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Synthesis Report on Landscape Response to Climate Change:

Draft Version 1

Prepared by Andrew Davidson

March 2005

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Contents

CONTENTS..........................................................................................................................................2

1 INTRODUCTION........................................................................................................................5

2 LANDSCAPE RESPONSES TO CLIMATE CHANGE .........................................................6

2.1 PERMAFROST .........................................................................................................................6 2.1.1 What is permafrost? ...........................................................................................................6 2.1.2 Factors influencing the occurrence of permafrost.............................................................7 2.1.3 Geographic distribution of permafrost ..............................................................................7 2.1.4 Permafrost and Climate Change........................................................................................9

2.1.4.1 Permafrost response to past climates .........................................................................9 2.1.4.2 Permafrost response to present climate....................................................................10

2.1.5 Predicting permafrost response to future climates ..........................................................12 2.1.5.1 The importance of predicting permafrost responses to future climates ...................12 2.1.5.2 A Geographic Information System (GIS) approach.................................................13 2.1.5.3 A process-based modeling approach........................................................................16 2.1.5.4 In situ monitoring of permafrost regions .................................................................19

2.1.6 Consequences of permafrost responses to a changing climate........................................20 2.1.6.1 The physical response of permafrost terrain ............................................................20 2.1.6.2 Responses of groundwater, river and lake systems..................................................21 2.1.6.3 Potential challenges to northern development .........................................................22 2.1.6.4 Case studies from Canada’s permafrost zone ..........................................................25 2.1.6.5 Potential feedbacks to the global carbon cycle ........................................................28

2.1.7 References ........................................................................................................................29

2.2 GLACIERS.............................................................................................................................36 2.2.1 What are Glaciers? ..........................................................................................................36 2.2.2 Factors influencing the distribution of glaciers...............................................................37 2.2.3 Geographic distribution of glaciers .................................................................................38 2.2.4 Glaciers and climate change............................................................................................39

2.2.4.1 Short-, medium- and long-term glacier responses to climate change .....................39 2.2.4.2 Glacier responses to past and present climates ........................................................40

2.2.5 Predicting glacier response to future climates ................................................................42 2.2.5.1 The importance of predicting glacier response to future climates ...........................42 2.2.5.2 In situ monitoring of glacier dynamics ....................................................................43 2.2.5.3 Remote sensing of glacier dynamics........................................................................44 2.2.5.4 A modeling approach ...............................................................................................56

2.2.6 Consequences of glacier response to a changing climate ...............................................58 2.2.6.1 Potential impacts on Canada’s freshwater resources ...............................................58 2.2.6.2 Potential impacts on surface-atmosphere energy exchange.....................................59

2.2.7 References ........................................................................................................................60

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2.3 NON-PERMANENT (SEASONAL) ICE AND SNOW COVER.....................................................65 2.3.1 Seasonal snow and ice cover............................................................................................65

2.3.1.1 Land surface ice and snow .......................................................................................65 2.3.1.2 Lake and river ice.....................................................................................................66 2.3.1.3 Sea ice ......................................................................................................................66

2.3.2 Snow and ice cover responses to past and present climates............................................67 2.3.2.1 Land surface ice and snow .......................................................................................67 2.3.2.2 Lake and river ice.....................................................................................................68 2.3.2.3 Sea ice ......................................................................................................................69

2.3.3 Predicting snow and ice response to future climates.......................................................71 2.3.3.1 The importance of predicting snow and ice response to future climates .................71 2.3.3.2 Remote sensing of ice and snow dynamics..............................................................71 2.3.3.3 A modeling approach ...............................................................................................76

2.3.4 Implications of changes in snow and ice cover................................................................79 2.3.4.1 Potential impacts of changes in land, lake and river snow and ice cover extent on Canada’s freshwater resources.................................................................................................79 2.3.4.2 Potential impacts of changes in land, lake and river snow and ice cover extent on surface-atmosphere energy exchange ......................................................................................79 2.3.4.3 Potential impacts of changes in sea ice ....................................................................80

2.3.5 References ........................................................................................................................80

2.4 WATER-DOMINATED LANDSCAPES .....................................................................................86 2.4.1 Canada’s water-dominated landscapes ...........................................................................86

2.4.1.1 Canada’s water resources.........................................................................................86 2.4.1.2 Groundwater.............................................................................................................86 2.4.1.3 Lakes and Reservoirs ...............................................................................................87 2.4.1.4 Rivers .......................................................................................................................88 2.4.1.5 Wetlands...................................................................................................................89

2.4.2 Potential impacts of climate change on the hydrological cycle.......................................90 2.4.2.1 The hydrological cycle and its components .............................................................90 2.4.2.2 Precipitation .............................................................................................................91 2.4.2.3 Evaporation / Evapotranspiration.............................................................................92 2.4.2.4 Groundwater.............................................................................................................93 2.4.2.5 Lakes and reservoirs.................................................................................................94 2.4.2.6 Streams and runoff ...................................................................................................95 2.4.2.7 Wetlands...................................................................................................................97

2.4.3 Predicting freshwater responses to changing climate: Regional perspectives................98 2.4.3.1 A National Perspective.............................................................................................98 2.4.3.2 A Regional Perspective - Atlantic region.................................................................98 2.4.3.3 A Regional Perspective - Quebec ..........................................................................102 2.4.3.4 A Regional Perspective - Ontario ..........................................................................104 2.4.3.5 A Regional Perspective - Prairie Provinces ...........................................................108 2.4.3.6 A Regional Perspective - The Arctic and the North ..............................................112 2.4.3.7 A Regional Perspective - British Columbia ...........................................................117

2.4.4 References ......................................................................................................................119

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2.5 COASTAL ZONES................................................................................................................132 2.5.1 The importance of Coastal Zone Ecosystems.................................................................132 2.5.2 Predicted changes in Canada’s coastlines under a changing climate ..........................132 2.5.3 Coastal Zone Impact Issues in Canada: A Regional Perspective..................................134

2.5.3.1 General Impacts .....................................................................................................134 2.5.3.2 Impacts on Arctic Coastlines .................................................................................134 2.5.3.3 Impacts on AtlanticCoastlines................................................................................137 2.5.3.4 Impacts on PacificCoastlines .................................................................................139

2.5.4 C-CIARN and the Identification of Canada’s Vulnerability to Water Level Change ....140 2.5.4.1 Water Levels ..........................................................................................................141 2.5.4.2 Mapping and Surveying .........................................................................................142 2.5.4.3 Vulnerability and Risk Assessment Mapping ........................................................143 2.5.4.4 Adaptation Options and Decision-Making.............................................................144 2.5.4.5 Education and Communication ..............................................................................145

2.5.5 References ......................................................................................................................146

2.6 GRASSLANDS......................................................................................................................147 2.6.1 The Importance of Grassland Ecosystems .....................................................................147 2.6.2 The Grasslands of North America..................................................................................147

2.6.2.1 The North American Grassland Biome..................................................................147 2.6.2.2 The Canadian Northern Mixed Grass Prairie.........................................................149

2.6.3 Predicting grassland response to future climates..........................................................149 2.6.3.1 Grassland responses to future climates ..................................................................149 2.6.3.2 In situ monitoring of grassland dynamics ..............................................................152 2.6.3.3 Remote sensing of grassland dynamics..................................................................157

2.6.4 Implications of climate changes on Canada’s grasslands.............................................166 2.6.4.1 Potential impacts on agriculture.............................................................................166 2.6.4.2 Potential feedbacks to the global carbon cycle ......................................................167

2.6.5 References ......................................................................................................................169

3 CONCLUSIONS ......................................................................................................................175

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1 Introduction

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2 Landscape Responses to Climate Change

2.1 Permafrost

2.1.1 What is permafrost?

Permafrost is defined as rock, sediment or any other earth material that remains at or below

0°C for at least two consecutive years (International Permafrost Association, 1998). Permafrost is

defined solely by temperature, not by the physical state of its soil moisture content (i.e. liquid water

vs. ice). In some materials, appreciable amounts of pore water can remain unfrozen at temperatures

several degrees below 0°C [Williams and Smith, 1989b]. This situation may occur where dissolved

salts or pressure effects depress the freezing point of water [Smith and Burgess, 2004]. As a result,

permafrost can contain significant amounts of ice, or practically no ice at all.

The uppermost layer of ground in a permafrost area

is called the active layer, and it is bounded at its base by the

permafrost table [Figure 1]. The active layer is only

seasonally frozen, and is the zone where maximum annual

temperatures exceed 0°C [Burgess and Smith, 2000]. The

layer immediately beneath the active layer is permafrost.

Permafrost develops where the depth of freezing in winter

exceeds the depth of thawing in summer. This layer thus

remains frozen through the summer months. Together, the

active layer and permafrost are the primary subsurface

components of the Arctic land-atmosphere system.

Almost all of the soil moisture in permafrost occurs in the form of ground ice. Ground ice is

usually concentrated in the upper-most layers of permafrost, and can occur in a variety of forms. It

can occur as structure-forming ice, which bonds the enclosing sediment, or as large bodies of almost

pure (massive) ice. The structure-forming ice comprises segregated ice, intrusive ice, reticulate vein

ice, ice crystals, and icy coatings on soil particles. The large bodies of more or less pure ice occur as

pingo cores, massive icy beds, and ice wedges [Smith and Burgess, 2004]. As a result, the spatial

distribution of ground ice is highly variable, and can range from almost 100% by volume at locations

Figure 1: Permafrost profile.

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where massive ice is present to virtually zero in dry permafrost [Briggs et al., 1993]. Ground ice is

one of the most important attributes of the terrain in permafrost regions. Its presence influences

topography, geomorphic processes, vegetation, and the response of landscape to environmental

changes.

2.1.2 Factors influencing the occurrence of permafrost

The global pattern of permafrost distribution is mainly determined by climate [Burgess and

Smith, 2000]. Regions underlain by permafrost generally share similar climatic characteristics: they

(a) have daily temperatures that are below 0°C for at least nine months of the year, and below -10°C

for at least six months of the year; (b) have summer temperatures that rarely exceed 20ºC; and (c)

experience low amounts of precipitation (less than 100mm in winter and 300mm in summer). The

large anticyclonic continental polar air masses are largely responsible for producing these conditions

of intense cold and aridity [Summerfield, 1991].

However, the thickness, temperature and stability of permafrost at any particular location are

also dependent on site-specific environmental factors that influence the amount of energy available

for heating the ground. Temperature in the ground is directly related to the balance between

incoming heat (mainly from incoming solar radiation, but also from geothermal energy arriving from

the earth’s interior) and the amount of heat re-radiated or reflected from the ground surface and

consumed by evapotranspiration [Burgess and Smith, 2000]. Factors such as slope, altitude, aspect,

vegetation type and density, snow cover, surficial materials, the presence or absence of organic

materials, soil moisture content and drainage can all influence the surface energy balance [Smith and

Burgess, 2004]. Many of these factors can change dramatically over time, thereby leading to the

degradation or aggradation of a permafrost layer [Summerfield, 1991]. Geothermal heat, in

conjunction with the thermal properties of subsurface materials, determines the maximum depth to

which permafrost can reach [Burgess and Smith, 2000].

2.1.3 Geographic distribution of permafrost

Permafrost currently underlies approximately 22.79 million km2 (about 24%) of the exposed

land surface of the northern hemisphere [Zhang et al., 1999]. Extensive permafrost is found in

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regions of Canada, Alaska, Russia, China, Greenland and Scandinavia. Only in the very coldest of

these regions is permafrost cover continuous (>90% permafrost cover). South of this zone lies a

broad zone of discontinuous permafrost, where areas underlain by permafrost coexist with areas of

unfrozen ground. Here, the proportion of frozen ground generally decreases as one moves

progressively southwards through the widespread (50-90% cover), sporadic (10-50% cover) and

localized (<10% cover) permafrost sub-zones [Smith and Burgess, 2004]. The widespread

distribution of permafrost makes it a significant component of the cryosphere [Williams and Smith,

1989a].

Permafrost covers approximately 42% of the Canadian landmass [Kettles et al., 1997]. It can

range from thin layers that have remained frozen from one winter to the next, to frozen ground

hundreds of metres in thickness, and thousands of years in age [Trenhaile, 1990]. Three broad

permafrost regions can be identified within the Canadian landmass [Figure 2]: (1) the continuous

permafrost zone extends from the Arctic Islands

as far south as southern Hudson Bay. Here,

unfrozen ground only occurs beneath large

bodies of water and in small areas of newly

deposited sediments [Smith et al., 2001].

Permafrost thicknesses in this zone vary from

more than 500 m in the north to 100m at its

southern limits [Briggs et al., 1993]; (2) the

discontinuous permafrost zone lies to the south

of the continuous permafrost zone. Here,

permafrost decreases southward in area and

thickness, until it is present only sporadically,

usually only in patches of terrain that display

locally favourable conditions for permafrost

formation (e.g. areas of elevated organic terrain) [Kettles et al., 1997; Smith et al., 2001]. At these

sites, permafrost is often only a few metres thick [Briggs et al., 1993]; and (3) alpine permafrost

occurs where conditions favouring the existence of permafrost also prevail at high altitudes. The

zone of discontinuous alpine permafrost typically occupies an elevation range of about 1500m within

Figure 2: The distribution of permafrost zonesin North America.

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the Rocky Mountains. At higher elevations, alpine permafrost tends to be continuous [Briggs et al.,

1993].

2.1.4 Permafrost and Climate Change

2.1.4.1 Permafrost response to past climates

Permafrost is a thermal condition, and thus changes in its historical distribution are strongly

linked to variations in climate. Climate, however, is not constant and has undergone significant

changes detectable at time scales ranging from decades to millennia. Under cooling conditions,

permafrost generally increases in its areal extent and thickness. In comparison, warming conditions

generally lead to an increase in active layer thickness, permafrost thinning, and in some cases, the

complete disappearance of permafrost [Smith and Burgess, 2004]. The surface buffer layer and the

natural damping effect of subsurface materials cause a lag in the permafrost response to changes in

surface conditions or climate [Smith et al., 2001].

During the glacial maxima of the Quaternary period (X BP), large areas of the continental

shelf were above sea level [Blasco et al., 1990]. These regions were thus exposed to air temperatures

that were as much as 18°C lower than those currently occurring in the seabed [Allen et al., 1988].

These conditions allowed the formation of permafrost up to 700m thick. These regions were then

covered by the Arctic Ocean during warmer interglacial periods. Because seabottom temperatures

range between –2 and 0°C, the thermal regime of the subsea permafrost is in disequilibrium with the

present marine environment [Taylor, 1991]. As a result, these sediments are warming gradually, and

the permafrost is slowly degrading.

A general warm period followed the disappearance of glacial ice. Temperatures peaked

during the Holocene between 9000BP and 6000BP. Zoltai [1995] presented evidence derived from

macrofossil analysis and the radiocarbon dating of peat cores to suggest that mean annual

temperatures were about 5ºC warmer than present and the southern limit of permafrost was 300-

500km northward of its current position. Much of the present discontinuous permafrost zone (see

next section) was likely free of permafrost during this time. Where permafrost did exist during this

time, active-layer thickness was probably greater than at present [see Burn et al., 1986].

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Permafrost became more extensive during cooler conditions following the mid-Holocene

warm period. Zoltai [1993] presented evidence to suggest that permafrost was established in

northwestern Alberta 3700 years ago and that the climate at this time probably resembled the present

climate regime. In the Mackenzie Delta region, permafrost aggradation and pingo development

occurred in response to cooling that began about 5000 BP [Vardy et al., 1998].

During the Little Ice Age (1550 – 1850 AD), temperatures were about 1°C cooler than

present and permafrost occurred farther south than it does today [Vitt et al., 1994]. These conditions

likely were responsible for forming the frozen peatlands that occur today at the southern margin of

the discontinuous permafrost zone. Much of the permafrost formed at this time has generally

degraded in response to warming, but has been preserved in some areas to due the insulating

properties of thick peat cover [Halsey et al., 1995].

There is also evidence for climate-changes to permafrost since the Little Ice Age. Analysis of

borehole temperatures in Alaska by Lachenbruch and Marshall [1986] indicate a general warming

trend during the last century. Furthermore, an observed warming of 1°C in the western arctic has

caused the eradication of thin permafrost and an apparent northward displacement in the southern

boundary of the discontinuous permafrost zone [Kwong and Gan, 1994], and an increase in

permafrost temperatures in Yukon Territory and western Northwest Territories [Halsey et al., 1995].

More recent observations have shown permafrost degradation in Manitoba and Quebec in the

southern margin of the permafrost region, especially where there is no peat layer [French and

Egerov, 1998; Laberge and Payette, 1995]. In the eastern arctic, however, recent cooling and

aggradation of permafrost has occurred. Records for northern Quebec show a decrease in air

temperature between 1947 and 1992 ranging from 0.02 - 0.03ºC per year. An analysis of ground

temperatures in the upper 20m for the period 1988-1993 indicates that permafrost also cooled over

this time period [Allard et al., 1995].

2.1.4.2 Permafrost response to present climate

Permafrost conditions are also dynamic under the current climate. Changes in the permafrost

environment can be attributed to both natural and anthropogenic causes. The natural factors affecting

permafrost include the inter-annual variability in rainfall, snowfall, or the occurrence of drought and

fires. Soil temperature and active-layer thickness may temporarily increase following a fire [Liang et

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al., 1991]. The increase in temperature is not due to the heat from the fire, but rather to the loss of the

shading effect of trees, the removal of the insulating organic layer, and a decrease in the reflectivity

of the surface, which results in a greater absorption of radiation by the ground surface [Johnson and

Viereck, 1983]. Within a few years, active-layer thickness may return to pre-burn levels if the fire

was not severe and vegetation regeneration occurs over a short period. Where vegetation

regeneration is slow, long-term permafrost degradation may continue [Burn, 1998]. Extreme climatic

events, for example higher than normal air temperatures associated with El Niño events, may also

affect the permafrost environment and can lead to increases in active layer thickness, thaw

settlement, and slope instability. The anthropogenic factors affecting permafrost primarily relate to

human-induced changes in the global climate. Since the onset of the industrial revolution,

atmospheric concentrations of CO2 have increased from 280 to 379 ppm. This change in atmospheric

composition has been accompanied by a steady increase in atmospheric temperature [IPCC1996].

This temperature increase is thought to be responsible for the previously described changes in

permafrost observed over the last century.

Further evidence for current changes in permafrost comes from the series of Canadian

Thermal Permafrost Monitoring Sites implemented by the Geological Survey of Canada (GSC) and

other government departments. Tarnocai et al., [2004], Romanovsky et al., [2002] and Smith et al.,

[2005] have used observations from the Mackenzie valley, Alert, Baker Lake and Iqaluit to

investigate the relationship between changes in permafrost temperature and changes in climate from

the 1980s to present. The results of these studies generally show that (a) the observed warming of

Canadian permafrost is consistent with regional changes in air temperature since the 1970s; (b) this

observed warming is spatially highly variable, and is highly dependence on local surface conditions

that influence the response of permafrost to changes in air temperature; (c) while permafrost

warming has occurred in the western Canadian arctic since the mid-to-late 1980s, the greatest

warming (0.3 – 0.6°C per decade) has occurred in the central and northern Mackenzie valley; and (d)

the warming of permafrost in the high and eastern Canadian arctic appears to have occurred later

than the western arctic, with the greatest warming occurring since the mid-1990s. Detailed

summaries of permafrost monitoring activity in Canada are provided by Burgess et al., [2001] and

the Canadian Permafrost Network Website (www.canpfnetwork.com}.

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2.1.5 Predicting permafrost response to future climates

2.1.5.1 The importance of predicting permafrost responses to future climates

It is predicted that a doubling of atmospheric CO2 over pre-industrial levels will occur

between 2050 and 2100 AD. This change in atmospheric composition is expected to bring about an

average global increase in air temperature ranging from 1 – 3.5ºC, a warming that is not expected to

be geographically uniform. It is expected that the greatest increases in air temperature will occur at

higher latitudes [Flato et al., 2000; Maxwell, 1997]. Indeed, the coupled general circulation model of

the Canadian Centre for Climate Modelling and Analysis (CCCma) of Environment Canada predicts

an increase in mean annual air temperature ranging from 2 - 6°C over the Canadian permafrost

region under a CO2-doubling scenario [Smith and Burgess, 1999; Smith et al., 2005]. If a similar

increase in mean ground surface temperature as that predicted for air temperature is assumed, as

much as 58% of Canada’s permafrost zone could experience rises in ground temperatures to above

0°C [Smith and Burgess, 1999]. In the continuous permafrost zone, such warming will likely lead to

lead to severe permafrost degradation. In the discontinuous permafrost zone, where ground

temperatures are within 1-2°C of the melting point of ice, such warming will lead to a thinning – and

ultimately the disappearance – of permafrost from the region [Smith et al., 2005]. Indeed, the 1995

report of the Intergovernmental Panel on Climate Change (IPCC) predicted the disappearance of

most of the ice-rich permafrost in the present discontinuous zone over the next century [IPCC1996].

Where ground ice contents are high, permafrost degradation will have associated physical impacts.

Of greatest concern are soils with the potential for instability upon thaw (thaw settlement, creep or

slope failure). Such instabilities may have implications for the landscape, ecosystems, and

infrastructure [Romanovsky et al., 2002]. Clearly, the assessment and prediction of the impact of

climate change on permafrost is necessary in order to determine whether adaptation measures will be

required. However, it is not possible to predict the response of permafrost to climate change by

simply applying the projected warming trend in the atmosphere to the ground [Zhang et al., in

review]. This is because changes in vegetation, snow cover, soil moisture, and other climatic

variables (e.g. precipitation, solar radiation and humidity) can all influence water and energy fluxes

on the ground surface and in the soil, and therefore, modulate the relationship between air and

ground temperature. Thus, predicting how permafrost will respond to such changes is a non-trivial

task.

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2.1.5.2 A Geographic Information System (GIS) approach

Another approach to predicting how permafrost will respond to future climates is to identify

and map areas of the Canadian permafrost region that are most sensitive to climate warming. Smith

and Burgess [1998; 1999; 2004] used a GIS approach to map the thermal and physical sensitivities of

Canada’s permafrost regions to climate warming. To achieve this, they combined raster maps

corresponding to the dominant factors influencing the thermal and physical responses of permafrost

to climate change in a GIS. (Here, the thermal response of permafrost to climate warming

corresponds to the relative rate and magnitude of ground temperature change, while the physical

response of permafrost to climate warming corresponds to the relative magnitude and impact of

permafrost thaw). The factors included in the analysis included snow cover, vegetation cover, soil

organic matter content, mineral soil component (for the thermal response of permafrost to climate

change) and surficial geology and peatland distribution (for the physical response of permafrost to

climate change). These factors are important because they act as thermal buffers between the

atmosphere and the ground; that is, their presence/absence/type determines how rapidly permafrost

will respond to a warming climate [Smith, 1988; Smith and Burgess, 1998]. For example, locations

containing vegetation, snow or organic material are more able to buffer the effects of a warming

climate compared to regions where vegetation cover, snow cover and organic material are absent

because the link between ground temperature and climate is more direct in the latter case [Smith and

Burgess, 1998].

The GIS approach is as follows. All input factor maps were provided at a spatial resolution of

10km, and covered the entire Canadian landmass. For each factor (i.e. raster map), pixels were

assigned a high ranking if they corresponded to a condition that would bring about the greatest

response to climate warming. All factors (i.e. all maps) were then combined using an added factor

analysis, where the ranks for each factor were added on a per-pixel basis. Two separate output maps

were produced using this method. These maps represented the per-pixel summation of the factors

influencing the thermal and physical responses to climate change. In these maps, the highest pixel

values corresponded to permafrost regions that were predicted to be most vulnerable to climate

change. The input data and statistical methods used in this approach are described in detail by Smith

and Burgess [2004].

