in the case of the break-up trend. Our trend analysis reveals a gradually delayed freeze-up onset in some areas as late as 7.7 days per decade, with a relatively earlier freeze-up in some areas as early as 4.5 days per decade. Analysis of fast-ice freeze-up trends across Chesterfield Inlet, Rankin Inlet, Arviat, Churchill in the western Hudson Bay shows a delayed onset of freezing in these areas, whereas trends across Sanikiluaq (located on the Belcher Islands), Inukjuak, Akulivik, and Ivujivik indicate a relatively earlier freezing event. A similar east–west differentiation was noted in James Bay as well, with fast-ice freeze-up delayed in Attawapiskat (western James Bay) but trending towards earlier in Chisasibi (eastern James Bay). Furthermore, a much delayed fast-ice freeze-up was noted around Moosonee in southern James Bay.

Contrasting trends of fast-ice break-up were observed in Hudson Bay and James Bay. Though fast ice breaks up relatively earlier in the eastern Hudson Bay compared to the northwestern part of the bay, western Hudson Bay shows a negative trend indicating an earlier break-up taking place across the region, whereas the eastern Hudson Bay displays a positive trend indicating towards a delayed break-up event. Historically, similar trends have been observed indicating an earlier break-up of the offshore sea ice (mobile pack ice) in the northwestern Hudson Bay compared to the eastern Hudson Bay (Tivy et al., 2011; Andrew et al., 2018; Bruneau et al., 2021). The strong, persistent westward winds are the primary reason for the mobile ice floes pushed further eastward from the western Hudson Bay region into the south and southeastern Hudson Bay, where the mass accumulation of ice causes sea-ice coverage to persist for a prolonged period of time before melting (Hochheim and Barber, 2014; Landy et al., 2017; Kirillov et al., 2020). The observations in our study support the gradual lengthening of fast-ice presence in the eastern part of the study area and shortening of fast ice in the northwestern and southwestern Hudson Bay divisions, with the larger regional sea ice trends, which further highlights the dominant role of wind direction.

In Figure 8, we highlight two transition period in the annual cycle of air temperature over the 15 study locations: when air temperatures drop below 0°C, between mid-September and early October; and when air temperatures rise above 0°C, between early April and late May (Tables S1–S4). Correlating FTD and day of fast-ice break-up at each of the 15 locations reveals both significant and non-significant relationships (Table 1). Generally, correlations are stronger during fall, indicating a strong interconnection between air temperature and fast-ice formation, whereas correlations are weaker during spring, indicating a greater contribution of various climatological and geophysical forcing to fast-ice break-up. An earlier study has shown an increase in air temperature recorded at weather stations above 60°N coinciding with the decreasing ice cover, with the most substantial warming experienced in the coastal and archipelago areas across the Arctic and sub-Arctic regions (Polyakov et al., 2012). This decreasing ice persistence resulted in the lengthening of the open water conditions. This ice-free condition facilitated oceanic absorbance of more energy from solar radiance; hence surface warming intensified across Arctic marginal seas including Hudson Bay (Gagnon and Gough, 2005; Steele et al., 2008; Steele et al., 2010; Galbraith and Larouche, 2011). A similar association between air temperature and ice persistence was noted in our study. Hence, while warmer air temperatures do delay ice formation, this relationship operates both ways, with delayed ice formation contributing to warmer temperatures by increasing the ocean heat flux to the atmosphere.

As the study area ranges from 51°N latitude (Moosonee) to 64°N latitude (Coral Harbour), climatic differences over this large latitudinal stretch of the Hudson Bay and James Bay combined are among the primary reasons for the observed location-specific variability in mean fast-ice persistence. The Hudson Bay and James Bay experience a wide air temperature regime, from an average air temperature of –22°C during January and February to 10°C in July and August (Figure 8). For the period examined, air temperatures began warming in mid-February and continued rising through August. This warming usually progressed from the south, advancing northwards along the coastline until August. Starting in August, the air temperature began decreasing again following essentially the reverse order to the warming regime (Figure S4). The trend of changing air temperature over the Hudson Bay and James Bay from 2000–2019 showcases a gradual warming pattern observed along the western part of Hudson Bay and southern James Bay, while the northern Hudson Bay experiences a gradual cooling trend (Figure S5).

Although thermodynamic growth is driven principally by air temperature, wind forcing may also play a large role in shaping the fast-ice cover. Larouche and Galbraith (1989) attributed the westward extension of fast ice in eastern Hudson Bay to the northwesterly winds over Hudson Bay. This conclusion was undoubtedly linked to the dynamic thickening of the ice cover in eastern Hudson Bay that Landy et al. (2017) ascribed to eastward transport of pack ice and the associated convergence of it towards eastern Hudson Bay. Past studies in the Canadian Arctic Archipelago have suggested that not only winds but also internal stress within the ice cover and surface currents contribute to the overall growth and stability of fast ice (Agnew et al., 2008). Other forcings such as strong winds and surface currents may play an important role in breaking up the ice cover. Hence, a more extensive examination of wind forcing, ocean temperature, ice drift, tidal currents and large-scale ocean circulation will be essential to theorise the cause of the observed spatial and temporal variations of the fast-ice regime in Hudson Bay and James Bay.