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Figures 3 and 4 show the spatial

variation in the thermal and physical

responses to climate warming within the

Canadian permafrost zone, respectively.

Figure 3 shows that >90% of the present

permafrost region would have a medium to

high thermal response to increases in air

temperature [Smith and Burgess, 2004].

Most of this area is located in the northern

portion of the permafrost region where there

is a minimal buffer layer [Smith and

Burgess, 1998] and surficial material

consists of coarser grained sediments or

bedrock [Smith and Burgess, 1999; Smith

and Burgess, 2004]. The thermal response to

climate warming is generally low to medium

in the southern portion of the permafrost

region where there is a substantial buffer

layer overlying organic or fine-grained

sediments [Smith and Burgess, 1999; Smith

and Burgess, 2004]. The magnitude and rate of response to climate warming will be lower in these

latter regions [Smith and Burgess, 1998]. However, it is important to note that most of the area that

would exhibit a high thermal response to warming is underlain by permafrost at temperatures colder

than -5°C, and therefore, has a lower potential for permafrost thaw [Smith and Burgess, 2004]. The

area with the greatest potential for permafrost thaw is thus >50% of the current permafrost zone

where ground temperatures > -2°C [Smith and Burgess, 2004]. Figure 4 shows that the physical

response to warming will be high in about 12% of the present permafrost region (excluding areas of

massive ice). Most of this is located in the southern region of the permafrost zone where ground

temperatures are warmer than -2°C and the potential of permafrost thaw is high. Throughout the rest

of the permafrost region where ground temperatures are lower, the impact of permafrost thaw is

generally low to medium. However, there are extensive areas where massive ice is present, and the

Figures 3 and 4: Physical and thermal responses of permafrost to climate warming

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impact of permafrost thaw in these areas would be more severe than indicated [Smith and Burgess,

1998].

However, the sensitivity of permafrost to warming consists of both the thermal and physical

responses to warming. Smith and Burgess [1999; 2004] combined these responses to create single

map that portrayed the overall sensitivity of permafrost to warming [Figure 5]. When creating this

map, more weight was given to the physical response of permafrost in the development of the

sensitivity index because the consequences of permafrost thaw are considered to be more important

than the thermal response [Smith and Burgess, 1999]. The sensitivity to climate warming can be low

in areas that show a high thermal response to warming, but also show a minimal physical response

due to the low ice content of the underlying materials. Areas classified as having a low thermal

response, but a high physical response to climate warming are considered to be more vulnerable to

climate warming.

Figure 5 highlights the following important trends [after Smith and Burgess, 2004]: (1)

Approximately 50% of the area within the zone containing warm permafrost is classified as having a

moderate to high sensitivity to warming. A significant portion of this area, however, is within the

sporadic and localized permafrost zones where permafrost may underlie less < 50% of the landscape,

often being limited to areas of organic terrain. Areas covered by organic cryosols are classified as

moderately sensitive to climate warming because although they are thaw-sensitive, their thermal

response to warming is generally low; (2) Much of the permafrost in the southern portion of the

discontinuous zone is not in equilibrium

with the present climate, but has been

preserved due to the insulating properties of

the peat. This permafrost should start to

degrade when mean annual air temperature

rises to above –3.5 °C [Halsey et al., 1995];

(3) Permafrost is considered to be

moderately to highly sensitive to warming

throughout the Mackenzie valley. This is

consistent with geothermal models that

suggest that a complete degradation of Figure 5: Overall sensitivity of permafrost to climate warming

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permafrost may occur in this region within the next 50 years in response to climate warming [e.g.

Burgess et al., 2004]; (4) Throughout the northern prairies and the Hudson Bay lowlands, the

sensitivity of permafrost to warming is considered to be moderate to high. This is mainly due to the

high ice-content soils associated with organic soils in this region; (5) Although permafrost is warm

and thermal response is high in the Cordilleran region, the overall sensitivity of permafrost to

warming is generally low. This is because the surficial material overlying bedrock is generally thin

and thaw stable in this region; (6) The sensitivity of permafrost to warming in Quebec and Labrador

is moderate to low except in a few patches of organic terrain or where moderate-to-high ice content

soils (silt, clay or till) are present; (7) While the potential for permafrost thaw is low where

permafrost is colder and thicker, such as in the Arctic Islands and the Canadian Shield area of

western Nunavut, progressive increases in active-layer thickness are expected where the thermal

response is considered to be high. The physical response is classified as moderate in a significant

portion of the continuous permafrost zone due to the lower structural ice content of surficial

materials, but there are extensive regions where massive ice may be present and where the impact of

climate warming may be severe. A substantial portion of the continuous permafrost zone, therefore,

it considered to be moderately to highly sensitive to warming.

2.1.5.3 A process-based modeling approach

Another approach to predicting how permafrost will respond to future climates is through

process-based modeling. Zhang et al., [2003] combined the strength of existing permafrost models

and land surface process models to develop a physically based model of Northern Ecosystem Soil

Temperature (NEST). The use of a process-based approach allowed the effects of climate,

vegetation, ground features and hydrological dynamics to be quantified and integrated on the basis of

energy and water transfer in the soil-vegetation-atmosphere systems. The NEST model was

developed to simulate the transient response of the soil thermal regime to climate change [Zhang et

al., 2003].

The NEST model explicitly considers the effects of different ground conditions, including

vegetation, snow cover, forest floor, peat layers minerals soils and bedrock. The dynamics of soil

temperature were simulated by solving the one-dimensional heat conduction equation, with upper

boundary conditions (ground surface or snow surface when snow is present) determined based on the

17

energy balance and the lower boundary conditions (at 35 depth) defined as the geothermal flux. The

snowpack was also divided into about 0.1 cm layers, and the number of snow layers and the

thickness of the snowpack were updated every day based on snow dynamics. Heat capacities of

ground and snow layers were calculated from the specific heat capacity of liquid water, ice, organic

materials, minerals and air, weighted according to their respective volumetric fractions. The thermal

conductivity of each ground layer was calculated as the geometric mean of the thermal conductivities

of the constituents. The profile of snow density was simulated considering compaction and

destructive metamorphism for each layer. The thickness of the snowpack was determined based on

snow density and the amount of snow on the ground (water equivalent), which was the accumulative

difference between snowfall and snowmelt. Snowmelt, sublimation, and evapotranspiration were

determined based on surface energy balance. Soil water dynamics were simulated considering water

input (rainfall and snowmelt), output (evaporation and transpiration) and distribution among soil

layers. The effects of thawing or freezing on soil temperature as well as the fractions of ice and

liquid water in a soil layer were determined based on energy conservation: latent heat released or

absorbed during freezing or thawing equals the amount of heat required or released for the apparent

temperature (soil temperature determined by heat conduction equation without considering the

thawing/freezing effects) change of the layer. The depth of thawing or freezing front was determined

based on the fractions of liquid water and ice in soil layers. Thus, the model integrated the effects of

atmospheric climate, vegetation, and ground strata (snow, forest floor, peat layers, mineral soils, and

bedrock) on soil thermal dynamics based on energy and water transfer in soil-vegetation-atmosphere

systems. The model was validated against measurements of energy fluxes, snow depth, soil

temperature and thaw depth. Detailed description and validation of the model can be found in Zhang

et al., [2003]

Inputs to the NEST model include information about vegetation (land cover types, leaf area

index), ground conditions (thickness of organic layers, texture of the mineral soils, SOC content in

mineral soils, ground ice content, and the geothermal flux) and atmospheric climate (air temperature,

precipitation, solar radiation, vapor pressure, and wind speed). Vegetation types (coniferous forest,

deciduous forest, mixed forest, crop/grass land, and shrub/tundra) were determined based on the land

cover map of Canada derived from the images of the Advanced Very High Resolution Radiometer

(AVHRR) [Cihlar et al., 1999]. Leaf area index (LAI) and its seasonal variation were derived from

AVHRR 10-day composition images [Chen et al., 2002]. To match the spatial resolution of climate

18

data, the 1 km resolution land cover map and LAI images were aggregated to 0.5° latitude/longitude

based on the dominant vegetation and average LAI in each pixel, respectively. Water bodies were

excluded in the calculation. Changes in vegetation types and LAI from year to year were not

considered due to lack of long-term data. Vegetation related parameters (i.e., height and wood

biomass) were estimated based on LAI and vegetation type. Soil texture, bulk density and organic

content were extracted from the soil landscape database of Canada [Shields, 1991; Tarnocai and

Lacelle, 1996]. Soil data were aggregated to 0.5° latitude/longitude grid cells based on the dominant

type in a grid cell for soil texture, and the averages for forest floor thickness, bulk density and soil

organic carbon content. Excess ground ice often exists in regions with permafrost and was

considered during model initialization utilizing the ground ice content presented on the permafrost

map of Canada [Heginbottom et al., 1995]. The geothermal heat flux measurements of Pollack et al.,

[1993] were interpolated to 0.5° latitude/longitude resolution for the whole of Canada.

The gridded climate dataset used in this study had a 0.5° latitude/longitude spatial resolution

globally and a monthly temporal resolution from 1901 to 1995 [New et al., 2000]. This dataset

included monthly means of air temperature, the diurnal range of air temperature, water vapor

pressure, cloudiness, monthly total precipitation, and monthly total wet-days. This dataset was

interpolated from station measurements considering topographic corrections and down-scaled to half

hourly intervals to accommodate the short time-step (e.g., 15-30 minutes) required by the NEST

model [see Chen et al., 2003]. The diurnal changes in vapor pressure were not considered in the

analysis. A wind speed of 3.0 m s-1 was assumed for the simulation based on climate station

measurements. The distribution of wet-days within a month was determined as a random

distribution, and the amount of precipitation on a wet-day was determined based on an exponential

distribution with less frequency for heavier precipitation events [Hann, 1977; Richardson, 1981]. A

detailed description of the down-scaling and its validity for simulating soil temperature were

presented by Chen et al., [2003]

Output from the NEST model was validated with ground observations from four sites in

Canada by Zhang et al., [2003]. Two sites were located in Saskatchewan near the southern boundary

of the permafrost region, about 50 km northwest of Prince Albert. One of these sites was covered by

deciduous forest of aspen, while the other was covered by a coniferous forest of jack pine. Two sites

were located in Yukon Territory in the Takhini River valley, about 50 km west of Whitehorse, and

19

were within the sporadic discontinuous permafrost zone. One of these sites was located in an area of

white spruce forest that was burned in 1958. Very little regeneration of forest vegetation had

occurred at this site since the time of the fire. The other site was located in an unburned area of white

spruce. The simulation results agreed with the measurements of energy fluxes, snow depth, soil

temperature, and thaw depth. These results indicate that this physically based model captured the

effects of climate, vegetation, and ground conditions on soil temperature and freezing/thawing

dynamics, and that the model is suitable to investigate the impacts of transient climate change on soil

thermal regimes and permafrost degradation and their consequent effects on ecosystem dynamics.

Indeed, Chen et al., [2003] and Zhang et al., [in review] used the NEST model to simulate the

changes in soil temperature across Canada during the twentieth century. While Zhang et al’s study

was carried out at a Canada-wide scale, Chen et al.’s study was limited to a region of western Canada

(115-95ºW, 50-65°N). Both studies showed that mean annual soil temperatures responded to changes

in air temperature during this time period, but that these responses were often complex. The complex

response of soil temperature to changes in air temperature has significant implications for the

impacts of climate change. Chen et al., [2003] also showed that permafrost active-layer thicknesses

increased by 79% in the isolated and sporadic discontinuous permafrost zones, 37% in the extensive

discontinuous permafrost zone, and 21% in the continuous permafrost zone from 1900 to 1995.

These results are consistent with field observations, as well as the modeled results of Zhang et al., [in

review], who showed a Canada-wide annual mean soil temperature increase (at 0.2m depth) of 0.6

°C over the same time period. Chen et al., [2003] also showed that 17% of the permafrost in the

discontinuous permafrost zone was lost between 1900 and 1940, and that another 22% was lost

between 1940 and 1995. This corresponded to a northward shift in the southern limit of permafrost

of about 200km since 1900. This shift is in agreement with estimates taken from aerial photographs.

2.1.5.4 In situ monitoring of permafrost regions

The previously described observed and predicted changes in permafrost stress the necessity to

monitor its dynamics (particularly its temperature) for timely assessment and predictions of the

possible negative impacts of permafrost degradation on ecosystems and infrastructure [Romanovsky

et al., 2002]. The Global Terrestrial Network for Permafrost (GTNet-P) was established by the

International Permafrost Association (IPA) in 1999 to organize and manage a global network of

20

permafrost observatories for detecting, monitoring, and predicting climate change. The network,

authorized under the Global Climate Observing System (GCOS) and its associated organizations,

consists of two observational components. These are (a) the Circumpolar Active Layer Monitoring

(CALM) network, which was established in 1990 to monitor changes in active layer thickness and

temperature, and (b) the permafrost thermal monitoring component, which was established to

monitor ground temperatures. GTNet-P will not only provide the long-term field observations

needed for the assessment of the impacts of climate change and permafrost, but is also critical for

testing, validating and improving predictive models and bettering the reliability of impact

assessments [Burgess et al., 2001]. The GTNet-P program, its goals and establishment, activities,

progress, and planned future goals are described in more detail by Burgess et al., [2000].

Because one third of the permafrost regions of the Northern Hemisphere lie within Canada,

the international community looks to Canada for leadership in climate related observations of

permafrost. Canada is indeed taking an active and proactive role at both the national and

international level. Through the Geological Survey of Canada (GSC), Canada is involved in the

development and implementation of GTNet-P. The GSC has identified 100 active layer and /or

thermal monitoring sites identified which are or may contribute to the GTNet-P. Sixty active layer

monitoring sites are located in the Mackenzie Valley/Delta, approximately 40 of which have air and

ground surface temperature sensors. An effective national and international monitoring strategy will

provide long-term field observations essential for the detection of the terrestrial climate signal and

for the assessment of its impact on permafrost, as well as indications of the spatial variability across

the Arctic.

2.1.6 Consequences of permafrost responses to a changing climate

2.1.6.1 The physical response of permafrost terrain

The physical response of the terrain to permafrost degradation is mainly dependent on the ice

content of the frozen material [Dyke et al., 1997]. Where ice-rich materials are present, an increase in

thaw-settlement and thermokarst activity will probably accompany climate warming. Soil strength

due to ice bonding will be reduced as unfrozen water content of the frozen ground increases in a

response to rising air temperature. This may lead to ground instability and an increased incidence in

slope failure. An increase in the frequency of wild fires may accompany climate warming. In

21

peatland areas, fires could consume some of the dry peat in the active layer of peat plateaus,

initiating widespread permafrost degradation [Zoltai, 1993] and thaw settlement. An increase in

active-layer detachment slides may also result from increased fire frequency and burn extent [Dyke,

2000].

2.1.6.2 Responses of groundwater, river and lake systems

Climate warming may have important effects on the hydrology of permafrost areas.

Permafrost provides an impermeable layer that impedes drainage, supports a high water table, and

constrains infiltration and groundwater movement to the active layer [Mackay and Loken, 1974].

Wetlands and ponds are therefore common in the permafrost region [Wright, 1979]. Patchy wetlands

may even exost in the polar desert of the high arctic because of the shallow active-layer and high

water table [Woo and Young, 1998]. Increases in active-layer thickness will improve drainage and

may lead to a loss of these wetlands. This can result in a change of vegetation patterns and the

potential loss of breeding habitats for wildlife [Michel and van Everdingen, 1994].

Under climate warming, groundwater will play a more important role in hydrological and

landscape processes, especially in areas currently underlain by continuous permafrost [Michel and

van Everdingen, 1994]. Frost heave in the active layer may increase due to a greater availability of

unfrozen water and this has important engineering implications. Frost blisters, which are formed

when frost heave occurs, may become more numerous in the arctic region. Icing activity, which can

present a serious road hazard, may also increase in the continuous permafrost zone. Greater

exchanges between surface water and groundwater may lead to a greater dissolved solids content in

rivers which may have an adverse affect on fish and other aquatic life. As permafrost thaws,

increased regional groundwater flow may promote further warming and thawing of permafrost.

Groundwater may also be discharged offshore (through the sea floor) where it may influence near

shore circulation and sea-ice cover.

A warming climate will also affect the way rivers in permafrost environments will respond to

snowmelt and rainfall events. Rivers usually exhibit a quick response to these events where

permafrost is present. In these regions, the active layer is easily saturated and most of the water

reaches streams as overland flow [Woo, 1976]. Drainage basins in the permafrost region therefore

have high runoff-to-rainfall ratios [Kane et al., 1998; Lilly et al., 1998]. Once the precipitation event

22

is over, however, stream flow quickly decreases because permafrost restricts groundwater flow to the

stream [Dingman, 1973]. As permafrost degrades and active layers thicken, subsurface flow will

become a more important contribution to baseflow, and streamflow will become more uniform

throughout the year [Ashmore and Church, 2001; Michel and van Everdingen, 1994; Woo, 1976]. An

unfrozen zone (talik) may develop between the base of the active layer and the permafrost table

allowing for streamflow to be sustained in winter [Hinzman and Kane, 1992]. However, enhanced

winter streamflow may result in more extensive ice-river formation and the possibility of more

serious flooding during break-up in the arctic river ice [Michel and van Everdingen, 1994]. A deeper

active layer will be associated with a greater variation in the amount of water stored in the soil as

well as an increase in the movement of subsurface water downslope [Hinzman and Kane, 1992].

2.1.6.3 Potential challenges to northern development

2.1.6.3.1 Freezing and thawing of soils

Potential challenges to northern development resulting from permafrost warming are

described at length by Couture et al., [2000], Robinson et al., [2001] and Smith et al., [2001]. In

general, the problems encountered during construction and development in permafrost regions are

classified as (a) those involving the thawing of ice-rich materials under unheated structures such as

roads, or heated structures such as buildings, (b) those resulting from resulting from frost action,

such as frost heave, and (c) those associated with the temperature of permafrost, such as the freezing

of buried water lines [Smith et al., 2001]. Since constraints to development in the north are often

associated with the melting of ground ice or freezing of soil water, it is important to understand what

happens as soils freeze and thaw. Thus, knowledge of the mechanical and physical properties of

freezing and thawing soils is also required to properly design structures for cold regions [Smith et al.,

2001].

As soil temperature falls below 0°C, water within the soil’s pores freeze gradually over a

range of temperatures below 0°C. Ice lenses may develop within the frozen soil as water accumulates

and freezes. As these lenses grow, the ground surface is displaced upwards, resulting in frost heave.

The amount of frost heave experience in any location depends on soil water conditions, the rate of

cooling of the soil, soil texture and overburden pressure. Fine-grained, silty soils are generally more

23

susceptible to frost heave. Frost heave may occur during permafrost formation or during seasonal

freezing. It may also occur beneath structures such as chilled pipelines. If conditions are uniform

beneath the entire structure, then the movement is uniform and can be tolerated by the structure.

However, if soil and moisture conditions are highly variable, differential movement can occur. This

can result in damages to structures such as roads, airstrips, buildings, buried utility lines and

pipelines [Smith et al., 2001].

As soil temperature rises towards 0°C, ice within the soil melts. The thawing of frozen soils

has a number of effects on the physical properties of soils. First, the strength of fine-grained soils

depends on the thermal characteristics of the soil. When frozen, such soils are strong because ice-

bonding contributes significantly to their overall strength. However, as these soils thaw, their

strength diminishes, then disappears as the soil thaws completely. Second, the strength of soil can be

further weakened where it contains excess ice in the form of lenses or bodies of massive ice. This is

because of an increase in pore-water pressures and a decrease in strength if excess water is unable to

drain from the soil profile. As the water eventually drains away, that settlement occurs, where the

volume formerly occupied by the excess ice is lost. Fine-grained soils that generally have high ice

contents (silt, clay, organic material) are generally the most susceptible to this process. Differential

thaw settlement produces irregular terrain called thermokarst topography. The potential development

hazard due to thaw settlement is thus greatest in regions of ice-rich permafrost. A substantial part of

the region containing permafrost where ground temperature is > -2°C has a high potential for thaw

instability [Smith et al., 2001].

2.1.6.3.2 Slope stability

The warming of ice-rich soils and melting ground-ice on slopes can cause instability resulting

in flows and slides. A flow is a landslide in which movement of debris has characteristics of a fluid.

A slide is a landslide movement in which the slide components move downslope as a block or series

of blocks. Flows include shallow active-layer detachments, deeper retrogressive thaw flows and

rapid debris flows. Active layer detachments are shallow slope failures involving the detachment and

downslope movement of ground within the active layer and associated vegetation mat. They may be

triggered by unusually warm temperatures or a disturbance to the vegetation cover, such as wildfire

or clearing for construction. Deeper retrogressive thaw flows occur in ice-rich terrain and consist of a

24

low-angle tongue of debris that extends downslope from a steep headwall. Debris flows occur in

regions of high relief, and are a rapid, destructive flow of water-saturated sediment. They may be

triggered by extreme precipitation events or unusually deep seasonal thaw in response to extremely

warm summer temperatures [Smith et al., 2001]. Slides are commonly triggered by the undercutting

of river banks, which results in enough loading to overcome the cohesion of frozen sediments. As ice

melts within the slide, it may deform and flow. Approximately 2,000 slides have been documented in

the Mackenzie Valley, with a further 1,000 retrogressive thaw-flows identified in the Mackenzie

Delta-Tuktoyaktuk Peninsula area. Landslide and flow events can cause direct damage to critical

infrastructure, including pipelines, roads, bridges and buildings. They can also indirectly damage

infrastructure by creating landslide dams that interrupt river navigation or cause flooding. They can

also increase siltation in streams that may affect fisheries.

2.1.6.3.3 Problems associated with development in the permafrost region

Ground instability is thus a major concern for development in permafrost regions, especially

in regions where permafrost temperatures are near 0°C. The removal of vegetation or insulating

organic cover and other disturbances to the ground surface can result in changes to the ground

thermal regime, leading to the warming and melting of permafrost. In addition, permafrost can be

warmed or melted by heat generated from heated buildings or buried water and sewage pipelines.

In areas of ice-rich, fine-grained sediment, failure to take proper precautions can have serious

consequences for development. For example, construction in Dawson (Yukon Territory) and Aklavik

(NWT) ignored the ice-rich soils underlaying these sites. In Dawson, buildings were similar in

design and construction to those in southern towns. The subsequent melting of ground ice and

differential settlement of the underlying ground damaged buildings, many of which became

uninhabitable. Roads became impassable because of differential settlement as underlying permafrost

thawed. Similar problems occurred in Aklavik, where buildings, roads, trails and other structures

became unusable because of ground subsidence due to permafrost thaw.

Because the presence of permafrost was often not recognized, construction in permafrost

terrain often used inappropriate practices. The consequences of these practices were later accepted or

addressed. Current engineering practices are more sophisticated and aim to minimize disturbance to

25

the terrain as well as impacts on structures. Modern northern infrastructure is designed to preserve

thaw-sensitive permafrost and therefore limit thaw settlement, located to avoid that-sensitive and

frost-sensitive soils, or designed to withstand the anticipated thaw settlement or frost heave. A

lengthy and detailed discussion of infrastructure design in permafrost landscapes (e.g. site

investigations; transportation, buildings and foundations, utilities, pipelines, dams and dykes) is

provided elsewhere [Smith et al., 2001].

2.1.6.4 Case studies from Canada’s permafrost zone

Natural Resources Canada (Geological Survey of Canada and Canada Centre for Remote

Sensing) has initiated community case studies examining infrastructure sensitivity to the impacts of

permafrost degradation under climate warming. Two separate reports provide a summary of present

permafrost conditions, surficial geology, and infrastructure conditions in Norman Wells [Robinson et

al., 2001] and Tuktoyaktuk [Couture et al., 2000], Northwest Territories. These reports include data

compilation, reviews of the communities, climatic conditions, terrain and permafrost conditions, and

infrastsructure facilities. They also provide a brief discussion on how climate change could affect the

infrastructure in these communities, and the possible costs associated with adaptation strategies. A

third study [Hoeve et al., in review] assesses the impacts of melting permafrost on community

infrastructure in 32 communities in the Northwest Territories. The aims of this study were to identify

and characterize the adverse impacts of climate change on building foundation systems in the

Northwest Territories, to which adaptation would be required. These documents form the basis of the

following discussion.

Long-term records from Norman Wells indicate that air temperatures in this region increased

by 1.3 °C from 1944 to 1998. A slightly higher increase (2 °C) was observed in Tuktoyaktuk over a

slightly shorter time period (1958 to 1998). These increases in are most striking in the winter and

spring seasons at both locations. Ground temperature measurements from a Norman Wells and

Tuktoyaktuk {source?} largely parallel the observed increase in air temperature. This warming trend

is predicted to continue. According to GCM results from the Canadian Centre for Climate Modelling

and Analysis (CCCma), winter warming in Norman Wells and Tuktoyaktuk is expected to increase

by 3 °C and 2.6 ºC, respectively (8 °C and 8.3 °C by 2080, respectively). In both cases, the

26

anticipated warming in other seasons is not as dramatic, yet it is nonetheless highly significant (4.5

°C by 2080 in Norman Wells; 6 °C by 2080 in Tuktoyaktuk).

The predicted changes in air temperature are expected to have significant effects on the

Norman Wells and Tuktoyuktuk communities. Much of the permafrost in these regions lies between

–2 and 0 °C. As a result, even small increases in air temperature will likely reduce the extent of

permafrost, increase the depth of the active layer, and cause melting of ground ice [Burgess et al.,

2004]. This warming will likely lead to more frequent geotechnical problems, causing important, and

certainly costly, problems or maintenance for infrastructure. Climate warming will likely affect the

performance of existing infrastructure and shorten its operating life. The design or maintenance of

future infrastructure should incorporate the potential for significant alterations to permafrost owing

to climate warming. Theses considerations are discussed in detail by Robinson et al., [2001] and

Couture et al., [2000].

More recently, Hoeve et al., [in review] assessed the impacts of melting permafrost on

community infrastructure in 32 communities in the Northwest Territories. This study used a multiple

accounts analysis (MAA) was used to rank the sensitivity of community building infrastructure to

climate change. Three factors – thermal sensitivity, physical sensitivity, and infrastructure sensitivity

– were considered in assessing the vulnerability of building foundation system to climate changes.

Thermal sensitivity is determined by the combination of climate trend and existing permafrost

conditions, physical sensitivity is determined by the susceptibility of the terrain to movement caused

by permafrost thawing, and infrastructure sensitivity is determined by vulnerability of building

foundations to damage as a result of climate change and to the ease with which adaptation measures

can be implemented.

The results of this study were that in the NWT (a) thermal, physical, and infrastructure

sensitivities exert significant effect on the sensitivity of building foundations to climate change; (b)

thermal and physical sensitivity are closely but inversely correlated, indicating that thermally

unfavourable conditions tend to be collocated with physically favourable conditions (and particularly

ground ice presence); (c) physical sensitivity and infrastructure sensitivity are weakly positively

correlated, and various infrastructure indicators are influencing this relationship; and (d) construction

practices (adfreeze piles) were a significant indicator for the thermal sensitivity-infrastructure

27

sensitivity relationship, suggesting a tendency for

thermally unfavourable conditions to be

compensated for by building foundation practices.

However, the relationship was generally weak.

The spatial pattern of the multiple account

scores is shown in Figure 6. For this purpose, the

communities have been grouped into three

categories of relative sensitivity (high, moderate and

low) by dividing the overall range of computed

values into thirds. These scores show that The

spatial pattern of MAA scores for each of the

communities in this study (Figure 6) shows that the

most sensitive communities are in the NW part of

the NWT due to high physical sensitivity (buildings

supported on relatively warm, ice- rich permafrost),

enhanced in some communities by high infrastructure sensitivity due to construction practices or

adaptation challenges. This pattern also shows that the communities with the least sensitivity are also

located in the south; these have moderate to high thermal sensitivities but generally low physical or

infrastructure sensitivities because relatively few buildings are supported on permafrost in this

region.

This study demonstrates that the MAA approach appears to be an appropriate tool for an

initial analysis of the sensitivity of infrastructure. It allows consideration of the most important

environmental and design variables that are likely to influence the response. It can also take

advantage of the level of understanding of the processes which may vary depending on the aspect of

the problem and the availability of data. It allows using available data in a way that reflects the

importance and likely accuracy of the data. The results may be used to focus attention to areas

identified as likely ‘hot spots’, and may be followed by more detailed investigations, modeling, field

studies, or consideration of remedial action. When used systematically, the MAA procedure can

avoid most of potential subjective bias, as it permits the initially identified indicators, their assigned

values, and selected account weights to characterize the overall problem.

Figure 6. MAA scores for communities considered in study [after Hoeve et al., in review].

28

2.1.6.5 Potential feedbacks to the global carbon cycle

Climate warming – and the predicted subsequent thawing of permafrost – may have

important feedback effects on the global carbon cycle. Terrestrial carbon sinks are an important, yet

poorly understood component of the global CO2 budget [Sundquist, 1993]. Northern peatlands are a

significant global carbon sink and it is estimated that 200 Gt to 500 Gt carbon are sequestered within

them [Kuhry and Zoltai, 1994]. The Canadian permafrost region currently represents an important

carbon sink, which may become a carbon source as climate warming progresses. Tarnocai [1998]

estimates that about 104 Gt of carbon are stored in soils (cryosols) within the permafrost region of

Canada, and that a substantial portion (47 Gt) of this is stored in frozen organic soils found mostly

within peatlands in western Canada and the Mackenzie Valley [Vitt et al., 1994]. Carbon that was

sequestered in organic matter during a warmer period has been preserved in the frozen material

[Oechel et al., 1995; Peteet et al., 1998; Vardy et al., 1973]. The amount of carbon stored in the

permafrost regions will likely change in response to climate warming, but the response will be

complex [Moore et al., 1998]. A change from a carbon sink to carbon a source has been suggested,

but this will not be caused directly by an increase in temperature. An increase in active-layer

thickness, enhanced drainage and soil aeration, and a decrease in water-table level will favour peat

decomposition and the release of CO2 to the atmosphere, constituting a potentially significant

positive feedback leading to an additional climate warming [Bubier et al., 1998; Oechel et al., 1995;

Schraeder et al., 1998]. Following an initial outgassing of CO2, an increase in primary productivity

related to tree and shrub growth may lead to a general increase in carbon storage [Oechel et al.,

1995; Waelbroeck et al., 1997]. If increased precipitation or poor drainage lead to higher water

tables, methane, which is also a greenhouse gas, may be released from northern peatlands [Roulet et

al., 1992]. The thawing of permafrost in dry plateaus will create collapsed bogs and fens which will

emit additional amounts of methane [Moore et al., 1998]. An increase in the incidence and severity

of fire would also modify the carbon balance of permafrost-affects peatlands. The role of fire in the

peatland carbon cycle, however, is poorly understood.

Large amounts of methane may currently be stored within and beneath permafrost as natural

gas hydrate bodies. These are ice-like solids that are stable under conditions of low temperature and

high pressure. Warming of terrestrial and marine sediments in the Mackenzie Delta-Beaufort Sea

29

region may result in gas-hydrate decomposition and the eventual release of methane to the

atmosphere, further enhancing climate warming [Kvenvolden, 1988; Kvenvolden, 1994; Nisbet,

1989]. Methane hydrate is generally only stable at depths greater than 200m, and hence, at locations

where permafrost thickens exceeds 200m. A substantial amount of time (hundreds of years) may be

required for gas-hydrate-bearing sediments to become sufficiently warm for decomposition to take

place. As permafrost thaws, however, potential increases in regional groundwater circulation would

further enhance the warming of deep-water sediments and degradation of gas hydrate [Michel and

van Everdingen, 1994]. Evidence also exists for the occurrence of gas hydrate at depths shallower

than the theoretical minimum determined for methane hydrate stability [Dallimore and Collet, 1995].

This gas hydrate could be affected by climate warming over a shorter time period. Permafrost may

act as an essentially impermeable barrier to the migration of hydrocarbons from below. Permafrost

thaw and the formation of taliks may create conduits for gases such as methane to reach the surface

and be released to the atmosphere.

2.1.7 References

Allard, M., B. Wang, and J.A. Pilon, Recent cooling along the southern shore of Hudson Strait Quebec, Canada, documented from permafrost temperature measurements, Arctic and Alpine Research, 27, 157-166, 1995.

Allen, D., F. Michel, and A. Judge, Paleoclimate and permafrost in the Mackenzie delta, in Proceedings of the fifth annual conference on permafrost, pp. 33-38, Trondheim, Norway, 1988.

Ashmore, P., and M. Church, The impact of climate change on rivers and river processes in Canada, pp. 58, Natural Resources Canada, Government of Canada, Ottawa, Ontario, Canada, 2001.

Blasco, S.M., G. Fortin, P.R. Hill, M.J. O'Connor, and J. Brigham-Grette, The late Neogene and Quaternary stratigraphy of the Canadian Beaufort continental shelf, in The artic ocean region, edited by A. Grantz, L. Johnson, and J.F. Sweeney, pp. 491-502, Geological Society of America, Boulder, Colorado, USA, 1990.

Briggs, D., P. Smithson, T. Ball, P. Johnson, P. Kershaw, and A. Lewkowicz, Fundamentals of Physical Geography, 692 pp., Copp Clark Pitman Ltd, Toronto, Ontario, Canada, 1993.

Bubier, J.L., P.M. Crill, T.R. Moore, K. Savage, and R.K. Varner, Seasonal patterns and controls on net ecosystem CO2 exchange in a boreal peatland complex, Global Biogeochemical Cycles, 12, 703-714, 1998.

Burgess, M.M., D.T. Desrochers, and R. Saunders, Potential changes in thaw depth and thaw settlement for locations in the Mackenzie Valley, in Physical environment of the Mackenzie Valley, Northwest Territories: a base line for the assessment of environmental change, edited

30

by L.D. Dyke, and G.R. Brooks, pp. 187-195, Natural Resources Canada, Government of Canada, Ottawa, Ontario, Canada, 2004.

Burgess, M.M., D.W. Riseborough, and S.L. Smith, Permafrost and glaciers/icecaps monitoring networks workshop: Report on the permafrost sessions, Geological Survey of Canada, Ottawa, ON, Canada, 2001.

Burgess, M.M., and S.L. Smith, Shallow ground temperatures, in The physical environment of the Mackenzie Valley, North West Territories: a base line for the assessment of environmental change, edited by L.D. Dyke, and G.R. Brooks, pp. 89-103, Geological Survey of Canada, Ottawa, Ontario, Canada, 2000.

Burgess, M.M., S.L. Smith, J. Brown, V. Romanovsky, and K. Hinkel, Global Terrestrial Network for Permafrost (GTNet-P): permafrost monitoring contributing to global climate observations, pp. 8, Geological Survey of Canada, Ottawa, Ontario, Canada, 2000.

Burn, C.R., The response (1958-1997) of permafrost and near-surface ground temperatures to forest fire, Canadian Journal of Earh Sciences, 35, 184-199, 1998.

Burn, C.R., F.A. Michel, and M.W. Smith, Stratigraphic, isotopic, and mineralogical evidence for an early Holocene thaw unconformity at Mayo, Yukon Territory, Canadian Journal of Earh Sciences, 23, 794-803, 1986.

Chen, J.M., G. Pavlic, L. Brown, J. Cihlar, S.G. Leblanc, H.P. White, R.J. Hall, D.R. Peddle, D.J. King, J.A. Trofymow, E. Swift, J. van der Sanden, and P. Pellikka, Validation of Canada-wide leaf area index maps using ground measurements and high and moderate resolution satellite imagery, Remote Sensing of Environment, 80 (1), 165-184, 2002.

Chen, W., Y. Zhang, J. Cihlar, S.L. Smith, and D.W. Riseborough, Changes in soil temperature and active layer thickness during the twentieth century in a region of western Canada, Journal of Geophysical Research, 108 (D22), 4696, 2003.

Cihlar, J., R. Beaubien, R. Latifovic, and G. Simard, Land Cover of Canada version 1995, Version 1.1 Digital Data Documentation, in Special Publication, NBIOME Project, Produced by the Canada Centre for Remote Sensing and the Canadian Forest Service, Natural Resources Canada, Ottawa, Ontario, Canada., 1999.

Couture, R., S.D. Robinson, and M.M. Burgess, Climate change, permafrost degradation, and infrastructure adaptation: preliminary results from a pilot community case study in the Mackenzie valley, pp. 9, Geological Survey of Canada, Ottawa, Ontario, Canada, 2000.

Dallimore, S.R., and T.S. Collet, Intra permaforst gas hydrates from a deep core hole in the Mackenzie Delta, Northwest Territories, Canada, Geology, 23, 527-530, 1995.

Dingman, S.L., Effects of permafrost on stream flow characteristics in the discontinuous permafrost zone of central Alaska, in Proceedings of the 2nd International Conference on Permafrost, North American contribution,, pp. 447-453, National Academy of Sciences, Washington, D.C., USA, 1973.

Dyke, L.D., Stability of permafrost slopes in the Mackenzie Valley, in The physical environment of the Mackenzie Valley, Northwest Territories: A baseline for the assessment of environmental change, edited by L.D. Dyke, and G.R. Brooks, pp. 177-186, Geological Survey of Canada, Ottawa, Ontario, Canada, 2000.

31

Dyke, L.D., J.M. Aylsworth, M.M. Burgess, F.M. Nixon, and F. Wright, Permafrost in the Mackenzie Basin, its influences on land-altering processes, and its relationship to climate change, in Mackenzie Basin Impact Study (MBIS), pp. 112-117, Environmental Adaptation Research Group, Ottawa, Ontario, Canada, 1997.

Flato, G.M., G.J. Boer, W.G. Lee, N.A. McFarlane, D. Ramsden, M.C. Reader, and A.J. Weaver, The Canadian Centre for Modelling and Analysis Global Coupled Model and its climate, Climate Dynamics, 16, 451–467, 2000.

French, H.M., and I.E. Egerov, 20th century variations in the southern limit of permafrost near Thompson, northern Manitoba, in International Conference on Permafrost, pp. 297-304, Université Laval, Québec, Canada, Yellowknife, Canada, 1998.

Halsey, L.A., D.H. Vitt, and S.C. Zoltai, Disequilibrium response of permafrost in boreal continental western Canada to climate change, Climate Change, 30, 57-73, 1995.

Hann, C.T., Statistical Methods in Hydrology, Iowa State University Press, Ames, Iowa, USA, 1977.

Heginbottom, J.A., M.A. Dubreuil, and P.A. Harker, Canada Permafrost, in National Atlas of Canada, Natural Resources Canada, Ottawa, Canada, 1995.

Hinzman, L.D., and D.L. Kane, Potential response of an arctic watershed during a period of global warming, Journal of Geophysical Research, 97 (D3), 2811-2820, 1992.

Hoeve, T.E., F. Zhou, A. Zhang, and J. Cihlar, Assessment of building foundation sensitivity to climate change in the Northwest Territories, Canada, Mitigation and Adaptation Strategies for Global Environmental Change, in review.

Intergovernmental Panel on Climate Change, Climate change 1995. Impacts, adaptations and mitigation of climate change: Scientific-Techincal Analysis Contributions of Working Group II to the second assessment report of the IPCC, 876 pp., WMO-UNEP, Geneva, Switzerland, 1996.

International Permafrost Association, Mult-language glossary of permafrost and related ground ice terms, version 2.0, 278 pp., University of Calgary, Calgary, Canada, 1998.

Johnson, L., and L. Viereck, recovery and active layer changes following a tundra fire in north-western Alaska, in Proceeding of the fourth International Permafrost Conference, pp. 543-544, National Academy Press, Washington, D.C., USA, 1983.

Kane, D.L., D.J. Soden, L.D. Hinzman, and R.E. Gieck, Rainfall runoff of a nested watershed in the Alaskan arctic, in Proceedings of the 7th International Conference on Permafrost, North American contribution,, pp. 539-543, Université Laval, Québec, Canada, Yellowknife, NWT, Canada, 1998.

Kettles, I.M., C. Tarnocai, and S.D. Bauke, Predicted permafrost distribution in Canada under a climate warming scenario, in Current Research 1997-E, pp. 109-115, Geological Survey of Canada, Ottawa, Ontario, Canada, 1997.

Kuhry, P., and S.C. Zoltai, Past climatic change and the development of peatlands: an introduction, Journal of Paleoliminology, 12, 1-2, 1994.

Kvenvolden, K.A., Methane hydrates and global climate, Global Biogeochemical Cycles, 2, 221-229, 1988.

32

Kvenvolden, K.A., Natural gas hydrate occurrence and issues, in International Conference on Natural Gas Hydrates, edited by E.D. Sloan, J. Happel, and M.A. Hnatow, pp. 232-246, Annals of the New York Academy of Sciences, 1994.

Kwong, Y.T., and T.Y. Gan, Northward migration of permafrost along the Mackenzie Highway and climate warming, Climate Change, 26, 399-419, 1994.

Laberge, M.-J., and S. Payette, Long-term monitoring of permafrost change in a palsa peatland in northern Quebec, Canada, Arctic and Alpine Research, 27, 167-171, 1995.

Lachenbruch, A.H., and B.V. marshall, Changing climate: geothermal evidence from permafrost in the Alaskan arctic, Science, 234, 689-696, 1986.

Liang, I.L., J.W. Zhou, and J.C. Wang, Changes to the permafrost environment after forest fire, Da Xi'an ridge, Gu Lian mining area, China, Permafrost and Periglacial Processes, 2, 253-257, 1991.

Lilly, E.K., D.L. Kane, L.D. Hinzman, and R.E. Gieck, Annual water balance for three nested watersheds on the north slope of Alaska, in Proceedings of the 7th International Conference on Permafrost, North American contribution,, pp. 669-674, Université Laval, Québec, Canada, Yellowknife, NWT, Canada, 1998.

Mackay, D.K., and O.H. Loken, Arctic Hydrology, in Arctic and Alpine Environments, edited by J.D. Ives, and R.G. Barry, pp. 111-132, Methuen and Co. Ltd, London, UK, 1974.

Maxwell, B., Responding to global climate change in Canada's arctic. Volume II of the Canada Country study:Climate impacts and adaptation, pp. 82, Environemntal Adaptation Research Group, Atmospheric Environment Service, Environment Canada, 1997.

Michel, F.A., and R.O. van Everdingen, Changes in hydrogeologic regimes in permafrost regions due to climate change, Permafrost and Periglacial Processes, 5, 191-195, 1994.

Moore, T.R., N.T. Roulet, and J.M. Waddington, Uncertainty in predicting the effect of climatic change on the carbon cycling of Canadian peatlands, Climatic Change, 40, 229-245, 1998.

New, M.G., M. Hulme, and P.D. Jones, Representing twentieth century space-time climate variability, Part II: Development of a 1901-1996 monthly terrestrial climate fields, Journal of Climate, 13, 2217-2238, 2000.

Nisbet, E.G., Some northern sources of atmospheric methane: production, history and future implications, Canadian Journal of Earh Sciences, 26, 1603-1611, 1989.

Oechel, W.C., G.L. Vourlitis, S.J. Hasings, and S.A. Bochkarev, Change in arctic CO2 flux over two decades: effects of climate change at Barrow, Alaska, Ecological Applications, 5, 846-855, 1995.

Peteet, D., A. Andreev, W. Bardeen, and F. Mistretta, Long-term arctic peatland dynamics, vegetation and climate history of the Par-Taz region, Western Siberia, Boreas, 27, 115-126, 1998.

Pollack, H.N., S.J. Hurter, and J.R. Johnson, Heat flow from the earths interior: analysis of the global data set, Reviews of Geophysics, 31, 267-280, 1993.

Richardson, C.W., Stochastic simulation of daily precipitation, temperature, and solar radiation, Water Resources Research, 17, 182-190, 1981.

33

Robinson, S.D., R. Couture, and M.M. Burgess, Climate change, permafrost and community infrastructure, pp. 68, Geological Survey of Canada, Ottawa, Ontario, Canada, 2001.

Romanovsky, V., S.L. Smith, K. Yoshikawa, and J. Brown, Permafrost temperature records: Indicators of climate change, EOS, Transactions of the American Geophysical Union, 83 (50), 593-594, 2002.

Roulet, N., T.R. Moore, J.L. Bubier, and P. Lafleur, Northern fens: methane flux and climatic change, Tellus, 44B, 100-105, 1992.

Schraeder, C.P., W.R. Rouse, T.J. Griffist, L.D. Boudreau, and P.D. Blenken, Carbon dioxide fluxes in a northern fen during a hot dry summer, Global Biogeochemical Cycles, 12, 729-740, 1998.

Shields, J.A., C. Tarnocai, K. W. G. Valentine, and K. B. MacDonald,, Soil Landscapes of Canada - Procedures Manual and Users Handbook, pp. 74, Land Resource Research Centre, Research Branch, Agriculture Canada, Ottawa, Canada, 1991.

Smith, S.L., The significance of climate change for the permafrost environment, in Proceedings of the fifth annual conference on permafrost, pp. 18-23, Trondheim, Norway, 1988.

Smith, S.L., and M.M. Burgess, Mapping the response of permafrost in Canada to climate warming, in Current Research 1998-E, pp. 163-171, Geological Survey of Canada, Ottawa, Ontario, Canada, 1998.

Smith, S.L., and M.M. Burgess, Mapping the sensitivity of Canadian permafrost to climate warming, in Interactions Between the Cryosphere, Climate and Greenhouse Gases (Proceedings of IUGG 99 Symposium HS2); IAHS Publication no. 256, pp. 71-80, Birmingham, UK, 1999.

Smith, S.L., and M.M. Burgess, Sensitivity of permafrost to climate warming in Canada, pp. 24, Natural Resources Canada, Government of Canada, Ottawa, Ontario, Canada, 2004.

Smith, S.L., M.M. Burgess, and J.A. Heginbottom, Permafrost in Canada, a challenge to northern development, in A Synthesis of Geological Hazards in Canada, edited by J.R. Brooks, pp. 241-264, Natural Resources Canada, Government of Canada, Ottawa, Ontario, Canada, 2001.

Smith, S.L., M.M. Burgess, D. Riseborough, and F.M. Nixon, Recent trends from Canadian permafrost thermal monitoring network sites, Permafrost and Periglacial Processes, 16, 1-13, 2005.

Summerfield, M.A., Global geomorphology. An introduction to the study of landforms, 537 pp., Longman Scientific and Technical, New York, NY, USA, 1991.

Sundquist, E.T., The global carbon dioxide budget, Science, 259, 934-941, 1993.

Tarnocai, C., The amount of organic carbon in various soil orders and ecological provinces in Canada, in Soil processes and the carbon cycle, edited by R. Lal, J.M. Kimble, R.L.F. Follett, and B.A. Stewart, pp. 81-92, CRC Press, New York, NY, USA, 1998.

Tarnocai, C., and B. Lacelle, Soil Organic Carbon Digital Database of Canada, Eastern Cereal and Oilseed Research Centre, Research Branch, Agriculture and Agro Food Canada, Ottawa, Canada, 1996.

34

Tarnocai, C., F.M. Nixon, and L. Kutny, Circumpolar-Active-Layer-Monitoring (CALM) sites in the Mackenzie valley, Northwestern Canada, Permafrost and Periglacial Processes, 15, 141-153, 2004.