Figure 9 and Table S5 summarize the fast-ice extent for each of the 19 ice seasons around the study area. These observations help to understand the inter-annual variability of the fast-ice extent around Hudson Bay and James Bay, and also highlight the formation of ice bridges connecting the Ottawa Islands (2015 and 2018), Belcher Islands (2007, 2009, 2012, 2013, 2014, 2015, 2017, 2018, 2019), Charlton Islands (2002, 2005, 2007, 2008, 2009, 2011, 2013, 2014, 2015, 2016, 2018, 2019) and Roes Welcome Sound (2007 and 2012) to the mainland. Spatially, the distance of the fast-ice extent varied immensely along the Hudson Bay and James Bay coast. Compared to the northwestern and southwestern Hudson Bay (Zones B and C), eastern Hudson Bay (Zone G) showed a higher variability of the fast-ice extent (Figure 5). As shown in Table 2, fast-ice extent in northwestern Hudson Bay (Zone B) was observed to be around 22 km compared to approximately 10 km in the southwestern Hudson Bay (Zone C), and 14 km in the southern Hudson Bay division (Zone D). This variability indicates a north–south contrast observed particularly in the western Hudson Bay, signifying a greater fast-ice extent with latitude. The exact opposite was observed for the eastern coast of Hudson Bay where the southern part showed a greater fast-ice extent as compared to the northern part. However, this particular observation could be attributed to the formation of ice bridges between Belcher Island and the mainland.

Another interesting observation is the west-to-east gradient of fast-ice extent in the Hudson Bay and James Bay region. Fast ice in the northwestern Hudson Bay division (Zone B) extended to an average distance of 22 km offshore, whereas in the eastern Hudson Bay division (Zone G) the average fast ice extent was 30 km. Similarly, a west-to-east gradient was observed in James Bay with a 15-km average fast-ice extent in western James Bay (Zone E) compared to 26 km in eastern James Bay (Zone F). The western coast of James Bay showcases less spatial variability of the fast-ice extent compared to the eastern coast. This difference can be attributed to the rather straight coastline and a steady offshore gradient of the underlying bathymetry, although the presence of Akimiski Island results in variation to the otherwise uniform ice extent. Fast ice forming around Akimiski Island with an ice bridge connecting the island to the mainland is the only example of an offshore feature that influenced fast-ice cover in the western James Bay (Zone E). These observations closely align with the latitudinal and west-to-east gradients previously discussed by Gagnon and Gough (2006).

Several past studies have discussed the relationship between fast-ice extent and bathymetry (e.g., Zubov, 1945; Jacobs et al., 1975; Barry, 1979; Divine et al., 2004; Mahoney et al., 2007; Jensen et al., 2020; Fraser et al., 2021). Common knowledge exists that the fast-ice edge stabilizes along a specific water depth. However, the isobath where the ice edge stabilizes has also been observed to be highly variable across the Arctic and Antarctic (Mahoney et al., 2014; Fraser et al., 2021). Bathymetry from GEBCO data indicates that the western and southern Hudson Bay has a shallower nearshore zone with a low offshore gradient compared to the eastern Hudson Bay. Similarly, in James Bay the west coast is relatively shallower and has a relatively lower offshore gradient compared to the east coast. Water depths, estimated at the fast-ice edge, corresponding to all 100-km intervals on the coastline, vary greatly. The average depth at the fast-ice edge in the Hudson Bay and James Bay was observed to be 20.5 m (Figure 10), which is similar to observation in the Beaufort Sea Shelf (Mahoney et al., 2014).

To investigate the possible relationship between offshore extent of fast ice and the underlying bathymetry, we correlated the distance of the fast-ice edge from the coastline with the distance of the 20-m isobaths from the coastline at each of the 100-km intervals (Figure 11). The generation of points along the coastline to measure fast-ice extent and distance of the 20-m isobaths from the coastline resulted in a different number of transects per zone. The impact of this unevenness on the overall analysis lies in the number of data points that feed into the various statistical analyses.