Taylor, A.E., Marine transgression, shoreline emergence: evidence in seabed and terrestrial ground temperatures of changing relative sea levels, Arctic Canada, Journal of Geophysical Research, 96 (B4), 6893-6909, 1991.

Trenhaile, A.S., The geomorphology of Canada: An introduction, 240 pp., Oxford University Press, Toronto, Ontario, Canada, 1990.

Vardy, S.R., B.G. Warner, and R. Aravena, Holocene climate effects on the development of a peatland on the Tuktoyaktuk Peninsula, Northwest Territories, Journal of Quaternary Research, 47, 90-104, 1973.

Vardy, S.R., B.G. Warner, and R. Aravena, Holocene climate and the development of a subarctic peatland near Inuvik, Northwest Territories, Canada, Journal of Quaternary Research, 47, 90-104, 1998.

Vitt, D.H., L.A. Halsey, and S.C. Zoltai, The bog landforms of continental western Canada in relation to climate and permafrost patterns, Arctic and Alpine Research, 26, 1-13, 1994.

Waelbroeck, C., P. Monfray, W.C. Oechel, S. Hastings, and G. Vourlitis, The impact of permafrost thawing on the varbon dynamics of tundra, Geophysical Research Letters, 24, 229-232, 1997.

Williams, P.J., and M.W. Smith, The Frozen Earth: Fundamentals of Geocryology, Cambridge University Press, Cambridge, UK., 1989a.

Williams, P.J., and M.W. Smith, The frozen earth: fundamentals of geocryology, 306 pp., Cambridge University Press, Cambridge, UK, 1989b.

Woo, M.-K., Hydrology of a small Canadian high arctic basin during the snowmelt period, Catena, 3, 155-168, 1976.

Woo, M.-K., and K.L. Young, Characteristics of patchy wetlands in a polar desert environment, in Proceedings of the 7th International Conference on Permafrost, North American contribution,, pp. 1141-1146, Université Laval, Québec, Canada, Yellowknife, NWT, Canada, 1998.

Wright, R.K., Preliminary results of a study on active layer hydrology in the discontinuous zone at Schefferville, Nouveau-Quebec, Géographie Physique et Quaternaire, 49, 45-54, 1979.

Zhang, T., R.G. Barry, K. Knowles, J.A. Heginbottom, and J. Brown, Statistics and characteristics of permafrost and ground ice distribution in the Northern Hemisphere, Polar Geography, 23 (2), 147-169, 1999.

Zhang, Y., W. Chen, and J. Cihlar, A process-based model for quantifying the impact of climate change on permafrost thermal regimes, Journal of Geophysical Research, 108 (D22), 4695, 2003.

Zhang, Y., W. Chen, S.L. Smith, D.W. Riseborough, and J. Cihlar, Soil temperature in Canada during the twentieth century: complex responses to atmospheric climate change, Journal of Geophysical Research, in review.

35

Zoltai, S.C., Cyclic development of permafrost in the peatlands of northwestern Alberta, Canada, Arctic and Alpine Research, 25, 240-246, 1993.

Zoltai, S.C., Permafrost distribution in the peatlands of west-central Canada during the Holocene warm period 6000 years BP, Géographie Physique et Quaternaire, 49, 45-54, 1995.

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2.2 Glaciers

2.2.1 What are Glaciers?

Glaciers are huge masses of ice, formed on land by the compaction and re-crystallization of

snow, that move very slowly outwards and downwards under the pressure of their own weight and

gravity. Glaciers form when the snow pack accumulated in winter (or, for tropical glaciers, the wet

season) does not usually melt entirely in summer; i.e. where the annual mass balance is positive for a

number of years [Oerlemans, 2001; Siegert, 2001]. The remaining snow is then transformed into ice

to form a glacier. Glaciers grow where climatic and topographic conditions exceed the losses, and

recede where the outputs are greater than the inputs.

The form and shape of glaciers are a function of climate and topography, and the gross

morphology of any one glacier is unique to its location on the Earth's surface. Therefore, there are a

wide variety of glacier types, which form a continuum of sizes from continental-scale ice sheets to

tiny niche glaciers found in mountain hollows. While glaciers can be classified based on their surface

climate (polar, temperate, or maritime) or on the conditions at their bed (wet, warm, or melting vs.

dry, cold, or freezing), they are often classified on the basis of their relationship to the underlying

bedrock topography. This classification system yields three main glacier types:

(1) Glaciers unconstrained by topography – e.g. ice caps and ice sheets – cover vast areas

and are of sufficient thickness to submerge the underlying landscape. Topography does not

play a major role in determining the extent of these glaciers. Thus, their direction of ice flow

reflects the size and shape of the glacier rather than the shape of the ground;

(2) Glaciers partially constrained by topography – e.g. ice shelves – have their shapes and

directions of ice flow only partly influenced by the underlying terrain. Ice shelves occur

where ice is forced to float by deeper water or where sea-ice thickens by surface

accumulation and bottom accretion. These glaciers are often partially constrained by the

shape of the coastline.

(3) Glaciers constrained by topography – e.g. ice fields, cirque glaciers, valley glaciers, other

small glaciers – are strongly influenced in form and direction of ice flow by the shape of the

ground. These are the glaciers that are found in rugged topography and are typically bound

within a valley or depression.

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2.2.2 Factors influencing the distribution of glaciers

Glaciers are located wherever topographic and climatic factors are suitable for snow to

accumulate and persist from year to year. The potential for perennial snow cover is primarily

determined by the rate of snow accumulation, which is largely a function of the amount of

precipitation falling as snow, and the rate of melting, which is largely a function of air temperature

[Summerfield, 1991].

The rate of snow accumulation in the mid-latitudes is determined by the seasonal distribution

of precipitation and its form (i.e. rain vs. snow), rather than the total annual amount of precipitation.

This is because even heavy precipitation is glacially ineffective if it falls as rainfall during summer

months, when high temperatures and prolonged periods of clear skies enable solar radiation to

achieve high melting rates at the snow or ice surface. A better indicator of glacier distribution is the

combination of total annual precipitation and the ratio of snowfall to total annual precipitation

(nivometric coefficient) [Sugden and John, 1990]. These factors are influenced by latitude, elevation,

aspect and, especially, continentality (i.e. the distance to the nearest ocean). Regions with a

nivometric coefficient approaching unity and a high total annual precipitation provide the best

conditions for nourishing glaciers (e.g. mountains of SW Greenland). Regions with a nivometric

coefficient approaching unity and a low total annual precipitation are less suitable for growth, but

more suitable for prolonged glacial survival (e.g. Antarctic ice sheet). Regions with medium

nivometric coefficients and high total annual precipitation are less suitable for glaciation, but may

nonetheless contain extensive ice sheets and valley glaciers (e.g. high-altitude mid-latitude mountain

massifs) [Sugden and John, 1990].

The rate of melting (ablation) at any given location is determined by the heat sources at that

location. Solar radiation is by far the most important heat input to glacier systems, and hence, is also

extremely important in determining their global distribution. Because melting can only occur at

temperatures greater than 0°C, it is the mean summer temperature rather than the mean annual

temperature that influences the amount of snow melt that will occur. In general, the lower the

incoming solar radiation, the greater the chances of glacier survival (providing that the previously

described precipitation conditions are met). The amount of solar radiation received at any given

location is influenced by continentality, altitude, slope aspect and, especially, latitude [Hattersley-

Smith, 1974; Summerfield, 1991]. The earth’s high latitude regions are thus the most favourable to

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the existence of glaciers. These regions receive low amounts of annual radiation and experience

prolonged winters with unbroken periods of temperatures below 0 °C. This is primarily the reason

why the North Polar ocean basin is covered with pack ice, and the South polar landmass is buried

beneath a continental ice sheet [Sugden and John, 1990].

2.2.3 Geographic distribution of glaciers

Glaciers cover approximately 14.9 million km2 (10%) of the earth’s surface [Briggs et al.,

1993]. Of this, about 12.5 million km2 is accounted for by the Antarctic ice sheet, and 1.7 million

km2 is accounted for by the Greenland ice sheet. The remainder (0.7 million km2) is distributed

mainly among high-latitude ice caps and high-altitude glaciers. Aside from the two ice sheets,

glaciation is concentrated in the northern hemisphere [Figure X], mostly on the islands of the North

Polar basin and on the uplands of the oceanic peripheries (e.g. Alaska and Scandinavia) [Sugden and

John, 1990]. Appreciable amounts of ice cover also occur in the highlands of the middle and low

latitudes, such as the Alps, Karakoram and Himalayan ranges. The areal extent of ice in the African

continent is negligible, and the only large ice caps in the southern hemisphere (outside of the

Antarctic ice sheet) are those of the southern Andes and the Antarctic Peninsula [Sugden and John,

1990]. Thus, present day ice cover is discontinuous and unbalanced between the two hemispheres

and major landmasses [Østrem, 1974].

Ice is an important part of the Canadian landscape. Indeed, only Antarctica and Greenland

contain more. It is estimated that glaciers and ice fields cover as much as 2% (200,000 km2) of

Canada’s total land area. However, at present the total number of these features is unknown. Glaciers

and ice fields are found in two regions of Canada – the Western Cordillera and mountains of the

eastern Arctic – where they are numerous and widely distributed [Figure X]. In the western

Cordillera, heavy snowfall from Pacific storms nourishes ice fields and valley glaciers on the western

slopes of British Columbia’s Coast Mountains. Although fewer glaciers exist on the drier eastern

slopes and inland ranges of the western Cordillera, they are fairly common in the Selkirk Range and

parts of the Rocky Mountains [Trenhaile, 1990]. The Arctic islands contain many glaciers and also

have many large icecaps. Indeed, the bulk of the ice in Canada is found here [Oerlemans, 2001].

Ellesmere, Baffin, Devon and Axel Heiberg islands contain huge icecaps that range can reach a

39

Wave length orderof climatic change

1. High frequency low magnitude oscillations.

2. Short term fluctuations

3. Medium term fluctuations

4. Medium term fluctuations

5. Long term fluctuations (Low frequency high magnitude events).

1-10

10-100

100-1,000

1,000-10,000

10,000-100,000

Small névé fieldsand snow beds

Cirque glaciers

Small valleyglaciers

Ice caps, ice fields,large valley glaciers

Ice sheets

Approx. timescale (years)

Glacier typeaffected

Table X. Relationship of different time scales and the climatic response level of glaciers (Sugden and John, 1990).

thickness of 1km. All of these ice features are remnants of the ice fields of the last ice age, which

peaked about 18,000 BP.

2.2.4 Glaciers and climate change

2.2.4.1 Short-, medium- and long-term glacier responses to climate change

Changes in glacier distribution are strongly

linked to short-, medium- and long-term

variations in climate detectable at time scales

ranging from decades to millennia (Table X.).

Glaciers grow and spread outwardly during

colder periods and they shrink in depth and

volume during warmer periods. However,

glaciers do not respond immediately to

changing climate, nor do they respond to this

change in exactly the same way. Changes in

climate are so rapid compared to the typical

response times of glacial systems that glaciers

probably never achieve true equilibrium with

their environments [Sugden and John, 1990]. Thus, glaciers are usually highly dynamic systems.

However, the exact nature of these dynamics will vary from glacier to glacier, and depend heavily on

glacier morphology.

The shortest-term glacier fluctuations (< 10 years) are caused by hourly, daily, weekly and

seasonal variations in weather patterns. These meteorological oscillations are generally low

magnitude events of high frequency. Some – such as the day/night melting cycle or the accumulation

season/melting season cycle – are predictable and have recognizable return periods, while others –

such as exceptionally sunny spells or periods of continuous rainfall – may be termed random.

However, the random element in this oscillatory behavior becomes more marked as the time scale

increases, making it more difficult to correlate glacier distributions with climatic events over decades

and centuries [Sugden and John, 1990]. Nonetheless, short-term glacier fluctuations (10 – 100 years)

have been linked to minor climatic oscillations. These include periods of climate cooling, increases

40

in total annual precipitation, heavier winter snowfalls and more prolonged springs [Sugden and John,

1990]. Over periods of centuries, glacier fluctuations have also been linked to increasingly vigorous

atmospheric circulation [Andrews et al., 1972]. Medium-term glacier fluctuations (100 – 10,000

years) have generally been linked to periods of cooling and warming, which in turn, are linked to

variations in the emission of solar radiation [Sugden and John, 1990]. Long-term glacier fluctuations

(10,000 – 100,000 years) are likely caused by a combination of factors. The primary forcing

mechanisms for climate change during the Quaternary period (the past 2 – 3 million years) are

thought to be changes in the earth’s orbital parameters [Imbrie et al., 1993]. Taken together, these

factors cause variations in the amount of solar radiation received throughout the year on different

parts of the earth’s surface, thus altering the most fundamental input to the earth’s climate system.

This external forcing mechanism brings about responses and chain reactions in the earth’s internal

elements (notably the atmosphere, the oceans, the hydrological cycle, vegetation cover, glaciers and

ice sheets) [Benn and Evans, 1998; Nesje and Dahl, 2000; Oerlemans, 2001].

Once glaciers and ice sheets are established, they can have considerable large-scale impacts

on regional climate and, in doing so, regulate their own existence [Benn and Evans, 1998;

Oerlemans, 2001]. Glaciers and ice sheets can exist for some time in a state of disequilibrium with

prevailing climate because their sheer sizes not only produce a lag in response time but also provides

an environmental buffer to outside forcing mechanisms [Benn and Evans, 1998]. The presence of

snow and ice is also able to modify its local climate to some extent. Once an ice sheet has grown, it

is able to modify its own temperature and precipitation patterns. For example, glaciers can increase

local precipitation and modify local surface radiation balances by increasing the amount of incoming

solar radiation reflected back into the atmosphere [Benn and Evans, 1998]. These feedbacks make

the relationship between glaciers and climate a complex one.

2.2.4.2 Glacier responses to past and present climates

The Quaternary is usually subdivided into glacial and interglacial periods, with further

subdivisions into stadials (shorter cold periods within interglacial or interstadial periods) and

interstadials (shorter mild episodes within a glacial stage) [Nesje and Dahl, 2000]. Glacial stages

correspond to cooler periods of time associated with major expansions of glaciers and ice sheets.

Interglacial stages correspond to warmer periods of time associated with glacier and ice sheet

41

recession. There have been between 30 and 50 glacial/interglacial cycles in the past 2.5 million years

[Ruddiman and Kutzbach, 1990]. The most recent of these glaciations, the Wisonsin glaciation,

occurred in North America between 75,000 and 10,000 years BP.

The Wisconsin glaciation can be subdivided into early (75,000 to 64,000 BP), middle (64,000

to 23,000 BP) and late (23,000 to 10,000 BP) substages. However, a detailed record only exists for

the most recent of these. During the glacial maximum during the Late Wisconsin, a continental ice

sheet was more or less continuous over the North American continent [Nesje and Dahl, 2000]. The

continental ice sheet consisted of two main parts: (a) the Laurentide ice sheet, and (b) the Cordilleran

glacial complex [Nesje and Dahl, 2000; Oerlemans, 2001; Trenhaile, 1990]. The Laurentide ice

sheet extended from the Arctic Ocean in the Canadian Arctic Archipelago to the mid-western U.S.

states in the south, and from the eastern slopes of the Rocky Mountains in the west. At its maximum

extent, this ice sheet covered most of the land lying north of the Missouri and Ohio rivers, as well as

northern Pennsylvania and all of New York and New England. North of 60°N, the ice sheet spread

over Ellesmere Island and was connected to the northwest Greenland Ice Sheet [Oerlemans, 2001].

The Cordilleran glacial complex was centered in the Coastal range and Rocky Mountains in the west.

At its maximum extent, this glacial complex stretched from the Pacific shores to the Laurentide ice

sheet in the east, and as far south as the Columbia River south of the Canada/US border. The

Laurentide ice sheet reached its maximum extents between 22,000 and 17,000 BP, while the

cordillera ice sheet reached its maximum extent between 15,000 and 14,000 BP [Nesje and Dahl,

2000]. These glacial periods were then followed by warmer periods of rapid ice retreat between

9,000 and 7,000 BP [Benn and Evans, 1998]. Remnants of the Laurentide Ice Sheet still exist today

in the form of the Barnes and Penny Ice Caps on Baffin Island, and continue to undergo retreat in our

present interglacial climate [Benn and Evans, 1998].

The recession of glaciers and ice sheets after the late Wisconsin was followed by a series of

less extensive glacier advances and formations. These “neoglaciations” [Porter and Denton, 1967]

have been reported from the Canadian Cordillera, the US Rocky Mountains, the Brooks Range in

Alaska and the Torngat Mountains in Labrador [Nesje and Dahl, 2000]. In Canada, neoglacial

advances took place between 6000 and 5000 BP, 4000 and 3000 BP, and 2500 and 1800 BP. The

most recent neoglacial advance, the Little Ice Age, started shortly after 900 BP and culminated in the

18th and 19th centuries. During the Little Ice Age, glaciers advanced on all continents. In North

42

America, major advances occurred during the late 17th to early 18th century, early to mid-19th

century, and late 19th to early 20th century. These advances were then followed by significant glacier

retreat, especially after 1910 to 1920, then stability (or even advance) around 1945 [Nesje and Dahl,

2000].

2.2.5 Predicting glacier response to future climates

2.2.5.1 The importance of predicting glacier response to future climates

It is predicted that a doubling of atmospheric CO2 over pre-industrial levels will occur

between 2050 and 2100 AD. This change in atmospheric composition is expected to bring about an

average global increase in air temperature ranging from 1 – 3.5ºC, a warming that is not expected to

be geographically uniform. It is expected that the greatest increases in air temperature will occur at

higher latitudes [Flato et al., 2000; Maxwell, 1997]. Indeed, the coupled general circulation model of

the Canadian Centre for Climate Modelling and Analysis (CCCma) of Environment Canada predicts

an increase in mean annual air temperature ranging from 2 - 6°C over the Canadian north under a

CO2-doubling scenario [Smith and Burgess, 1999; Smith et al., 2005].

It is generally believed that glaciers in northern regions will respond quickly to this warming

trend. Significant ice losses have already been observed on the Greenland Ice Sheet [Krabill et al.,

1995] and several Alaskan glaciers [Arendt et al., 2002]. The assessment and prediction of the

impacts of climate change on glaciers is necessary to determine whether adaptation measures will be

required. This is necessary at both local and global scales. At a global scale, glaciers impact upon the

global climate and sea level. Glaciers and ice caps outside of Greenland and Antarctica have the

potential to raise sea levels by an estimated 0.5m [Church et al., 2001]. At a local scale, glaciers

impact upon their nearby surrounding, which often include human habitats. Glacierized areas supply

meltwater for hydroelectricity, irrigation, industry, domestic use, and the development and

maintenance of stream-associated habitat [Demuth and Pietronomo, 1999]. Meltwater outbursts and

rapid ice advances can result in the loss of pasture lands, human property and even human life

[Grove, 1988]. Predicting future glacier response to climate change involves the continued

monitoring and modeling of glacier systems.

43

2.2.5.2 In situ monitoring of glacier dynamics

2.2.5.2.1 The role of in situ monitoring

In situ measurements of glaciological and meteorological quantities are important for

understanding the major processes in controlling glacier dynamics [Oerlemans, 2001].

Measurements on glaciers have not, as yet, been set up as part of a global program. However, an in

situ monitoring program does exist in Canada, run jointly by interested Government and University

researchers (see following sections). In recent years, in situ observations have become increasingly

valuable for calibrating satellite data and supporting and validating models.

2.2.5.2.2 Automated stations at Western Canadian glacier sites

Munro et al., [2004] describe the automated weather station (AWS) program for Western

Canadian glaciers. The AWS program is a work in progress since the installation of the first AWS at

the Geological Survey of Canada (GSC) base camp site adjacent to the Peyto Glacier, in 1987. After

a few years of trial and error and data interruptions, the station evolved into a reasonably reliable

collector of hourly precipitation, temperature, solar (shortwave) radiation, long-wave radiation, air

temperature, relative humidity, windspeed and direction data. In 2003, AWS measurements began at

the Place Glacier as the AWS program continues to expand. The development of the program has

been possible due to collaboration among the GSC, University of Toronto and the University of

British Columbia, where research funding has been gathered from various sources throughout the

years. Among these sources are the National Sciences and Engineering Research Council (NSERC)

of Canada, but most currently through the Cryospheric System (CRYSIS) to assess global change in

Canada, a program of the Meteorological Service of Canada.

2.2.5.2.3 Automated stations in the Canadian Arctic Archipelago

Labine and Koerner [2004] describe the automated weather stations set up on the Canadian

Arctic Archipelago. Automatic weather stations have been running in the Canadian Arctic

Archipelago since 1987. A small collection of AWS are located on the Meighen, Melville and

Agassiz ice caps, as well as some of their offspring glaciers. Most of the stations contain only an air

44

temperature sensor as well as an ultra sonic snow depth sensor because these are the most effective

measurements given the nature of the nature of the ongoing research. Issues surrounding the

continuous year-round automatic collection of data have generally been resolved at these stations.

The emphasis now is to understand the quality of the data and to be able to identify instrumentation

and technique errors. Detailed methods for measuring mass balances from these data are described

by Koerner [Unpublished].

2.2.5.3 Remote sensing of glacier dynamics

2.2.5.3.1 Passive vs. Active remote sensing

Areas of extensive snow and ice accumulation are typically remote and inaccessible. Thus, the

direct measurement and monitoring of desired parameters is extremely difficult or impossible. Over

the past few decades, researchers have directly addressed this problem by using airborne and satellite

observation systems to monitor glacier dynamics. The remote observation of glaciers and ice sheets

has been accomplished using a variety of methods, including airborne photography [Lachapelle,

1962] and digital imagery derived from airborne and orbital visible/near-infrared sensors [Williams

et al., 1991]. While passive satellite optical sensors such as SPOT and LANDSAT-TM are

commonly used to monitor snow, glacier ice, and other hydrological parameters [Hall et al., 1995],

they have limited applicability if polar regions because of the extent and duration of the polar night,

cloud cover, and the saturation of visible bands by snow and ice [Oerlemans, 2001; Short and Gray,

in review]. These restrictions can be overcome by the use of airborne or orbiting active imaging

systems – such as spaceborne aperture radar (SAR) – can acquire data independently of weather,

season and time of day [Lillesand et al., 2004]. It is the interferometric capabilities of SAR, often

referred to as satellite radar interferometry (InSAR), that are especially important for monitoring

glacial ice dynamics [Goldstein et al., 1993].

Orbital SAR data have revolutionized glacier monitoring by enabling frequent observations of

remote ice masses and by contributing new techniques to monitor glacier ice dynamics [Oerlemans,

2001; Short and Gray, in review]. These data have become increasingly available with the launches

of the European Earth Resources Satellites (ERS-1 and –2), the Japanese Earth Resources Satellite

(JERS-1), and more recently RADARSAT-1 [Demuth and Pietronomo, 1999]. RADARSAT-1 has

45

been shown to be particularly appropriate for monitoring glacier ice dynamics at high latitudes. The

SAR sensor aboard RADARSAT-1 operates at a single frequency (C-band; wavelength = 5.6cm) and

horizontal-like polarization (horizontal transmit-horizontal receive, HH) resulting in one intensity

measure per pixel. Unlike other SAR orbital platforms, the RADARSAT antenna can be configured

to allow for more frequent imaging of an area of interest [Demuth and Pietronomo, 1999]. The

satellite orbits at an altitude of 798 km, and has a 24-day repeat orbit (repeat coverage of 1 day over

the arctic; 3 days at mid-latitudes), an inclination of 98° and a spatial resolution of ~9m [Demuth and

Pietronomo, 1999; Short and Gray, in review]. The launch of a second generation RADARSAT

satellite, RADARSAT-2, is planned for December 2005. Among other differences [see Short and

Gray, 2004], RADARSAT-2 will have easily interchangeable left and right imaging modes and a

finer spatial resolution (~3m in ultra-fine mode) than RADARSAT-1. The use of SAR to monitor

and map glacier dynamics is well documented. For example, Adam et al., [1997] and Partington

[1998] showed that transitions between dry snow, wet snow, ice facies and bedrock could be

identified and mapped using SAR at various frequencies and polarizations. Smith et al., [1997]

showed that SAR was effective in monitoring temporal trends in transient snow line variations, the

presence of melting conditions and the delineation of glacier facies. However, little attention has

been given to the significance of the operational use of this data for glacier mass balance and ice

velocity studies [Demuth and Pietronomo, 1999; Short and Gray, in review]. It is thus important that

future research efforts address these research issues directly.

2.2.5.3.2 Case study: Inferring the glacier mass balance of Peyto Glacier, Canada

Demuth and Pietronomo [1999] investigated the potential of RADARSAT-1 data to assess the

glacier facies configuration and snow-line/accumulation area mapping of Peyto Glacier, Canada for

1996. Peyto Glacier is located in the Canadian Rocky Mountains, and contributes flow to the

Mistaya River Catchment and the North Saskatchewan River Basin. Peyto glacier presently ranges in

elevation from 2140m to 3180m above sea level. The glacier consists of three extensive

accumulation basins that feed into a valley glacier configuration.

The authors collected two RADARSAT-1 SAR images for dates in September and November

1996. Both scenes were imaged at a nominal incidence angle at the scene centre of approximately

42º using standard beam mode S6 [RSI, 1997]. The pixel dimensions for the raw data were

46

approximately 20m in range and 27m in azimuth. The images were georeferenced over the full scene

using standard Canadian National Topographic Series (NTS) 1:50,000 scale UTM map sheets and

orthocorrected using digital elevation information derived from 1986 stereo vertical photos (British

Columbia Terrain Resources Information Management (BCTRIM) data).

The orthorectification procedure followed the methods of Toutin [1995] which has been shown

to be effective in mapping snow-lines in glacier environments with ERS data [Adam et al., 1997].

The radar brightness image was calculated for the original raw scene and the ortho-rectification

procedure was then applied using a gamma filter (3 x 3 pixel window) during pixel resampling to a

25m resolution UTM base map and DEM. The satellite root mean square (RMS) positional error

was less than a pixel for both image dates. Radar brightness training areas (representative sample

sites) were delineated over typical zones of contrasting glacier facies for the September image. The

November image, having identical orbit and imaging parameters was used as a reference image. The

reference image was used to guide the training area delineation so as not to introduce regions of

radar shadow or layover. Areas of ortho-modeling distortion caused by DEM errors were also

avoided.

Numerical data representing the dielectric and scattering attributes for the training areas were

used to generate the first (at the time of writing) radar brightness signatures for glacier facies using

RADARSAT-1. The radar brightness signatures were applied to automatically classify the

September image and delineate the extent of the transient snowline using an image-analysis

classification routine. Based on the intended application, the expected differences in the dielectric

and scattering responses of the known facies configuration and the limitation of having data derived

from only a single-intensity measurement (i.e. a single frequency and polarization), a simple

minimum-distance-to-means decision rule was chosen. For comparison purposes, the snowline was

first delineated manually on the image with the aid of the reference image, photographs and limited

field documentation.

Results of the study showed (a) that the bulk of the transient snow-line could be readily

distinguished and corresponds well to available field documentation and ground-based photography

(although the position of the snow-line in some regions appeared to be obscured by the effects of

steep topography or radar shadow); and (b) that RADARSAT-1 SAR data could distinguish between

late-summer firn and bare ice facies. These results show, despite some unresolved issues, that

47

RADARSAT-1 data shows promise for the estimation of glacier mass balance of temperate alpine

glaciers. The authors suggest that more confident facies delineation over wider areas and regions of

glacier cover may require the coupling of SAR with fine-resolution optical imagery as suggested by

the studies of Rott [1994] and Hall et al., [1995].

2.2.5.3.3. Case study: Large scale albedo trends from ISCCP Satellite data over the Coast Mountain

Ranges of British Columbia, Canada (1983-2000)

Trishchenko et al., [2005] used ISCCP satellite data to investigate the long-term changes in

surface albedo over the Coast Mountain Ranges of British Columbia, Canada (1983-2000). This data

set was created as part of the ISCCP Project from 1983 to 2000. Forty-four parameters were

generated as part of this project, including SW and LW upward and downward fluxes at the ground

surface and TOA. A detailed description of this dataset and its processing is provided by Rossow and

Zhang [1995] (ftp://www.giss.nasa.gov/pub/fc/documents/fluxdoc.ps).

Despite quite substantial inter-annual variability, there exist well-defined negative trends in SW

albedo over Coast Mountain Ranges detected from coarse resolution satellite data which may serve

as potential indicator of changing climate . For 11 selected regions, that cover areas from 45N to

65N, the average observed trends derived from ISCCP data varied from –3.3 to –8.6 % per decade

(Figure X). the average trend is –6.1% per decade. Trends observed from the ISCCP were cross

checked against Polar Pathfinder and Canada Centre for Remote Sensing datasets. Trends from Polar

Pathfinder data agree very well with ISCCP derived trends. Trends from CCRS AVHRR data show

similar results for northern areas and slight difference for southern region due to significantly shorter

period. High spatial resolution images from Landsat TM confirms that major reason for reduced

albedo is due to shrinking area of snow/ice covered mountain area. Contribution of cloud-screening

effect, inter-annual variability, satellite calibration and surface BRDF effects still require further

investigation.

48

Figure X. Large scale albedo trends from ISCCP Satellite data over the Coast Mountain Ranges of British Columbia, Canada (1983-2000)

49

2.2.5.3.4 Case study: Elevation changes of ice caps in the Canadian Arctic Archipelago

Abdalati et al., [2004] conducted precise airborne laser surveys in the spring of 1995 using

NASA’s Airborne Topographic Mapper (ATM-1) on board the NASA P-3 aircraft to assess the

current mass balance of the Canadian ice caps. These measurements were repeated using a

commercial twin otter platform in the spring of 2000. The survey trajectories were designed to cross

the broadest and longest portions of the major ice caps, and where possible, were planned to cover

some of the more significant outlet glaciers.

The ATM combined precise ranging capability with global positioning systems (GPS)

techniques to retrieve surface elevation to a root-mean-square error of 10cm or better [Krabill et al.,

1995]. As such, it provides a valuable tool for measuring changes in the ice cap surface elevations by

means of repeat surveys when there is sufficient time between measurements. The surveys were

conducted over a five year period in an effort to minimize the effects of year-to-year variability. The

aim of the surveys was to provide a spatially broad assessment of ice cap thinning and thickening

rates over the five year time interval. These measurements complement the in situ measurements of

accumulation and mass balance that have been ongoing as part of the Canadian glaciology program

[see Koerner, 2002; Koerner, ; Munro et al., 2004].

The surveys were carried out in the springs of 1995 and 2000. In 1995, a first generation ATM

(ATM-1) was used. A 250 µJ laser pulse was directed towards the ground by a mirror angles at 10°.

The mirror scan rate was 10Hz, and the laser pulse repetition frequency was 1kHz. These parameters

gave a 5m ground spacing between laser shots, with the size of each spot being approximately 0.5m.

The 2000 were conducted with a second generation ATM (ATM-2). The ATM-2 had a laser pulse

energy of 125 µJ, operated at a scan rate of 20 Hz, with a 15° scan angle and a pulse repetition

frequency of 3 kHz. This gave a denser sampling, with laser spots being spaces at approximately 2

m. Change estimates were made by comparing each individual laser pulse surface return from the

2000 campaign with the returns from the 1995 season located within a 1 m horizontal search radius.

More than 90% of the surveys yielded useful (non cloud-covered) data for comparison.

50

The results of this study showed that: (1)

with only a few exceptions, thinning was

more evident on the warmer southern ice

caps of Baffin island, Penny, and Barnes than

on the more northerly ones. In general,

thinning was evident at the lower elevations

near the ice cap edges but thickening was

observed in more central regions [Figure X];

(2) while most of the individual ice caps in a

specific region exhibited a clear inverse

relationship between elevation and thinning

or thickening rate, a strong quantitative

relationship could not be easily extracted; (3)

the different relationships existed between

annual elevation changes and elevation were

evident in many of the regions surveyed. One

reason for this was the varied topography

found in some of the more mountainous regions, which strongly influences atmospheric circulation

characteristics and creates highly localized effects. As a result, annual elevation changes can vary

widely for a given elevation, even on an individual ice cap; (4) the absence of a consistent

relationship between annual elevation change and elevation makes it difficult to extrapolate these

results to a highly accurate mass balance estimate for the ice caps as a whole. They do, however,

offer strong indications of the overall state of the ice cap balance; (5) these results were consistent

with field data from Devon and Agassiz ice caps.

Abdalati et al., [2004] provide the following climatological interpretation of their results. The

thinning of ice caps at lower elevations by 0.5m a-1 was generally explained by recent warm

temperature between 1995 and 2000. These temperature anomalies, however, did not explain the

thinning trend on Baffin Island where thinning was most pronounced. Indeed, temperatures in this

region during this time period were not particularly warm, and in some cases they were slightly cool.

Thus, the authors contend that the observed changes here of 1 m a-1 were not directly attributable to

climate conditions of the late 1990s but are most likely part of the ongoing response to a longer-term

Figure X. Elevation changes (dh/dt) as a function of surface elevation for different ice sheets. Thinning is evident in most areas and is most pronounced at the

lower elevations. At high elevations, there is consistently negligible thinning or even thickening. The more southerly ice caps show greater thinning

rates than the northerly ones.

51

deglaciation in recent centuries [Jacobs et al., 1993; Jacobs et al., 1997]. The thickening of ice cap

interiors between 1995 and 2000 was explained by changes in climatic conditions. Of particular

importance were the combined effects of increases in annual precipitation and the fraction of

precipitation that falls as snow (i.e. nivometric coefficient) during the latter half of the twentieth

century [Zhang et al., 2000]. Together, these trends provided conditions suitable for accumulation on

the ice caps. The elevation dependence of estimated elevation change was consistent with in situ

measurements on the Meighen, Agassiz, Devon, ice caps in Nunavut, and the Melville ice cap in the

Northwest Territories.

In situ measurements on several ice caps on the Queen Elizabeth Islands showed that mass

balance at the field sites between 1995 and 2000 was more negative than observed over the

preceding 2 or 3 decades and that the changes were strongly driven by melt. The relationship

provides further evidence that the survey period was anomalously warm and that these warm

temperatures were responsible for some of the thinning at the lower elevations. However, these

longer in situ data records also suggest that the thinning of the late 1990s appears to be a

continuation of a phenomenon that began in the mid-1980s.

2.2.5.3.5 Case study: Glacier dynamics in the Canadian high Arctic

Short and Gray [in review] used speckle tracking of RADARSAT-1 SAR data to monitor the

velocities of 11 glaciers in the Queen Elizabeth Islands of the Canadian high Arctic between 2000

and 2004. Speckle tracking of SAR data has emerged as a powerful velocity measuring technique. In

this technique, the relative displacement of small image chips in a repeat-pass pair of SAR images is

measured in order to estimate surface movement. If strong image features, such as from crevasses,

are present then the technique will track the displacement of these features instead. Hence, the

technique is also referred to as intensity tracking [Strozzi et al., 2002]. The successes of speckle

tracking in polar regions, mostly on fast flowing or surging glaciers, are described in detail elsewhere

[Gray et al., 2001; Joughin, 2002; Joughin et al., 1999; Murray et al., 2003b; Murray et al., 2002;

Strozzi et al., 2002].

The authors collected RADARSAT-1 data over the Queen Elizabeth Islands between 2000 and

2004. Most data were obtained at a spatial resolution of ~9 m resolution, which was optimal for the

52

high relief and relatively narrow glaciers of the Canadian Arctic. Only large glaciers were selected

for study since the size of the image chips used in the speckle tracking technique dictated that sites

should be > 1.5 km in width.

Speckle tracking began with the acquisition of two coarsely registered repeat-pass single look

complex SAR images. The speckle tracking routine then used a two-dimensional cross-correlation

function with small detected image chips to determine relative displacement in the slant range

(across-satellite track) and azimuth (along-satellite track) directions [Gray et al., 2001; Joughin,

2002; Strozzi et al., 2002]. A digital elevation model (DEM) – 1:250,000 Nunavut DEM – was used

to remove the topographic component in the slant range displacement. The DEM was also used to

convert the measurements from slant range to ground range. Both the ground range and azimuth

displacements were then calibrated using a stationary reference area, such as a rock outcrop or valley

bottom. Finally, movement parallel to the ice surface was assumed [Joughin et al., 1998] and down-

slope surface velocity estimates were calculated from the corrected displacements, surface slopes and

the time separation of the data acquisitions (24 days for these RADARSAT-1 data).

Although Short and Gray [in review] could not give an

overall error estimate for the results of speckle tracking,

they classified glaciers on the basis of whether they (a) were

surging, (b) were non-surging, but nonetheless had elevated

flow velocities, or (c) had slow rates of flow.

Dramatic surges were observed in the Middle, Mittie

and Otto glaciers during the observation period. Middle

Glacier is a land terminating outlet glacier on the west side

of the Müller Ice Cap. Observed surface velocities showed

that in 2000 speeds over the lower reaches of the glacier

were up to ~180 m a-1, and although there had been a slight

slow down by 2004, speeds were still ~150 m a-1 at the same

location. Speckle tracking measurements here indicated

velocities of 41 m a-1 in 2000, reducing to 25 m a-1 by 2004.

These velocities suggest that the Middle Glacier

accumulation area may have been losing mass. Mittie

Figure X. Surface velocities of Mittie Glacier in 2003 with the profile marked

at 5 km intervals.

53

Glacier drains north from the Manson Ice Cap on south east Ellesmere Island. Copland et al., [2003]

identified the glacier as surging in 2000. While the observed velocities in 2003 were comparable to

those of 2000, 2004 saw a slight deceleration in velocities. The profiles suggest that the surge

activity is concentrated in the lower 32 km of the glacier [Figure X]. Otto Glacier is a tidewater

glacier on the north west coast of Ellesmere Island. The surge activity for Otto was concentrated in

the lower 15 km of the glacier. Ice velocities suggested that the Otto Glacier accumulation area may

have been losing mass in 2002. Unfortunately, the equilibrium line was not covered in the data sets

from other years. Calving rates for Otto Glacier in 2002, 2003 and 2004 were shown to vary

significantly from year to year, with 2003 to 2004, a year of retreat, producing 0.61 ± 0.15 km3 of

icebergs, approximately five times the volume of icebergs of 2001-2002, when the glacier was

advancing.

While not actively surging, the Tuborg, Trinity, Antoinette, Ekblaw and Wykeham glaciers

demonstrated elevated and variable rates of ice flow. Tuborg glacier is on the north west side of the

Agassiz Ice Cap. Observed flow velocities for this glacier were consistently high (150 – 300 m a-1).

Trinity is a major glacier also draining the east side of the Prince of Wales Ice Cap. Trinity appears

to be a fast flowing glacier with speeds of between 600 and

730 m a-1 measured over its tongue, although these are

potentially under-estimated by < 18 m a-1 due to the tidal

effect [Figure X]. Flow velocities here have been

approximately constant for 2003 and 2004, with minor

acceleration over the lower 10 km. Antoinette Glacier lies

just to the south of Tuborg. Antoinette shows a large slow

down between 2003 and 2004. Ekblaw Glacier is one of

several glaciers draining the east side of the Prince of Wales

Ice Cap. It produced speeds of ~ 100 - 300 m a-1 during the

study period. Velocity measurements showed a significant

acceleration over the main trunk between 2000 and 2003

(potentially underestimated by 30 m a-1 due to the tidal

effect). The position of the terminus was constant in the

2000 and 2003 imagery, and the calving rate was estimated

to be ~ 0.29 ± 0.07 km3 a-1. Wykeham Glacier is adjacent to

Figure X. Surface velocities of Trinity and Wykeham glaciers in 2004. Profiles

are marked at 5 km intervals.

54

Trinity Glacier in Talbot Inlet. Observations suggest that the lower reaches of Wykeham Glacier

were flowing at between 240 and 375 m a-1 in 2003 and 2004 (although these figures were

potentially under-estimated by < 18 m a-1 in both years due to the tidal effect) and there was some

acceleration of flow in 2004 [Figure X].

Slow flow rates were observed for Iceberg and

d’Iberville glaciers. Iceberg Glacier drains the west side

of the Müller Ice Cap and is a tidewater glacier with

clear evidence of past surging [Ommanney, 1969]. In

2000 the lower 15 km of the glacier flowed slowly (75 -

130 m a-1) but by 2004 velocities had dropped to < 25 m

a-1. D’Iberville Glacier drains the west side of the

Agassiz Ice Cap. The velocity of 2003 and 2004 are

consistent, where 35 – 80 m a-1 are normal speeds for the

entire 33 km of the profile. Only Canon glacier, located

on the south west side of the Agassiz Ice Cap hints at

being stable. Velocity measurements showed a fairly

steady flow rates ranging from 75 – 200 m a-1 for 2003

[Figure X]. Tentative calculations of a balance velocity indicate that this glacier is in balance at the

equilibrium line.

The results described above suggest that: (1) speckle tracking is an effective technique for

monitoring the dynamics of large high Arctic glaciers; (2) many Canadian Arctic glaciers flow

considerably faster than the 10-50 m a-1 previously assumed [Koerner, 2002]; (3) at any one time,

several QEI glaciers are behaving in a surge-like manner, with active phases of at least 5-9 years, and

quiescent phases potentially as short as 35 years. (4) surge front propagation may be absent from

tidewater surge glaciers [Murray et al., 2003a; Murray et al., 2003b]; (5) significant flow

fluctuations occur even in glaciers not considered to be surge type that experience multi-year

periodic, pulse-like increases in speed, but where the pulses are not large enough to produce the

major ice displacements of full surges [Mayo, 1978]; (7)

While Short and Gray’s [in review] data set does not allow the identification of the pulse

mechanisms, it does identify areas to study and continued monitoring would enable the discernment

Figure X. Surface velocities of Canon glacier in 2003. Profiles are marked at 5 km

intervals.

55

of temporal cycles. Raymond [1987] suggests that surge behaviour may be an extreme end-member

in a continuum of possible pulsating flow. The Ekblaw, Trinity and Antoinette Glaciers might be

interesting cases of pulsating flow to study and to test this suggestion. The possible role of tidewater

cycles [Paterson, 1994] influencing velocity fluctuations should also be investigated. The causes of

all these fluctuations will be important in understanding how ice dynamics might change with

climate.

The limitations of this study are as follows. Attempts to calculate balance velocities are limited

by a lack of glacier thickness data and sometimes come close to the margins of error. However, other

glaciers do better with a reliable estimate of ice thickness and results that do clear the margins of

error. It seems that at current SAR resolutions, speckle tracking velocities cannot be considered

accurate enough to compare with balance velocities in the accumulation regions of small ice caps. In

areas of slightly faster flow the method is more reliable but more ice thickness measurements and

estimates of local ablation rates are needed to develop this application. Finer resolution SAR data

from the planned RADARSAT-2 mission may improve the accuracy of speckle tracking sufficiently

to monitor the slow areas of Arctic ice caps.

Short and Gray [2004] assessed the potential of speckle tracking of RADARSAT-2 data for

monitoring glacier dynamics. Their study suggests that RADARSAT-2 will improve upon the

accomplishments of RADARSAT-1 by having easily interchangeable left and right imaging modes

and a finer spatial resolution. The left-imaging mode will open up more areas of Antarctica to study

and will increase opportunities to obtain multiple pairs to combine differential phase for slow

glaciers. Multiple pairs will also mitigate the effects of shadow in areas of high relief. The ultra-fine

spatial resolution of RADARSAT-2 will make the speckle tracking technique described in previous

paragraphs suitable for smaller glaciers and should improve the accuracy to tens of centimeters, thus

making it possible to monitor much slower glaciers.

2.2.5.3.6 Case study: Investigating subglacial water transport in the West Antarctica Ice Sheet

Gray et al., [200X] used RADARSAT-1 interferometric data to investigate the characteristics of

the basal water system of the West Antarctica Ice Sheet (WAIS). This study is important because

56

improved knowledge of the basal water system is necessary for understanding changes in the WAIS

mass balance and predicting its future contributions to sea level rise [Bougamont et al., 2004].

Gray et al., [200X] collected RADARSAT-1 data over Antarctica during a 30-day period in

1997. Satellite radar interferometry requires repeat coverage with the same geometry, but this was

only possible with RADARSAT-1 every 24 days. As a result, most areas in the WAIS were covered

by only one image pair, some had no repeat coverage at all, but a few had coverage by two pairs.

Where only one interferometric SAR pair existed, ice velocity was estimated by assuming that the

flow vector was parallel to the ice surface then combining displacements in the line-of-sight direction

with less accurate estimated of along-track displacement made using the speckle tracking technique

described earlier [Gray et al., 2001; Joughin, 2002]. However, for the few areas in the WAIS for

which there were both ascending and descending pass pairs it was possible to solve for the 3-

dimensional displacements without using the surface-parallel-flow assumption.

The interferometric data collected during this study implied a vertical surface movement of up to

~2 cm per day. These results represent a first indication that the anomalous changes in surface

elevation on WAIS ice streams may be linked to subglacial water transport. The results also suggest

that imaging all three components of surface ice displacement (north, east, vertical) in the WAIS will

help establish links between basal water transport in the subglacial environment, and ice dynamics

and mass transport to the ocean. Gray et al., [200X] suggest that the RADARSAT-2 satellite should

help to provide such a capability.

2.2.5.4 A modeling approach

2.2.5.4.1 The importance of modeling

The modeling of glacier dynamics is important because it allows researchers to quantify

processes associated with glacier systems, and to put these processes into perspective [Oerlemans,

2001]. Glacier models are based on physical laws and, as simplified abstractions of reality, are

limited in their degree of detail. Remote sensing observations are becoming an increasingly

important source of input to models. The accuracy of model-generated output is usually compared to

in situ observations so that the overall performance of the model can be assessed.

57

2.2.5.4.2 Case study: Modeling snowline migration and discharge patterns of Place glacier, Canada

Marosz-Wantuch [2004] used a spatially distributed surface melt model to analyze glacier

snowline migration and discharge patterns (Place Glacier, Canada). Place Glacier is located in the

Coast Mountains of the Canadian Cordillera. The region is influenced by a maritime climate, which

causes heavy snow accumulation and moderates the annual air temperatures. The Place Glacier

covers a 3.23 km2 area within a 12.04 km2 basin. Its elevation range is 1950-2530 m above sea level.

The model was run using an hourly timestep for a 97 day period during the 2002 melt season. This

period was chosen to coincide with the time period for which satellite data was available. The hourly

interval of the model allowed the detection of runoff variations on a fine temporal scale. The

objectives of this study were to (a) develop a snowpack distribution based on the winter mass

balance; (b) estimate variations in the surface cover during the melt season through the classification

of satellite images, (c) improve a glacier surface melt model for modeling snowline migration and

surface melt using automatic weather station, topographic and satellite data, and (d) validate model

results by comparing modeled snowline extent to snowlines classified from satellite data, and

matching modeled meltwater production to catchment hydrographs.

This modeling approach was based on a one-dimensional point glacier surface model that

assumed a bulk, vertical energy exchange between the atmosphere and the glacier surface. The

model was initialized with a snowpack estimated by winter mass balance measurements and driven

by surface radiation, air temperature, relative humidity, windspeed and precipitation data. These

meteorological data were collected hourly at an automated weather station located approximately

300m north of the glacier snout. Hourly discharge data were derived from a gauging station located

approximately 4km away from the snout of the glacier. The point model was spatially distributed

over the surface according to the variations in radiative energy transfer influenced by topographic

characteristics of the terrain. The spatially variable parameters used in the model included elevation,

slope angle, slope azimuth, the albedo assigned to six surface cover classes (water, rock, vegetation,

dirty ice, clean ice, dark snow, white snow) and altitudinally-dependent meteorological variables.