The relationship observed through this analysis is complex, with an overall correlation of r = 0.306 (p < 0.05) (Table 2). Examining this relationship for the different zones of Hudson Bay reveals a strong correlation in western James Bay (Zone E; r = 0.89) and southwestern Hudson Bay (Zone C; r = 0.55), which are both areas with a fairly consistent fast-ice edge. In other zones the correlation between fast-ice extent and water depth was weak. Thus, this analysis shows that water depth alone does not control the physical stability of the ice extension. Complex geometry of the coastline and the variable underwater topography are playing major roles in stabilizing the fast-ice edge as well. Similar observations were made by Porter-Smith et al. (2021) who indicated that complexity of the coast can be an important factor responsible for the spatial variability pattern and persistence of fast ice.

The presence of offshore islands, lagoons and barrier islands in certain areas is associated with a typically stable and far extended fast-ice cover and the frequent formation of ice bridges. With ice bridges predominantly forming around Ottawa Islands, Belcher Islands, Charlton Islands and Roes Welcome Sound, the fast-ice edge at these locations was extended much further than in the rest of the bay (Figure 9). The presence of islands provides enhanced stability for the floating fast-ice extensions, further facilitating the formation of a continuous fast-ice cover from the islands to mainland. In such locations the offshore extent of fast ice depends more on the proximity of the offshore features than on the underlying bathymetry. Fast ice forming in these regions is protected and well anchored to its surrounding coastline or submerged features hence allowing the fast-ice edge to extend further offshore. Nearshore regions of the Coats Island, Mansel Island, and parts of the southern Hudson Bay coastal region are dominated by underwater ridges, which serve as stabilizing platforms for the fast ice, thereby facilitating a further seaward extension.

We find that the entire Hudson Bay and James Bay coastline can be divided broadly into three categories, Types I to III (Figure 12). Type I represents a relatively straight and flat coastline with a steady offshore gradient, which is best exemplified by southern Hudson Bay and western James Bay. Type II represents a more complex coast, best exemplified by eastern James Bay and a majority of eastern Hudson Bay with the presence of numerous small and large islands and rocky outcrops close to the mainland. The third category, Type III, represents coasts with an unsteady offshore gradient due to the presence of submerged mounds and ridges. Type III is best exemplified by the southwestern coasts of Coats and Mansel Islands and parts of southern Hudson Bay (Cape Tatnum), as well as eastern Hudson Bay (south of Inukjuak).

A general representation of the mean fast-ice edge occurring over the 15 locations across the Hudson Bay and James Bay selected for this study (Figure S6) highlights the water depth to which they are limited. Previously (Figures 10 and 12), we established that the limit of the fast-ice edge and its stability is majorly controlled by the type of coast and coastal topography. As other physical forcings like wind and tidal currents destabilize the fast-ice extensions, complex coastal features shelter the ice against the action of these forcings, as observed in Coral Harbour, Chisasibi and Moosonee. Raised underwater features and offshore features provide substratum for the ice, stabilizing it in place and allowing it to extend further offshore, as seen in Chesterfield Inlet, Rankin Inlet, Attawapiskat, Moosonee, Chisasibi, Sanikiluaq, Ivujivik and Akulivik.

In this study we have provided a comprehensive analysis of fast-ice regimes and their variability along approximately 10,000 km of coastline in Hudson Bay and James Bay. Our results show that between 2000 and 2019 trends towards delayed freeze-up and earlier break-up led to an overall shorter fast-ice season in western Hudson Bay. Conversely, the duration of fast ice in eastern Hudson Bay increased during this time. A similar pattern was observed in James Bay with fast-ice duration decreasing along the western and southern coast while fast-ice duration increased along the east coast. Trends towards later freeze-up generally correlated with increasing air temperatures, while trends towards earlier break-up did not correlate with air temperatures, highlighting the influence of dynamics (winds and currents) on fast-ice break-up.

The mean fast-ice extent also displayed regional variability, as the fast-ice edge was observed to extend further offshore in eastern Hudson Bay and James Bay compared to western Hudson Bay and James Bay. Further analysis also revealed a decreasing trend in fast-ice extent in southwestern Hudson Bay and western James Bay regions. Even though several climatological and geophysical factors influence the variability in fast-ice extent, in this study we found that variations in fast-ice duration and extent corresponded with changing air temperatures and variability in coastal topography. The fast-ice edge is limited in coastal zones with straight coastlines and a steady offshore slope, whereas the presence of offshore features and variable underwater topography supports the stabilization of fast ice and leads to greater extent offshore. Not only do the complex coasts affect fast-ice growth but these features also protect the ice against dynamic forces like wind, waves and currents, which can prevent fast ice from stabilizing and promote break-up. The variation in coastal topography and its implications to fast-ice growth and stability can be one of the key reasons behind the notable east–west contrast of fast-ice extent and break-up events in the Hudson Bay and James Bay system. Areas exhibiting a simple coastline orientation and steady bathymetric gradient like southwestern Hudson Bay and western James Bay showed strong correlation (0.55 and 0.89, respectively) between fast-ice extent and water depth, whereas correlations remained weak in areas with a complex coastline orientation and an inconsistent bathymetric gradient. Considering the role that coastal orientation can play in affecting fast-ice stability, we conclude that it has considerable influence on fast-ice persistence as well. Areas with complex coastlines and presence of offshore features like the eastern shore of Hudson Bay and James Bay exhibited an enhanced fast-ice persistence compared to other locations like southern and western Hudson Bay and James Bay where the coast is rather open and planer. This enhanced persistence can be attributed to the sheltering effect that the coastal features provide to the fast-ice cover, protecting fast ice against dynamic forces like wind and waves and allowing a much more stable fast-ice cover of longer extension. This east–west gradient is strengthened further by the ice drift patterns observed in the bay (Kirillov et al, 2020), where ice drifts away from the west and clogs in the east resulting in increased ice thickness and longer ice persistence.