The topographically-related parameters were derived from two Digital Elevation Models (DEM) at

30m and 15m spatial resolutions. Landsat Thematic Mapper (TM) satellite images were used to

58

classify the surface cover types across the study region using a supervised classification approach to

which surface albedo values were assigned.

Modeling the snowline migration patters produced results comparable to the surface

classification derived from the satellite images. The snow retreated up the glacier as the summer

progressed. However, the details in spatial pattern left room for further improvement, mainly related

to the initial snowpack distribution. The modeled discharge patterns show a response to precipitation

events and, related to, variations in glacier albedo. The model also reflected changes in surface

albedo caused by snowline migration. These values did not show an exact correspondence, which

was partially caused by the lack of consideration of the off-glacier water storage and routing. The

main factors responsible for the discrepancies were the snowpack initialization, meltwater storage

and the turbulent transfer coefficient and fresh snow albedo.

The results of this study support the idea discharge patterns are strongly affected by the

snowline retreat patterns which depend on long term climatic and short term meteorological

conditions, as well as the topography of the basin and the surrounding terrain. Climatic trends affect

the size of the glacier and the spatial extent of the ablation and accumulation areas, which determine

the absorptive potential of glacier ice, one of the main factors affecting glacier mass balance.

2.2.6 Consequences of glacier response to a changing climate

2.2.6.1 Potential impacts on Canada’s freshwater resources

The previously described responses to climatic warming – increases in the fraction of

precipitation received as rainfall, reduced snow and ice cover duration, disappearance of mountain

glaciers – will have important consequences for regional hydrological cycles [Rouse et al., 1997].

However, quantifying this response and its impact on the hydrological cycle is difficult because the

cryosphere is characterized by complex linkages and feedbacks, as well as variable time lags and

storages [Brown et al., 2004]. Of particular interest are the effects of warming-induced glacier

responses on Canada’s freshwater resources.

Canada’s freshwater resources will likely be impacted considerably by warming-induced

glacier melt. Climate warming is expected to have a significant impact on the amount and timing of

59

water discharged in glacierised basins. Increases in air temperature, radiation flux and rainfall bring

about increased surface melting, leading to greater rates of runoff than would occur on a snow-free

catchment [Benn and Evans, 1998]. Consequently, meltwater discharge is generally high during

warmer periods when storage decreases and glaciers retreat, and highest when deglaciation is at its

most rapid. Discharge then declines as glaciers retreat and the supply of stored water is exhausted

[Benn and Evans, 1998]. The reduction in glacier size reduces the potential for meltwater yield. As a

result, continued climate warming may have far-reaching consequences for communities depending

on such sources of water during the summer months. The watersheds of Western Canada that provide

freshwater for domestic, agricultural and industrial use in the western Canadian prairies are one such

region [Demuth and Pietronomo, 2003]. A number of researchers [e.g. Demuth and Pietronomo,

2003; Dornes et al., in preparation; Pietronomo et al., in preparation] have assessed the impacts of

warming-induced glacier melt on the eastern flowing basins of the Rocky Mountain, particularly the

North and South Saskatchewan River Basins. These studies are discussed in greater detail in section

2.4 (Canada’s water-dominated landscapes).

2.2.6.2 Potential impacts on surface-atmosphere energy exchange

Land surface albedo – the fraction of incident (shortwave) solar radiation reflected in all

directions by the land surface [Pinty and Verstraete, 1992] – is one of the most important parameters

controlling the earth’s climate. Albedo is important because it directly determines the amount of

solar energy absorbed by the ground, and hence, the amount of energy available for heating the

ground and lower atmosphere and evaporating water [Rowe, 1991]. It also affects the global climate

system indirectly by controlling the ecosystem energy, water and carbon processes that regulate

greenhouse gas exchange [Wang et al., 2002]. The high albedo of snow and ice means that the

disappearance of glaciers and ice sheets will lead to significant changes in the optical properties of

the earth's surface. It is predicted that these changes will have a considerable impact on local surface

energy balances, and ultimately, lead to further atmospheric heating [Cess et al., 1990; Cess et al.,

1991]. However, significant uncertainties exist over the size of the effect, particularly with sea-ice

[Ingram et al., 1989]. Understanding and predicting the response of glaciers to warming climate is

thus an important ongoing component of climate change research.

60

2.2.7 References

Abdalati, W., W. Krabill, E. Frederick, S. Manizade, C. Martin, J. Sonntag, R. Swift, R. Thomas, J. Yungel, and R.M. Koerner, Elevation changes of ice caps in the Canadian Arctic Archipelago, Journal of Geophysical Research, 109 (F04007), 2004.

Adam, S., T. Toutin, A. Pietronomo, and M.M. Brugman, Using ortho-rectified SAR imagery acquired over rugged terrain for thematic applications in glacier hydrology, Canadian Journal of Remote Sensing, 23 (1), 76-80, 1997.

Andrews, J.T., R.G. Barry, R.S. Bradley, G.H. Miller, and L.D. Williams, Past and present glaciological responses to climate in eastern Baffin Island, Quaternary Research, 2, 303-314, 1972.

Arendt, A.A., K.A. Echelmeyer, W.D. Harrison, C.S. Lingle, and V.B. Valentine, Rapid wastage of Alasks glaciers and their contribution to rising sea level, Science, 297 (382-386), 2002.

Benn, D.I., and D.J.A. Evans, Glaciers and Glaciation, 734 pp., John Wiley and Sons, London, UK, 1998.

Bougamont, M.S., S. Tulaczyk, and I. Joughin, Numerical investigations of the slow-down of Whillans Ice Stream, West Antarctica: is it shutting down likw Ice Stream C?, Annals of Glaciology, 37, 239-246, 2004.

Briggs, D., P. Smithson, T. Ball, P. Johnson, P. Kershaw, and A. Lewkowicz, Fundamentals of Physical Geography, 692 pp., Copp Clark Pitman Ltd, Toronto, Ontario, Canada, 1993.

Brown, R.D., M.N. Demuth, B.E. Goodison, P. Marsh, T.D. Prowse, S.L. Smith, and M.-K. Woo, Climate variability and change - Cryosphere, in Threats to water availability in Canada, NWRI Scientific Assessment Report Series No. 3 and ACSD Science Assessment Series No. 1, pp. 107-116, National Water Research Institute, Meteorological Service of Canada, Environment Canada, Burlington, ON, Canada, 2004.

Cess, R.D., G.L. Potter, J.P. Blanchet, G.J. Boer, A.D. Del Genio, M. Déqué, V. Dymnikov, V. Galin, W.L. Gates, S.J. Ghan, J.T. Kiehl, A.A. Lacis, H. Le Treut, Z.-X. Li, X.-Z. Liang, B.J. McAvaney, V.P. Meleshko, J.F.B. Mitchell, J.-J. Morcrette, D.A. Randall, L. Rikus., E. Roeckner, J.F. Royer, U. Schlese, D.A. Sheinin, A. Slingo, A.P. Sokolov, K.E. Taylor, W.M. Washington, R.T. Wetherald, I. Yagai, and M.-H. Zhang, Intercomparison and interpretation of climate feedback processes in nineteen atmospheric general circulation models, Journal of Geophysical Research, 95, 16601-16615, 1990.

Cess, R.D., G.L. Potter, M.-H. Zhang, J.P. Blanchet, S. Chalita, R. Colman, D.A. Dazlich, A.D. Del Genio, V. Dymnikov, V. Galin, D. Jerret, E. Keup, A.A. Lacis, H. Le Treut, X.-Z. Liang, J.-F. Mahouf, B.J. McAvaney, V.P. Meleshko, J.F.B. Mitchell, J.-J. Morcrette, P.M. Norris, D.A. Randall, L. Rikus., E. Roeckner, J.F. Royer, U. Schlese, D.A. Sheinin, A. Slingo, A.P. Sokolov, K.E. Taylor, W.M. Washington, R.T. Wetherald, and I. Yagai, Interpretation of snow-climate feedback as produced by seventeen general circulation models, Science, 253, 888-892, 1991.

Church, J.A., J.M. Gregory, P. Huybrechts, M. Kuhn, K. Lambeck, M.T. Nhuan, D. Qin, and P.L. Woodworth, Changes in sea level, in The Scientific Basis. Contribution of Working Group 1

61

to the Third Assessment Report of the Intergovernmental Panel on Climate Change., edited by J.T. Houghton, Y. Ding, D.J. Griggs, M. Noguer, P.J. Van der Linden, X. Dai, K. Maskell, and C.I. Johnson, pp. 639-694, Cambridge University Press, Cambridge, UK, 2001.

Copland, L., M. Sharp, and J. Dowdeswell, The distribution and flow characteristics of surge-type glaciers in the Canadian High Arctic, Annals of Glaciology, 36, 82-90, 2003.

Demuth, M.N., and A. Pietronomo, Inferring glacier mass balance using RADARSAT: Results from Peyto Glacier, Canada, Geografiska Annaler, 81 (A4), 521-540, 1999.

Demuth, M.N., and A. Pietronomo, The impact of climate change on the glaciers of the Canadian Rocky Mountain eastern slopes and implications for water resource-related adaptation in the Canadian prairies. Phase I - Headwaters of the North Saskatchewan River Basin, pp. 96, Climate Change Action Fund - Prairie Adaptation Research Collaborative, Ottawa, Canada, 2003.

Dornes, P.F., A. Pietronomo, and M.N. Demuth, Esimation of the glacier-melt contributions to the North and South Saskatchewan rivers, in to be submitted at the Symposium on the contribution from glaciers and snow cover to runoff from mountains in different climates (ICSI), in preparation.

Flato, G.M., G.J. Boer, W.G. Lee, N.A. McFarlane, D. Ramsden, M.C. Reader, and A.J. Weaver, The Canadian Centre for Modelling and Analysis Global Coupled Model and its climate, Climate Dynamics, 16, 451–467, 2000.

Goldstein, R.M., H. Englehardt, B. Lamb, and R.M. Frolich, Satellite radar interferometry for monitoring ice sheet motion: application to an Antarctic ice stream, Science, 262, 1525-1530, 1993.

Gray, A.L., I. Joughin, S. Tulaczyk, V.B. Spikes, R. Bindschadler, and K.C. Jezek, Evidence for subglacial water transport in the West Antarctica Ice Sheet through three-dimensional satellite radar interferometry, 200X.

Gray, A.L., N. Short, K.E. Mattar, and K.C. Jezek, Velocities and Flux of the Filchner Ice Shelf and its Tributaries Determined from Speckle Tracking Interferometry, Canadian Journal of Remote Sensing,, 27 (3), 193-206., 2001.

Grove, J.M., The little ice age, Methuen, London, UK., 1988.

Hall, D.K., R.S.J. Williams, and O. Sigurdsson, Glaciological observations of Brúarjökull, Iceland, using synthetic aperture radar and thematic mapper satellite data, Annals of Glaciology, 21, 217-276, 1995.

Hattersley-Smith, G., Present arctic ice cover, in Arctic and Alpine Environments, edited by J.D. Ives, and R.G. Barry, pp. 195-223, Methuen and Co. Ltd, London, UK, 1974.

Imbrie, J., A. Berger, and N.J. Shackleton, Role of orbital forcing: a two-million year persepctive, in Global changes in the perspective of the past, edited by J.A. Eddy, and H. Oeschger, pp. 263-277, John Wiley, Chichester, UK., 1993.

Ingram, W.J., C.A. Wilson, and J.F.B. Mitchell, Modelling climate change: an assessment of sea ice and surface albedo feedbacks, Journal of Geophysical Research, 94, 8609-8622, 1989.

62

Jacobs, J.D., R. Heron, and J.E. Luther, Recent changes at the northwest margin of the Barnes ice cap, Baffin Island, N. W. T., Canada, Arctic and Alpine Research, 25, 341-352, 1993.

Jacobs, J.D., E.L. Simms, and A. Sims, Recession of the southern part of Barnes ice cap, Baffin Island, Canada, between 1961 and 1993, determined from digital mapping of Landsat TM, Journal of Glaciology, 43, 98-102, 1997.

Joughin, I., Ice-Sheet Velocity Mapping: a Combined Interferometric and Speckle Tracking Approach, Annals of Glaciology, 34, 195-201, 2002.

Joughin, I., L. Gray, R. Bindschadler, S. Price, D. Morse, C. Hulbe, K. Mattar, and C. Werner, Tributaries of West Antarctic Ice Streams Revealed by RADARSAT Interferometry, Science, 286, 283-286, 1999.

Joughin, I., R. Kwok, and M. Fahnestock, Interferometric Estimation of Three-dimensional Ice Flow Using Ascending and Descending Passes, IEEE Transactions on Geoscience and Remote Sensing, 36 (1), 25-37, 1998.

Koerner, R.M., Glaciers of the High Arctic Islands, in Satellite Image Atlas of Glaciers of the World, North America, edited by R. Williams, and J. Ferrigno, pp. 111-146, United States Geological Survey, 2002.

Koerner, R.M., Mass balance of glaciers in the Queen Elizabeth islands, Nunavut, Canada, Unpublished.

Krabill, W.B., R.H. Thomas, C.F. Martin, R.N. Swift, and E.B. Frederick, Accuracy of airborne laser altimetry over the Greenland Ice Sheet, International Journal of Remote Sensing, 16, 1211-1222, 1995.

Labine, C., and R.M. Koerner, Autostations networks in the Canadian Arctic Archipelago: Twenty years later, issues and problems, in Automatic weather stations on glaciers, edited by J. Oerlemans, and C.H. Tijm-Reijmer, pp. 66-71, Institute for Marine and Atmospheric Research Utrecht, Utrecht University, The Netherlands, Pontresina, Switzerland, 2004.

Lachapelle, E.R., Assessing glacier mass budgets by reconassance aerial photography, Journal of Glaciology, 4, 290-297, 1962.

Lillesand, T.M., R.W. Kiefer, and J.W. Chipman, Remote sensing and image interpretation, pp. 763, Wiley, New York, New York, USA, 2004.

Marosz-Wantuch, M., Modelling snowline migration and runoff response for Place Glacier basin, M.Sc thesis, University of Toronto, Toronto, 2004.

Maxwell, B., Responding to global climate change in Canada's arctic:, pp. 82, Environemntal Adaptation Research Group, Atmospheric Environment Service, Environment Canada, 1997.

Mayo, L.R., Identification of unstable glaciers intermediate between normal and surging glaciers, Mat. Glyatsiologicheskikh Issled. Khronica Obsuzhdeniya, 33, 133-135, 1978.

Munro, D.S., M.N. Demuth, and D. Moore, Operating the AWS at western Canada glacier sites, in Automatic weather stations on glaciers, edited by J. Oerlemans, and C.H. Tijm-Reijmer, pp. 72-75, Institute for Marine and Atmospheric Research Utrecht, Utrecht University, The Netherlands, Pontresina, Switzerland, 2004.

63

Murray, T., A. Luckman, T. Strozzi, and A.-M. Nuttal, The initiation of glacier surging at Fridtjovbreen, Svalbard, Annals of Glaciology, 36, 110-116., 2003a.

Murray, T., T. Strozzi, A. Luckman, H. Jiskoot, and P. Christakos, Is there a single surge mechanism? Contrasts in dynamics between glacier surges in Svalbard and other regions, Journal of Geophysical Research, 108 (B5), 2237, 2003b.

Murray, T., T. Strozzi, A. Luckman, H. Pritchard, and H. Jiskoot, Ice dynamics during a surge of Sortebræ, East Greenland, Annals of Glaciology, 34, 323-329, 2002.

Nesje, A., and S.O. Dahl, Glaciers and environmental change, 203 pp., Oxford University Press, New York, NY, USA, 2000.

Oerlemans, J., Glaciers and climate change, 148 pp., A.A. Balkema, Exton, PA, USA, 2001.

Ommanney, S., A Study in Glacier Inventory, The Ice Masses of Axel Heiberg Island, Canadian Arctic Archipelago, McGill University, Montreal, Canada, 1969.

Østrem, G., Present alpine ice cover, in Arctic and alpine environmnents, edited by A.R. Ives, and R.G. Barry, Methuen, London, UK, 1974.

Partington, K.C., Discrimination of glacier facies using multi-temporal SAR data, Journal of Glaciology, 44 (146), 42-53, 1998.

Paterson, W.S.B., The Physics of Glaciers, Pergamon Press, Oxford, UK, 1994.

Pietronomo, A., M.N. Demuth, P.F. Dornes, N. Kouwen, A. Bingeman, C. Hopkinson, and R.B. Brua, Streamflow shifts resulting from past and future glacier fluctuations in the eastern flowing basins of the Rocky Mountains - Interim Report for Year 2, pp. 71, to be submitted to Alberta Environment Climate Change Resources Users Group, Ottawa, Canada, in preparation.

Pinty, B., and M. Verstraete, On the design and validation of surface bidirectional reflectance and albedo model, Remote Sensing of Environment, 41, 155-167, 1992.

Porter, S.C., and G.H. Denton, Chronology of neoglaciation in the North American Cordillera, American Journal of Science, 265, 177-210, 1967.

Raymond, C.F., How do glaciers surge? A review, Journal of Geophysical Research, 92 (B9), 9121-9134, 1987.

Rossow, W.B., and Y.-C. Zhang, Documentation of Radiative Flux Dataset (FC), Online document, ftp://www.giss.nasa.gov/pub/fc/documents/fluxdoc.ps, 1995.

Rott, H., Thematic studies in alpine areas by means of polarimetric SAR and optical imagery, Advances in Space Research, 14 (3), 217-226, 1994.

Rouse, W., M. Douglas, R. Hecky, A. Hershey, G. Kling, L. Lesack, P. Marsh, P. McDonald, M. Nicholson, N. Roulet, and J. Smol, Effects of climate change on the fresh waters of Arctic and subarctic North America, Hydrological Processes, 11, 873-902, 1997.

Rowe, C.M., Modeling land-surface albedos from vegetation canopy architecture, Physical Geography, 12 (2), 93-114, 1991.

64

RSI, RADARSAT-data product specifications, RADARSAT International, Richmond, BC, Canada, 1997.

Ruddiman, W.F., and J.E. Kutzbach, Late Cenozoic plateau uplift and climate change, Transactions of the Royal Society of Edinburgh: Earth Sciences, 81, 301-314, 1990.

Short, N., and A.L. Gray, Glacier dynamics in the Canadian High Arctic from RADARSAT-1 data, Canadian Journal of Remote Sensing, in review.

Short, N.H., and A.L. Gray, Potential for RADARSAT-2 interferometry: glacier monitoring using speckle tracking, Canadian Journal of Remote Sensing, 30 (3), 504-509, 2004.

Siegert, M.J., Ice sheets and late Quaternary environmental change, 231 pp., John Wiley and Sons, Chichester, England, UK, 2001.

Smith, L.C., R.R. Forster, B.I. Isacks, and D.O. Hall, Seasonal climatic forcing of alpine glaciers revealed with orbital synthetic aperture radar, Journal of Glaciology, 43 (145), 480-488, 1997.

Smith, S.L., and M.M. Burgess, Mapping the sensitivity of Canadian permafrost to climate warming, in Interactions Between the Cryosphere, Climate and Greenhouse Gases (Proceedings of IUGG 99 Symposium HS2); IAHS Publication no. 256, pp. 71-80, Birmingham, UK, 1999.

Smith, S.L., M.M. Burgess, D. Riseborough, and F.M. Nixon, Recent trends from Canadian permafrost thermal monitoring network sites, Permafrost and Periglacial Processes, 16, 1-13, 2005.

Strozzi, T., A. Luckman, T. Murray, U. Wegmüller, and C.L. Werner, Glacier Motion Estimation Using SAR Offset-Tracking Procedures, IEEE Transactions on Geoscience and Remote Sensing, 40 (11), 2384-2391, 2002.

Sugden, D.E., and B.S. John, Glaciers and Landscape, 376 pp., Edward Arnold, New York, NY, USA, 1990.

Summerfield, M.A., Global geomorphology. An introduction to the study of landforms, 537 pp., Longman Scientific and Technical, New York, NY, USA, 1991.

Toutin, T., Multi-source data fusion with an integrated and unified geometrc modelling. EARSELJ, Journal of Advances in Remote Sensing, 4 (2), 12, 1995.

Trenhaile, A.S., The geomorphology of Canada: An introduction, 240 pp., Oxford University Press, Toronto, Ontario, Canada, 1990.

Trishchenko, A., 2005.

Wang, S., W. Chen, and J. Cihlar, New calculation methods of diurnal distributions of solar radiation and its interception by canopy over complex terrain, Ecological Modelling, 155, 191-204, 2002.

Williams, R.S.J., D.K. Hall, and C.S. Benson, Analysis of glacier facies using satellite technologies, Journal of Glaciology, 37 (125), 1991.

Zhang, X., L.A. Vincent, W.D. Hogg, and A. Niitsoo, Temperature and precipitation trends in Canada during the 20th century, Atmosphere-Ocean, 38, 395–429, 2000.

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2.3 Non-Permanent (Seasonal) Ice and Snow Cover

2.3.1 Seasonal snow and ice cover

2.3.1.1 Land surface ice and snow

Many parts of Canada experience snow cover for 5-6 months of the year (Figure X). The

structure and dimensions of snow cover are highly variable in both space and time, depending on a

host of factors such as the location of storm tracks, weather conditions between storms, wind action,

surface topography and vegetation cover [CRYSYS, 2003b]. Snow cover is an integrated response to

both temperature and precipitation and exhibits strong negative correlations with air temperature in

most areas with a seasonal snow cover [SOCC, 2005d]. Snow cover is important for several reasons.

Its presence (a) provides a major store of water which is released as air temperatures rise during the

spring melt period, (b) insulates the ground, thereby preventing the soil to freeze to great depths, and

(c) significantly increases the albedo of the ground surface, thereby causing more incoming solar

radiation to be reflected back into the atmosphere. The combination of (b) and (c) mean that snow

cover dramatically alters the surface-atmosphere energy balance. As a result, soil temperatures are

usually typically 5-10ºC warmer than

mean air temperatures when the ground

is covered by snow. The influence that

snow cover has on agriculture (e.g.

depth of frost penetration and soil

moisture recharge), water resource

management (e.g. changes in the timing

and amount of water from snowmelt)

and ecosystems (alterations in the sub-

nival environment for plants and

animals) mean that predicting the

response of snow cover to climate

change has become a priority for many

scientists [CRYSYS, 2003b].

Figure X. Mean duration of snow cover (days) over the 1972-94 period, as computed from satellite-derived maps of weekly snow cover extent. Source: R. Brown, Environment Canada

(data supplied by D. Robinson, Rutgers University).

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2.3.1.2 Lake and river ice

The timing and amount of lake and river ice cover is greatly influenced by regional climate

forcing. The timing of lake and river icing and melting, and hence annual ice cover duration, is

controlled by air temperature, snow and ice thickness, surface albedo and mechanical action (i.e.

wind) [Hofmann et al., 1998]. Lake ice is important for a number of reasons. Its presence (a)

influences surface-atmosphere interactions, (b) cuts the lake water off from direct atmospheric

influences, thereby restricting the movement of gases across the water-air interface and, in doing so,

altering the water’s chemical characteristics, and (c) determines the amount of sunlight that can

penetrate into the water below [CRYSYS, 2003a]. River ice is important for a number of reasons.

Like lake ice, its presence regulates flow aeration and oxygen concentrations under the ice

[Chambers et al., 1997; Prowse, 1994] and water chemistry [Cunjak et al., 1998]. However, river

ice-induced flooding also supplies the flux of sediment, nutrients and water that are essential to the

health of freshwater delta ecosystems [Lesack et al., 1991; Prowse and Conly, 1998]. Changes in ice

freeze and breakup will thus have numerous ecological impacts, including considerable impacts on

aquatic biota and productivity. As a result of these influences, predicting lake and river ice responses

to climate change is important.

2.3.1.3 Sea ice

The timing and amount of sea ice cover is directly controlled by the atmosphere (temperature,

radiation, wind) and the ocean (heat transport and mixing, and surface currents) [Stocker et al.,

2001]. The annual chronology of sea ice formation and decay is described in detail by Environment

Canada [Environment Canada, 2001a; 2001b]: The formation of sea ice begins in mid-September in

the Canadian Arctic and advances southward through to the onset of winter. Sea ice begins to form in

the St. Lawrence estuary around January 1st and advances from coastal inlets into the Gulf of St.

Lawrence. Sea ice in Canada normally reaches a maximum extent at the beginning of March. At that

time, sea ice is usually present in coastal waters of Canada except for those of British Columbia

where warm ocean currents from the south prevent the formation of sea ice. The melting of sea ice

begins in spring in the Gulf of St. Lawrence and East Newfoundland, retreating northward towards

the Labrador coast. In June openings appear in Baffin Bay and the Beaufort Sea, while clearing is

already underway in Hudson Bay. Break-up continues throughout the summer months, reaching a

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minimum extent around mid-September. The Arctic Ocean sea ice cover thus evolves from a highly

reflective snow covered surface with few openings in May to a lower-reflecting decaying sea-ice

cover, mottled with melt ponds and interrupted by many openings in July [Stocker et al., 2001]. Sea

ice is important for several reasons. Its presence (a) moderates heat exchange between the ocean and

atmosphere at high latitudes, especially by controlling the heat flux through openings in the ice; (b)

feedbacks to the broader climate system via the ice albedo feedback, which amplifies projected

climate warming at high latitudes, and by oceanic feedbacks involving ice growth and melt and the

fresh water balance at the ocean surface, and (c) leads to the transport of fresh water from the Arctic

southward, influencing the density structure of the upper ocean in the Nordic, Labrador and Irminger

Seas [Curry and Webster, 1999]. The role of sea ice in enhancing climate warming sensitivity in the

polar regions makes its dynamics particularly important to understand.

2.3.2 Snow and ice cover responses to past and present climates

2.3.2.1 Land surface ice and snow

Like all components of the climate system, snow cover exhibits considerable variation from one

year to the next in response to the natural variability of atmospheric circulation patterns which affect

both snowfall and temperature [SOCC, 2005c]. Sensitivity studies using climate data collected

during the 20th century show a strong inverse correlation between northern hemispheric air

temperatures and corresponding snow cover extent [Brown, 2000; Groisman et al., 1994]. This

response is thought to be caused by a positive albedo feedback mechanism, where a decrease in snow

cover leads to the absorption of more solar radiation at the earth’s surface, leading to faster snow

melt, and ultimately, even less snow. The albedo feedback mechanism is strongest in the spring

period when incoming solar radiation input to the snowpack is highest, and exerts a measurable

feedback to the earth radiative balance [Groisman et al., 1994].

However, North American snow cover data does not reveal the same clear temperature response

as that described for the Northern Hemisphere. This is because winter snow cover extent and snow

water equivalent has increased in parts of North America and southern Canada [Brown, 2000] in

response to higher winter precipitation [Frei and Robinson, 1995; Groisman and Easterling, 1994;

Groisman et al., 1994; Hughes and Robinson, 1993; Mekis and Hogg, 1999; SOCC, 2005c]. As a

result, North American snow cover extent tends to be more closely correlated to atmospheric

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circulation patterns than large-scale temperatures [Brown, 1995; Brown and Braaten, 1998].

Canadian daily snow depth observations (1955-) and satellite observations (1972-) reveal that snow

cover over Canada increased during the 20th Century, reaching a maximum in the mid-1970's. The

1980s and 1990s were characterized by a rapid reduction in snow cover over many regions of

Canada, particularly over western Canada, where spring melt has advanced at a rate of close to half-

a-day per year over the period since 1955 [Figure X, SOCC, 2005c]. In these regions, a reduction of

snowpack water storage and earlier spring melt is expected to bring about a lower fresh water pulse

for recharge of soil moisture and reservoirs, and increased potential for evaporation loss. These

observed changes are consistent with climate model simulations forced with increasing levels of

greenhouse gases [SOCC, 2005d].

2.3.2.2 Lake and river ice

The earliest observations on the timing of lake and river freezing and thawing in the Northern

Hemisphere predate observations made by scientists by many years. These measurements, which

were taken primarily for religious, cultural and transportation reasons, can provide a seasonally

integrated view of river ice dynamics in regions where early temperature measurements are sparse

[Magnuson et al., 2000]. Because, ice freeze-up and break-up changes, on average, by approximately

1 day for each 0.2°C change in air temperature, lake and river ice information can provide a valuable

proxy record of climate change [SOCC, 2005a]. Magnuson et al., [2000] used northern hemisphere

lake and river ice observations spanning a 150 year period (1846-1995) to show that changes in

freeze dates averaged 5.8 days per 100 years later, and that changes in breakup dates averaged 6.5

days per 100 years earlier, translating to an average lengthening of the ice-free season by 12.3 days

per 100 days and a 1.2ºC increase in air temperature per 100 years. This lengthening of the ice-free

season is consistent with lake ice records from North America, as observed for Wisconsin lakes

(1968-1988; Anderson et al.,, [1996]), the Experimental Lakes Area in northern Ontario (1970-;

Schindler at al.,, [1996], the Great Lakes (1870-1940, Williams [1971]; 1965-1990, Hanson et al.,

[1992]) lakes and reservoirs in southern Canada and the upper-midwest USA (1980-1994; Wynne et

al., [1998]), and rivers in western Canada (1969-1996, 1957-1996 and 1947-1996; Zhang et al.,

[1999]), where the sensitivity of the ice cover cycle to air temperature fluctuations generally

69

increases with latitude [Walsh, 1995]. However, it should be noted that Zhang et al., [1999] also

found that rivers in eastern Canada experienced later ice breakup over the time periods analyzed.

2.3.2.3 Sea ice

Our understanding of historical trends in sea ice extent and concentration (the fraction of local

area covered by ice) over the past half century are based primarily on regional ice charts (pre-1970s)

and satellite observations (1970s-). While limited, these observations show (a) an annual cycle,

where sea ice extent varies from a maximum of about 15 million square kilometers in March to a

minimum of about 7 million square kilometres in September, and (b) a longer-term trend, where sea

ice extent and thickness declines from the 1970s to their present day values (a 9% and 40% decline,

respectively) (Figure X; Stocker et al., [2001] and SOCC, [2005b]). The observed sea ice retreat and

thinning in the Arctic spring and summer over the last few decades is consistent with an increase in

spring temperature, and to a lesser extent, summer temperatures at high latitudes. However, due to

limited sampling, uncertainties are difficult to estimate, and the influence of decadal to multi-decadal

variability cannot yet be assessed [Stocker et al., 2001].

While it is generally accepted that Arctic Sea Ice extent has reduced over recent decades, on a

more local scale sea ice extent has actually increased in some areas, such as along the Newfoundland

coast. Such variability reflects important changes in the regional climate of eastern Canada and the

western North Atlantic [SOCC, 2005b]. Sea ice observations taken off the coast of Newfoundland

show below-normal sea ice extents from about 1920 to 1970, then an increase to normal or above-

normal extents in the 1980s and 1990s. The causes of these changes have generally been attributed to

large scale atmospheric circulation and air temperature over the North Atlantic reflected in the North

Atlantic Oscillation (NAO) [Chapman and Walsh, 1993; Fang and Wallace, 1993; Jacobs and

Newell, 1979; Marko et al., 1994; Prinsenberg et al., 1997], although other causes, such as salinity

anomalies in the North Atlantic [Mysak et al., 1990] and solar activity [Hill and Jones, 1990] have

also been investigated.

70

Figure X: Monthly variability of sea ice extent over the Northern Hemisphere since 1954 [SOCC, 2005b].

71

2.3.3 Predicting snow and ice response to future climates

2.3.3.1 The importance of predicting snow and ice response to future climates

The pattern of snow cover, as determined by snow accumulation and snowmelt processes, has a

significant impact upon climate processes, surface hydrological cycles and ecological processes

within northern biomes [Simic et al., 2004]. Snow cover exerts a dominating influence on the energy

budget of the lower atmosphere and the surface. Snow influences the surface-atmosphere radiation

balance in three ways [CIRES, 2004]. First, it has the highest albedo of any natural surface and is

able to reflect a considerable amount of solar radiation that would otherwise be absorbed at the

ground surface, heating the ground and atmosphere above. Second, it has a higher emissivity than

snow-free ground, resulting in the further reduction of temperature due to enhanced infrared cooling.

Third, it represents a significant heat sink during the spring melt period because of its relatively high

latent heat of fusion. Snow cover is also strongly linked to basin-averaged snowmelt [Luce et al.,

1998] and local hydrological processes [Cherkauer et al., 2003]. The timing of the transition

between the snowmelt and leaf appearance period is critical for terrestrial ecosystem functioning and

management of both understory and overstory vegetation, and has an effect on annual net ecosystem

productivity. Therefore, the realistic simulation of snow cover in climate and hydrologic models is

critical for the accurate representation of surface energy balance, as well as for predicting winter

water storage and year-round runoff [CIRES, 2004]. However, because in situ observations are rarely

sufficiently dense to calibrate, execute and validate these models, satellite remote sensing data are

often used to provide the spatially integrated data sets needed for these purposes [CIRES, 2004].

2.3.3.2 Remote sensing of ice and snow dynamics

The in situ measurement of snow is a labour intensive process and only provides a small sample

of the snow cover over an area. Hence, major use of active and passive microwave satellite data has

been undertaken due to their all-weather ability to map snow water equivalent over selected types of

terrain [CRYSYS, 2003b]. The unique spectral characteristics of snow are usually used in snow-cover

classification algorithms to map snow cover extent [Maurer et al., 2003; Xiao et al., 2002].

The accuracy of snow-cover maps is of particular importance in remote sensing applications to

hydrological and coupled hydrological–atmospheric models. The different spatial resolutions,

72

geographic extents and snow classification algorithms affect the accuracy of satellite-based snow-

cover maps. Cloud–snow confusion is one of the major impediments for snow classification. Certain

types of cloud, such as cirrus, low stratus, and small cumulus, are hard to discriminate from snow-

and ice-covered surfaces [Simpson et al., 1998]. Forest areas represent another obstacle to accurate

snow-cover mapping in remote sensing applications. Forest canopies obscure snow from the view of

both visible and passive-microwave satellite sensors. Ultimately, less snow is detected underneath a

forest canopy when the sensors view at off-nadir angles rather than at nadir [Hall et al., 1998]. To

enhance snow classification, Vikhamar and Solberg [2003] specified the narrow spectral ranges for

which snow is most distinctive in forest areas. The spatial scaling effect from point to pixels plays an

important role in the validation process and may explain some of the differences between satellite

and in situ snow-cover distribution. Conifer canopies block understory snow cover from solar

radiation and wind during snowmelt and, therefore, snow persists longer than in open areas where in

situ observations are commonly performed [Brown and Goodison, 1996; Vikhamar and Solberg,

2003]. In addition, snow depth and snow metamorphosis influence the reflectance of snow and the

validation process [Romanov et al., 2003; Vikhamar and Solberg, 2003].

The constant comparison of satellite-derived snow maps and surface measurements is vital for

improvement of snow-mapping algorithms. Yet, a lack of ground measurements commonly results in

two major limitations: (1) the assessments are performed within small areas, which have available

local surface measurements, and/or (2) the assessments are based on other satellite data, primarily

Landsat. Maurer et al. [2003] evaluated the Terra Moderate Resolution Imaging Spectroradiometer

(MODIS) and the National Operational Hydrologic Remote Sensing Centre (NOHRSC) snow daily

products with snow surface observations over the Columbia River basin and Missouri River basin

during winter and spring of 2001. Their results suggested that the MODIS product exhibits a greater

agreement rate than the NOHRSC snow daily product. Similar results were found by Klein and

Barnett [2003] in the validation of daily MODIS snow-cover maps of the Upper Rio Grande River

basin. Xiao et al. [2002] used Landsat Enhanced Thematic Mapper data in the validation process of

the SPOT-4 VEGETATION (VGT)-derived snow cover based on the normalized difference

snow/ice index (NDSII) approach. Snow-cover dynamics were found to be consistent between the

fine-resolution data and the VGT product over an alpine region in Asia. Romanov et al. [2002]

recently validated snow-cover distribution over the North American continent. Their automated snow

mapping technique, based on Geostationary Operational Environmental Satellites (GOES) and the

73

Special Senser Microwave Imager (SSM/I) sensor measurements, was found to have 88% agreement

with surface observations from approximately 1000 meteorological stations.

Simic et al., [2004] compared three daily snow-cover products of different spatial resolution

(VGT—S1 product; MODIS, MOD10A1 product (Version 3.0); NOAA GOES+ SSM/I) with daily

surface snow depth observations during winter and spring season over Canada, then highlighted the

difference in agreement between satellite and surface snow-cover data within different land cover

types (evergreen forest, deciduous forest, herb dominated and lichen). The validation was based on

surface snow depth observations from almost 2000 meteorological stations across Canada. The

assessment was performed on a daily basis for the period of 160 days from January to June of 2000

(SPOT-4) and 2001 (MODIS and NOAA). Direct comparisons between the snow products were not

attempted because of the limited availability of the data over the same period. Therefore, the main

intention of the validation was to compare the individual satellite products with ground data and not

to each other.

This study showed that the SPOT4 VGT product was not applicable for the retrieval of snow

cover in Canada because of high omission errors. It is likely that thresholds within the VGT products

are too restrictive for mapping snow cover. Better agreement between the satellite observations and

surface measurements was seen with both the MODIS and NOAA products, ranging from

approximately 80 to 100% on a monthly basis (Figure X). Evergreen forests exhibited the lowest

percentage agreement of all land cover types. Percentage agreement was generally reduced during

the beginning of the snow season and during snowmelt, particularly within the forested areas, for

both the MODIS and NOAA snow products. As the validation of the MODIS product between sparse

and dense conifer regions indicated, the scaling differences between point snow cover at the in situ

locations and the areal snow cover over the corresponding pixel likely contributed on the order of

10% disagreement. The regional validation of the MODIS product indicated that the level of

agreement between the MODIS product and surface data was generally lower over the mountainous

and forested regions. The NOAA product showed a more equal omission–commission errors ratio

over forest areas and had a higher retrieval rate (cloud-free pixels) than the MODIS product. The

difference between the retrieval rates for the NOAA product (100%) and MODIS product (below

74

Figure X. Ratio between omission and commission errors for the SPOT VGT snow cover product (a), and the MODIS snow cover product (b) within four land cover

types (Simic et al.,[2004]).

(a)

(b)

75

40%) suggests that the use of the microwave sensor within the NOAA product is advantageous in the

performance of the product. Furthermore, the NOAA product was found to be most consistent among

land cover types.

Simic et al., [2004] concluded that although additional assessments are necessary in order to

evaluate the snow products in applications such as hydrological modelling and, therefore, to evaluate

the effectiveness of the increased MODIS resolution, the results in this study indicate that the coarser

resolution NOAA product may be sufficient for basin- or sub-basin-scale applications. The

difference in spatial resolutions between the NOAA and MODIS products did not result in

substantially different agreement trends, and the NOAA product did not exhibit considerable

variability in agreement statistics as a function of forest cover density.

Figure X. Ratio between omission and commission errors for the NOAA GOES+ SSM/I snow cover product (c) within four land cover types

(Simic et al.,[2004]).

(c)

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2.3.3.3 A modeling approach

2.3.3.3.1 Land surface ice and snow

Because of the close association between air temperature and snow cover extent, most GCM

simulations project widespread reductions in snow cover over the next 50-100 years in response to

greenhouse gas-related climate warming [SOCC, 2005d]. However, there is considerable variability

among model-predicted regional patterns of snow cover change, largely because of the difficulties

associated with simulating precipitation over high elevation and high latitude regions and the role of

atmospheric circulation patterns [Frei et al., 2003]. For example, Moore and McKendry [1996]

showed that snowpack conditions in southern British Columbia were dominated by atmospheric

circulation patterns linked to decadal-scale variability in Pacific Ocean sea surface temperatures.

They also found that an abrupt change to less winter snow accumulation after 1976 – resulting in

reduced snow cover, earlier runoff and more negative glacier mass balances over much of western

North America – coincided with a shift in the Pacific-North America (PNA) teleconnection pattern to

more positive values [SOCC, 2005d]. Similarly, Brown [1998] showed that ENSO was responsible

for significant anomalies in regional snow cover over western Canada. The results of this study

demonstrated that El Niño was associated with below-average winter snow cover extent, while La

Niña was associated with above-average snow extent [SOCC, 2005d]. Such abrupt shifts in

atmospheric circulation and a recent tendency toward more frequent El Niño events add an additional

level of uncertainty onto the regional snow cover response to large scale climate warming [SOCC,

2005d].

Climate models project rising air temperatures to have a considerable impact on the amount and

timing of snow and ice extent across Canada’s land surface. The CCCma first and second generation

global coupled models (CGCM1 and CGCM2) have been used to produce climate simulations from

1900 to the present and projections from the present to the end of the next century [Hengeveld,

2000]. These models make projections using a scenario for future concentrations of greenhouse gases

and sulphate aerosols similar to the Intergovernmental Panel on Climate Change (IPCC) IS92a GHG

+ A emission scenario. In the CGCM1 and CGCM2 scenarios, CO2 concentrations increase by 1 %

per year after 1990, and sulphate aerosol concentrations increase until 2050 [Meteorological Service

of Canada, 2003]. While these models project a northward recession of the snowline in North

America as acclimate warms, and with it, a decrease in snow cover extent in southern and western

77

Figure X. Observed and modeled variation of annual averages of Arctic sea-ice extent, based on Vinnikov et al. [1999]. Observed data are from Chapman and Walsh [1993] and Parkinson et al. [1999]. Sea-ice curves

are produced by GFDL low-resolution R15 climate model and by HADCM2 climate model, both forced by CO2 and aerosols.

Figure X. Canadian Centre for Climate Modelling and Analysis (CCCma) prediction of sea ice extent during the 21st century [Hengeveld, 2000].

78

Canada, it is not so clear what other aspects of the snow cover may do [SOCC, 2005d]. For example,

increased precipitation may lead to increased snow accumulation in cold climate regions [Brown,

2000], while warming will be accompanied by increased frequencies of mixed precipitation and rain-

on-snow events which have implications for snowmelt, snow depth and snow density [SOCC,

2005d]. Nonetheless, despite these projections, the natural variability of the climate system and

atmospheric circulation will ensure that snow is still an important component of the winter climate in

many regions of Canada [SOCC, 2005c].

2.3.3.3.2 Lake and river ice

Models also predict that climate warming will have a considerable impact on the amount and

timing of lake and river ice cover [Hofmann et al., 1998]. Assel [1991] modeled lake ice cover for

Lake Erie and Lake Superior under a CO2-doubling warming scenario, and projected a decrease in

mid-lake ice formation and a reduction in ice cover duration of 5-13 weeks on Lake Superior and 8-

13 weeks on Lake Erie [Bruce et al., 2000]. These projections are consistent with many observations.

Earlier spring break ups attributed to warmer spring air temperatures and below average snow covers

have been observed for Wisconsin lakes (1968-1988; Anderson, [1996]), the Experimental Lakes

Area in northern Ontario (1970-; Schindler, [1996] and the Great Lakes (1870-1940, Williams

[1971]; 1965-1990, Hanson et al., [1992]). Temperate regions with intermittent river ice cover, such

as British Columbia and southwestern Ontario, may experience more intermittent periods of ice, or

even total ice disappearance. In regions where ice is more permanent, such as the far north, a

warming climate may not raise temperatures enough to cause winter ice break-up, but will likely

bring about a significant decrease in ice thickness [Clair et al., 1997] and lengthen the ice-free

season [Hofmann et al., 1998].

2.3.3.3.3 Sea ice

Sea ice is particularly difficult to simulate in climate models because it is influenced directly by

atmospheric and the oceanic components, and because many of the relevant processes require high

grid resolution or must be parameterized [Stocker et al., 2001]. While coupled climate model

projections of the future distribution of sea ice differ quantitatively from one model to another

79

[Cubasch et al., 2001], they agree that sea ice extent and thickness will decline over the 21st

century as the climate warms [Stocker et al., 2001]. For example, Canada’s CGCM1 model projects

major changes in sea ice coverage in the Northern Hemisphere, with annual mean coverage

decreasing by about 40% by 2050 and virtually disappearing by 2100 [Figure X, Hengeveld, 2000].

These projections are consistent with those of other models (e.g. the Hadley Centre and the

Geophysical Fluid Dynamics Laboratory (GFDL) models [Vinnikov et al., 1999]). GCM simulations

for Arctic sea ice predict that warming will cause a decrease in maximum ice thickness of about

0.06m per °C and an increase in open water duration of about 7.5 days per °C [Flato and Brown,

1996].

2.3.4 Implications of changes in snow and ice cover

2.3.4.1 Potential impacts of changes in land, lake and river snow and ice cover extent on Canada’s freshwater resources

Canada’s freshwater resources will likely be impacted considerably by changes in land, lake

and river snow cover extent. The observed and projected decreasing snow cover extent in western

Canada and southern Ontario will affect the timing and amount of peak water discharge in spring, as

well as the amount of water available for irrigation and industrial and domestic uses. As a result,

continued climate warming will likely have far-reaching consequences for communities, such as the

watersheds of Western Canada that provide freshwater for domestic, agricultural and industrial use in

the western Canadian prairies [Demuth and Pietronomo, 2003]. The effects of climate on Canada’s

freshwater resources are discussed in greater detail in section 2.4 (Water-dominated landscapes).

2.3.4.2 Potential impacts of changes in land, lake and river snow and ice cover extent on surface-atmosphere energy exchange

Land surface albedo is one of the most important parameters controlling the earth’s climate.

Albedo is important because it directly determines the amount of solar energy absorbed by the

ground, and hence, the amount of energy available for heating the ground and lower atmosphere and

evaporating water [Rowe, 1991]. It also affects the global climate system indirectly by controlling the

ecosystem energy, water and carbon processes that regulate greenhouse gas exchange [Wang et al.,

80

2002]. The high albedo of snow and ice means that a reduction in snow cover extent will lead to

significant changes in the optical properties of the earth's surface. It is predicted that these changes

will have a considerable impact on local surface energy balances, and ultimately, lead to further

atmospheric heating [Cess et al., 1990; Cess et al., 1991]. However, significant uncertainties exist

over the size of the effect, particularly with sea-ice [see next sub-section, Ingram et al., 1989].

Understanding and predicting the response snow cover extent to warming climate is thus an

important ongoing component of climate change research.

2.3.4.3 Potential impacts of changes in sea ice

Changes in Arctic sea ice will have a number of impacts on regional climate patterns, Arctic

ecosystems and coastal environments and communities. A decline in Arctic sea ice will (a) increase

wave heights, exposing the Arctic coast to severe weather events such as storm surges that cause

increased annual erosion, inundation, and threat to structures [Ansimov et al., 2001]; (b) result in a

greater amount of open water, which will lead to higher water evaporation, and hence, greater

precipitation [Hengeveld, 2000]; (c) have considerable impacts on Arctic biology through the entire

food chain, from algae to higher predictors such as seals and polar bears [Ansimov et al., 2001]; (c)

impact upon the albedo feedback mechanism that reflects incoming solar radiation back into the

atmosphere, thereby decreasing the amount of energy absorbed by the ground surface; and (d) favor

increased shipping along high-latitude routes, and could lead to faster and cheaper ship transport

between eastern Asia, Europe and eastern North America.