Through this study we have observed that this variability in fast-ice regime across the study area can be attributed to different morphological and climatic regimes. Several communities reside on the coastal margins of Hudson Bay and James Bay. Changes in the fast-ice cover and duration impact the livelihood of the coastal communities as well as marine biota thriving on or close to the coastal margins. People residing in these communities depend on the fast-ice cover for transport and hunting. The current scenario of a changing fast-ice regime is forcing people to adapt to an increasingly unsafe and unpredictable ice cover situation (Ford et al., 2008; Druckenmiller et al., 2010). Better observations and predictions of fast-ice cover will help the communities to be well equipped for vulnerabilities on ice and adapting to the changing ice regime (Furgal and Seguin, 2006). As fast-ice regimes depend highly on an array of environmental forcings like surface temperature, wind, surface water currents and ice drift, studying the variations of these forcings is important to create a concise understanding of the fast-ice regime leading to a better approach for predictive and monitoring studies. Regional factors like coastline orientation and bathymetry can attenuate the impact of climatological forcings, as seen in the east–west contrast of Hudson Bay and James Bay. Fast-ice variability in areas with an open and planer coastline and steady bathymetric slope like southern Hudson Bay and western James Bay are more impacted climatically than the rest of the bay; hence these areas provide better prospects to be considered for climate change indicator-based studies.

The sources of the primary datasets used in the study are as follows:

• Weekly ice charts from The Canadian Ice Service. (https://iceweb1.cis.ec.gc.ca/Archive/page1.xhtml?)

• Air temperature reanalysis products from MERRA2. (https://disc.gsfc.nasa.gov/datasets/M2TMNXSLV_5.12.4/summary?keywords=MERRA_)

• Bathymetry products from GEBCO Global Bathymetry. (https://neo.sci.gsfc.nasa.gov/view.php?datasetId=GEBCO_BATHY)

Other datasets include satellite images from Landsat series (https://earthexplorer.usgs.gov/) and daily observations from the online viewing platform NASA Worldview (https://worldview.earthdata.nasa.gov/).

The supplemental files for this article can be found as follows:

Figures S1–S6. Tables S1–S5. Docx

The authors thank all collaborators of the BaySys program and Manitoba Hydro. Thank you Kevin Sydor and Karen Wong from Manitoba Hydro for providing comments on the manuscript. This work is a contribution to the Arctic Science Partnership (ASP, asp-net.org) and ArcticNet. Thank you R. Galley for providing much needed support while developing this manuscript.

This work is a contribution to the Natural Sciences and Engineering Council of Canada (NSERC) Collaborative Research and Development project: BaySys (CRDPJ 470028-14) led by D. Barber (Academic PI) and K. Sydor (Industry PI). Individual support from NSERC has been provided to D. Barber, D. Babb, A. Mukhopadhyay and J. Ehn. Additional support was provided to D Babb by the Canadian Meteorological and Oceanographic Society.

The authors have no competing interests to declare.

• Contributed to original data acquisition: KG, AM.

• Contributed to analysis and interpretation of the data: All authors.

• Drafted the article: KG, JE, DBabb.

• Revised the article: All authors.

• Approved the submitted version for publication: All authors.

How to cite this article: Gupta, K, Mukhopadhyay, A, Babb, DG, Barber, DG, Ehn, JK. 2022. Landfast sea ice in Hudson Bay and James Bay: Annual cycle, variability and trends, 2000–2019. Elementa: Science of the Anthropocene 10(1). DOI: https://doi.org/10.1525/elementa.2021.00073

Domain Editor-in-Chief: Jody W. Deming, University of Washington, Seattle, WA, USA

Associate Editor: Stephen F. Ackley, Department of Geological Sciences, University of Texas at San Antonio, TX, USA

Knowledge Domain: Ocean Science

Part of an Elementa Special Feature: The Hudson Bay System Study (BaySys)