2.3.5 References

Anderson, W.L., Roberson, D. and Magnuson, J., Evidence of recent warming and El Nino related variations in ice breakup of Wisconsin Lakes, Limnology and Oceanography, 41 (5), 815-821, 1996.

Ansimov, O., B. Fitzharris, J.O. Hagen, R. Jefferies, H. Marchant, F. Nelson, T. Prowse, D.G. Vaughan, I. Borzenkova, D. Forbes, K.M. Hinkel, K. Kobak, H. Loeng, T. Root, I. Shiklomanov, B. Sinclair, P. Skvarca, Q. Dahe, and B. Maxwell, Polar Regions (Arctic and Antarctic), in Climate Change 2001: Impacts, Adaptation and Vulnerability, pp. 803-841, Cambridge University Press, Cambridge, U.K, 2001.

Assel, R., Implications of Carbon Dioxide Global Warming on Great Lakes Ice Cover, Climatic Change, 18 (4), 377-396, 1991.

81

Brown, R.D., Spatial and temporal variability of North American snow cover, 1971-1992, in Proceedings of the Eastern Snow Conference, 52nd Annual Meeting, pp. 69-78, Toronto, ON, Canada, 1995.

Brown, R.D., El Niño and North American snow cover, in Proceedings of the 55th Eastern Snow Conference, pp. 165-172, Jackson, NH, USA, 1998.

Brown, R.D., Northern hemisphere snow cover variability and change, 1915–97, Journal of Climate, 13 (13), 2339-2355, 2000.

Brown, R.D., and R.O. Braaten, Spatial and temporal variability of Canadian monthly snow depths, 1946-1995, Atmosphere-Ocean, 36, 37-45, 1998.

Brown, R.D., and B.E. Goodison, Interannual variability in reconstructed Canadian snow cover, Journal of Climate, 9, 1915–1992, 1996.

Bruce, J.P., I. Burton, H. Martin, B. Mills, and L. Mortsch, Water sector: vulnerability and adaptation to climate change, pp. 144, Global Change Strategies International Inc. and Meterological Service of Canada, Ottawa, Canada, 2000.

Cess, R.D., G.L. Potter, J.P. Blanchet, G.J. Boer, A.D. Del Genio, M. Déqué, V. Dymnikov, V. Galin, W.L. Gates, S.J. Ghan, J.T. Kiehl, A.A. Lacis, H. Le Treut, Z.-X. Li, X.-Z. Liang, B.J. McAvaney, V.P. Meleshko, J.F.B. Mitchell, J.-J. Morcrette, D.A. Randall, L. Rikus., E. Roeckner, J.F. Royer, U. Schlese, D.A. Sheinin, A. Slingo, A.P. Sokolov, K.E. Taylor, W.M. Washington, R.T. Wetherald, I. Yagai, and M.-H. Zhang, Intercomparison and interpretation of climate feedback processes in nineteen atmospheric general circulation models, Journal of Geophysical Research, 95, 16601-16615, 1990.

Cess, R.D., G.L. Potter, M.-H. Zhang, J.P. Blanchet, S. Chalita, R. Colman, D.A. Dazlich, A.D. Del Genio, V. Dymnikov, V. Galin, D. Jerret, E. Keup, A.A. Lacis, H. Le Treut, X.-Z. Liang, J.-F. Mahouf, B.J. McAvaney, V.P. Meleshko, J.F.B. Mitchell, J.-J. Morcrette, P.M. Norris, D.A. Randall, L. Rikus., E. Roeckner, J.F. Royer, U. Schlese, D.A. Sheinin, A. Slingo, A.P. Sokolov, K.E. Taylor, W.M. Washington, R.T. Wetherald, and I. Yagai, Interpretation of snow-climate feedback as produced by seventeen general circulation models, Science, 253, 888-892, 1991.

Chambers, P.A., G.J. Scrimgeour, and A. Pietroniro, Winter oxygen conditions in ice-covered rivers: the impact of pulp and paper mill effluents, Canadian Journal of Aquatic Sciences, 54, 2796-2806, 1997.

Chapman, W.L., and J.E. Walsh, Recent Variations of Sea Ice and Air Temperature in High Latitudes, 74, 1, 33-47, 1993.

Cherkauer, K.A., L.C. Bowling, and D.P. Lettenmaier, Variable infiltration capacity cold land process model updates, Global and Planetary Change, 804, 1-9, 2003.

CIRES, Global snow cover products derived from satellite remote sensing: Analysis and validation of long-term climatologies and development of near real-time products from current satellite sensors, Online document (http://cires.colorado.edu/science/projects/aoms-armstrong R01.html), Last accessed February 2005, 2004.

Clair, T.A., S. Beltaos, W. Brimley, and A. Diamond, Ecosystem science and water resources, climate change sensitivities of Atlantic Canada's hydrological and ecological systems, in

82

Climate change and climate variability in Atlantic Canada, edited by R.H. Shaw, pp. 67-82, Halifax, NS, Canada, 1997.

CRYSYS, Learn more about lake ice, Online document (http://www.msc.ec.gc.ca/crysys/ education/lakeice/lakeice_edu_e.cfm), Last accessed March 2005, Cryosphere System in Canada, 2003a.

CRYSYS, Learn more about snow, Online document (http://www.msc.ec.gc.ca/crysys/ education/snow/snow_edu_e.cfm), Last accessed March 2005, Cryosphere System in Canada, 2003b.

Cubasch, U., G. Meehl, G.J. Boer, R.J. Stouffer, M. Dix, A. Noda, C.A. Senior, S. Raper, K.S. Yap, A. Abe-Ouchi, S. Brinkop, M. Claussen, M. Collins, J. Evans, G. Fischer-Bruns, G. Flato, J.C. Fyfe, A. Ganopolski, J.M. Gregory, Z.-Z. Hu, F. Joos, T. Knutson, R. Knutti, C. Landsea, L. Meams, C. Milly, J.F.B. Mitchell, T. Nozawa, H. Paeth, J. Räisänen, R. Sausen, S. Smith, R. Stocker, A. Timmermann, U. Ulbrich, A. Weaver, J. Wegner, P. Whetton, T. Wigley, M. Winton, F. Zwiers, J.-W. Kim, and J. Stone, Physical Climate Processes and Feedbacks, in Climate Change 2001: The Scientific Basis, pp. 527-582, Cambridge University Press, Cambridge, U.K, 2001.

Cunjak, R.A., T.D. Prowse, and D.L. Parrish, Atlantic salmon in winter: "the season of parr discontent", Canadian Journal of Fisheries and Aquatic Sciences, 55 (1), 161-180, 1998.

Curry, J.A., and P.J. Webster, Thermodynamics of Atmospheres and Oceans, 465 pp., Academic Press, New York, NY, USA, 1999.

Demuth, M.N., and A. Pietronomo, The impact of climate change on the glaciers of the Canadian Rocky Mountain eastern slopes and implications for water resource-related adaptation in the Canadian prairies. Phase I - Headwaters of the North Saskatchewan River Basin, pp. 96, Climate Change Action Fund - Prairie Adaptation Research Collaborative, Ottawa, Canada, 2003.

Environment Canada, Sea Ice Climatic Atlas: East Coast of Canada, 1971-2000, Canadian Government Publishing, Ottawa, ON, Canada, 2001a.

Environment Canada, Sea Ice Climatic Atlas: Northern Canadian Waters, 1971-2000, Canadian Government Publishing, Ottawa, ON, Canada, 2001b.

Fang, Z., and J. Wallace, The relationship between the wintertime blocking over Greenland and the sea ice distribution over N. Atlantic, Advances in Atmospheric Sciences, 10 (4), 453-464, 1993.

Flato, G.M., and R.D. Brown, Variability and climate sensitivity of landfast Arctic sea ice, Journal of Geophysical Research, 101, 25767–25777, 1996.

Frei, A., J.A. Miller, and D.A. Robinson, Improved simulations of snow extent in the second phase of the Atmospheric Model Intercomparison Project (AMIP-2), Journal of Geophysical Research - Atmospheres, 108 (D12), 4369-4386, 2003.

Frei, A., and D.A. Robinson, Regional signals in Northern Hemisphere snow cover during autumn and spring, in Proceedings of the Fourth Conference on Polar Meteorology and Oceanographt, pp. 127-131, American Meteorological Society, Dalla, TX, USA, 1995.

83

Groisman, P.Y., and D. Easterling, Variability and trends of total precipitation and snowfall over the United States and Canada, Journal of Climate, 7, 184-205, 1994.

Groisman, P.Y., T.R. Karl, R.W. Knight, and G.L. Stenchikov, Changes of snow cover, temperature and radiative heat balance over the Northern Hemisphere, Journal of Climate, 7 (11), 1633-1656, 1994.

Hall, D.K., J.L. Foster, D.L. Verbyla, A.G. Klein, and C.S. Benson, Assessment of Snow-Cover mapping accuracy in a variety of vegetation-cover densities in Central Alaska, Remote Sensing of Environment, 66 (2), 129-137, 1998.

Hanson, H., C. Hanson, and B. Yoo, Recent Great Lakes ice trends, Bulletin of the American Meteorological Society, 73 (5), 577-584, 1992.

Hengeveld, H.G., Projections for Canada's climate future: A discussion of recent simulations with the Canadian Global Climate Model, Climate Change Digest Series, CCD 00-01, pp. 27, Environment Canada, Ottawa, ON, Canada, 2000.

Hill, B.T., and S.J. Jones, The Newfoundland Ice Extent and the Solar Cycle from 1860 to 1988, Journal of Geophysical Research, 95 (C4), 5385-5394, 1990.

Hofmann, N., L. Mortsch, S. Donner, K. Duncan, R. Kreutzwiser, S. Kulshrestha, A. Pigott, S. Schellenberg, B. Schertzer, and M. Slivitzky, Climate Change and Variability: Impacts on Canadian Water. Volume VIII of the Canada Country Study, pp. 1-120, Government of Canada, Ottawa, Canada, 1998.

Hughes, M.G., and D.A. Robinson, Creating temporally complete snow cover records using a new method for modelling snow depth changes., Vol. 25, Glaciological Data, World Data Center A for Glaciology (Snow and Ice), 150-163, 1993.

Ingram, W.J., C.A. Wilson, and J.F.B. Mitchell, Modelling climate change: an assessment of sea ice and surface albedo feedbacks, Journal of Geophysical Research, 94, 8609-8622, 1989.

Jacobs, J.D., and J.P. Newell, Recent year-to-year variations in seasonal temperatures and sea ice conditions in the eastern Canadian Arctic, Arctic, 32 (4), 345-354, 1979.

Klein, A.G., and A.C. Barnett, Validation of daily MODIS snow cover maps of the Upper Rio Grande River basin for the 2000–2001 snow year, Remote Sensing of Environment, 86, 162-176, 2003.

Lesack, L.F.W., R.E. Hecky, and P. Marsh, The influence of frequency and duration of flooding on the nutrient chemistry of Mackenzie Delta lakes, in Mackenzie Delta, Environmental Interactions and Implications of Development. NHRI Symposium No. 4, edited by J. Dixon, pp. 19-36, National Hydrology Research Institute, Environment Canada, Saskatoon, SK, Canada, 1991.

Luce, C.H., D.G. Tarboton, and K.R. Cooley, The influence of the spatial distribution of snow on basin averaged snow melt, Hydrological Processes, 12, 1671-1683, 1998.

Magnuson, J.R., D.M. Robertson, B.J. Benson, R.H. Wynne, D.M. Livingstone, T. Arai, R.A. Assel, R.G. Barry, V. Card, E. Kuusisto, N.G. Granin, T.D. Prowse, K.M. Stewart, and V.S. Vuglinski, Historical trends in lake and river ice cover in the Northern Hemisphere, Science, 289 (5485), 1743-1747, 2000.

84

Marko, J.R., D.B. Fissel, P. Wadhams, P. Kelly, and R. Brown, Ice berg severity off Eastern North America: Its Relationship to Sea Ice Variability and Climate Change, Journal of Climate, 7, 1335-1351, 1994.

Maurer, E.P., J.D. Rhoads, R.O. Dubayah, and D.P. Lettenmaier, Evaluation of the snow-covered area data product from MODIS, Hydrological Processes, 17, 59-71, 2003.

Mekis, É., and W.D. Hogg, Rehabilitation and analysis of Canadian daily precipitation time series, Atmosphere-Ocean, 37 (1), 53-85, 1999.

Meteorological Service of Canada, Climate Models, Online document (http://www.msc-smc.ec.gc.ca/education/scienceofclimatechange/understanding/climate_models/index_e.html), 2003.

Moore, R.D., and I.G. McKendry, Spring snowpack anomaly patterns and winter climatic variability, British Columbia, Canada, Water Resources Research, 32 (623-632), 1996.

Mysak, L.A., D.K. Manak, and R.F. Marsden, Sea-ice anomalies observed in the Greenland and Labrador Seas during 1901-1984 and their relation to an interdecadal Arctic climate cycle, Climate Dynamics, 5, 111-133, 1990.

Parkinson, C., D. Cavalieri, P. Gloersen, H. Zwally, and J. Comiso, Arctic sea ice extents, areas, and trends, 1978-1996, Journal of Geophysical Research - Oceans, C9, 20837-20856, 1999.

Prinsenberg, S., I.K. Peterson, S. Narayanan, and J.U. Umoh, Interaction between atmosphere, ice cover, and ocean off Labrador and Newfoundland from 1962 to 1992, Canadian Journal of Fisheries and Aquatic Science, 54 (1), 30-39, 1997.

Prowse, T.D., Environmental significance of ice to streamflow in cold regions, Freshwater Biology, 32, 241-259, 1994.

Prowse, T.D., and M. Conly, Impacts of climatic variability and flow regulation on ice-jam flooding of a northern delta, Hydrological Processes, 12 (10-11), 1589-1610, 1998.

Romanov, P., G. Gutman, and I. Csiszar, Satellite-derived snow cover maps for North America: accuracy assessment, Advances in Spatial Research, 30 (11), 2455–2460, 2002.

Romanov, P., D. Tarpley, G. Gutman, and T. Carroll, Mapping and monitoring of the snow cover fraction over North America, Journal of Geophysical Research, 108 (D16), 8619, 2003.

Rowe, C.M., Modeling land-surface albedos from vegetation canopy architecture, Physical Geography, 12 (2), 93-114, 1991.

Schindler, D.W., S. Bayley, B. Parker, K.G. Beaty, D.R. Cruikshank, E. Fee, E. Schindler, and M.P. Stainton, The effects of climate warming on the properties of boreal lakes and streams in the Experimental Lakes Area, Northwestern Ontario, Limnology and Oceanography, 41 (5), 1004-1017, 1996.

Simic, A., R. Fernandes, R. Brown, P. Romanov, and W. Park, Validation of VEGETATION, MODIS and GOES+ SSM/I snow-cover products over Canada based on surface snow depth observations, Hydrological Processes, 18, 1089-1104, 2004.

Simpson, J.J., J.R. Stitt, and M. Sienko, Improved estimates of the areal extent of snow cover from AVHRR data, Journal of Hydrology, 204, 1–23, 1998.

85

SOCC, Historic Variability of Lake Ice, Online document (http://www.socc.ca/lakeice/ lakeice_hist_e.cfm), Last accessed March 2005, State of the Canadian Cryosphere, 2005a.

SOCC, Historical Variability of Sea Ice, Online document (http://www.socc.ca/seaice/ seaice_hist_e.cfm), Last accessed March 2005, State of the Canadian Cryosphere, 2005b.

SOCC, Is snow cover changing in Canada?, Online document (http://www.socc.ca/snow/ snow_synopsis_e.cfm), Last accessed March 2005, State of the Canadian Cryosphere, 2005c.

SOCC, Snow, water resources and climate change, Online document (http://www.socc.ca/snow/ variability/index.cfm), Last accessed March 2005, State of the Canadian Cryosphere, 2005d.

Stocker, T.F., G.K.C. Clarke, H. Le Treut, R.S. Lindzen, V.P. Meleshko, R.K. Mugara, T.N. Palmer, R.T. Pierrehumbert, P.J. Sellers, K.E. Trenberth, J. Willebrand, R.B. Alley, O.E. Anisimov, C. Appenzeller, R.G. Barry, J.J. Bates, R. Bindschadler, G.B. Bonan, C.W. Böning, S. Bony, H. Bryden, M.A. Cane, J.A. Curry, T. Delworth, A.S. Denning, R.E. Dickinson, K. Echelmeyer, K. Emanuel, G. Flato, I. Fung, M. Geller, P.R. Gent, S.M. Griffies, I. Held, A. Henderson-Sellers, A.A.M. Holtslag, F. Hourdin, J.W. Hurrell, V.M. Kattsov, P.D. Killworth, Y. Kushnir, W.G. Large, M. Latif, P. Lemke, M.E. Mann, G. Meehl, U. Mikolajewicz, W. O’Hirok, C.L. Parkinson, A. Payne, A. Pitman, J. Polcher, I. Polyakov, V. Ramaswamy, P.J. Rasch, E.P. Salathe, C. Schär, R.W. Schmitt, T.G. Shepherd, B.J. Soden, R.W. Spencer, P. Taylor, A. Timmermann, K.Y. Vinnikov, M. Visbeck, S.E. Wijffels, M. Wild, S. Manabe, and P. Mason, Physical Climate Processes and Feedbacks, in Climate Change 2001: The Scientific Basis, pp. 419-470, Cambridge University Press, Cambridge, U.K, 2001.

Vikhamar, D., and R. Solberg, Subpixel mapping of snow cover in forests by optical remote sensing, Remote Sensing of Environment, 84 (1 SU -), 69-82, 2003.

Vinnikov, K.Y., A. Robock, R.J. Stouffer, J.E. Walsh, C.L. Parkinson, D.J. Cavalieri, J.F.B. Mitchell, D. Garrett, and V.F. Zakharov, Global warming and Northern Hemisphere sea ice extent, Science, 286, 1934-1937, 1999.

Walsh, J.E., Long-term observations for monitoring of the cryosphere, Climatic Change, 31 (2-4), 369-394, 1995.

Wang, S., W. Chen, and J. Cihlar, New calculation methods of diurnal distributions of solar radiation and its interception by canopy over complex terrain, Ecological Modelling, 155, 191-204, 2002.

Williams, G., Predicting the Date of Lake Ice Breakup, Water Resources Research, 7 (2), 323-333, 1971.

Wynne, R.H., T.M. Lillesand, M.K. Clayton, and J.J. Magnuson, Satellite monitoring of lake ice breakup on the Laurentian Shield (1980-1994), Photogrammetric Engineering and Remote Sensing, 64, 607-617, 1998.

Xiao, X., B. Moore, X. Qin, Z. Shen, and S. Boles, Large-scale observations of alpine snow and ice cover in Asia: using multi-temporal VEGETATION sensor data, International Journal of Remote Sensing, 23 (11), 2213–2228, 2002.

Zhang, X., K.D. Harvey, W.D. Hogg, and T.R. Yuzuk, Trends in Canadian Streamflow, pp. 38, Meteorological Service of Canada, Environment Canada, Ottawa, Canada, 1999.

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2.4 Water-dominated Landscapes

2.4.1 Canada’s water-dominated landscapes

2.4.1.1 Canada’s water resources

Canada is a freshwater-rich country: it contains approximately 9% of the world's freshwater

supply, but has less than 1% of the world's total population [Lemmen et al., 2004]. This water is

distributed unevenly across Canada’s landmass as glaciers, snowpacks, permafrost, groundwater,

lakes and reservoirs, rivers, and wetlands. The non-uniform distribution of Canada’s water means

that some regions often suffer from large water deficits (i.e. droughts), while others often suffer from

large water surpluses (i.e. floods). This disparity is expected to widen as the resource becomes

increasingly stressed by increased withdrawals by an ever-growing population, and is only likely to

be further enhanced by climatic warming. As a result, the management of Canada’s water resources

needs to consider not only the effects of natural

climate variability and increased population

pressure, but also what appears to be a change in

climate brought about by human activity. This

section considers Canada’s freshwater resources in

groundwater, rivers and streams, lakes and reservoirs

and wetlands. The effects of climate change on ice

and snow are discussed in sections 2.1 (Permafrost)

and 2.2 (Glaciers) and the effects of climate change

on sea level is discussed in section 2.6 (Coastal

zones).

2.4.1.2 Groundwater

Groundwater refers to the water below the water table where saturated conditions exist

[Botkin and Keller, 1998]. Groundwater is frequently collected in large subterranean areas called

aquifers. Once collected, groundwater flows downhill with the slope of the water table. Eventually,

percolating water may leave the groundwater system and eventually drains into streams, rivers, lakes

and the oceans. Although Canada’s groundwater represents about thirty seven times the total amount

Figure X. Canada’s drainage basins (Atlas of Canada, 2004)

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of water contained in rivers and lakes in Canada, we know very little about its precise location,

extent, quantity and quality [Draper, 2002; Hofmann et al., 1998].

About 30% of the Canadian population (8.9 million people) relies on groundwater for

domestic use [Environment Canada, 2004b]. Approximately 82% of these users live in rural areas. In

many regions, wells produce more reliable and less expensive water supplies than those obtained

from nearby lakes, rivers and streams. The remaining users are located primarily in smaller

municipalities where groundwater provides the primary source for their water supply systems. For

instance, 100% of Prince Edward Island's population and over 60% of the population of New

Brunswick and Yukon rely on groundwater to meet their domestic needs [Draper, 2002]. The

predominant use of groundwater varies by province. In Ontario, Prince Edward Island, New

Brunswick, and the Yukon, the largest users of groundwater are municipalities; in Alberta,

Saskatchewan, and Manitoba, the agricultural industry for livestock watering; in British Columbia,

Quebec and the Northwest Territories, industry; and in Newfoundland and Nova Scotia, rural

domestic use. Prince Edward Island is almost totally dependent on groundwater for all its uses

[Environment Canada, 2004b].

In recent years, however, a number of events affecting groundwater quality have contributed

to a heightened public awareness and concern about the importance and vulnerability of the resource

[Environment Canada, 2004b]. Even where groundwater is not used directly as a drinking water

supply, it must still be protected. This is because it can carry contaminants and pollutants from the

land into the lakes and rivers from which other people get a large percentage of their freshwater

supply.

2.4.1.3 Lakes and Reservoirs

A lake is generally defined as an inland body of fresh or saline water, appreciable in size (i.e.

larger than a pond), and too deep to permit vegetation (excluding submergent vegetation) to take root

completely across its expanse [Environment Canada, 2004a]. Environment Canada estimates that

Canada has more than 1 million lakes, covering as much as 7.6% of the country’s area. While 578 of

these lakes are considered large (they have a surface area > 100 km2), substantial amounts of water

are contained in small to medium size lakes. Lakes in the highly resistant rocks of the Canadian

Shield tend to be clear and long-lived. By contrast, Prairie lakes, which are often formed by melted-

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out glacial deposits, tend to be shallower and contain more sediment. Lakes in the mountainous areas

of BC and the Yukon are typically confined to deep glaciated valleys. Additional and considerable

water resources are also found in more than 900 of Canada’s major reservoirs. It is estimated that the

storage capacity of Canada’s 849 largest reservoirs is sufficient to hold 25% of the volume of the

five Laurentian Great Lakes [Environment Canada, 2004a].

Lakes and reservoirs have many important ecological and economical values and functions.

The freshwater produced by these resources is critical for the ability of ecosystems to generate and

sustain services to society [Jansson et al., 1999]. The ecological services provided by lakes and

reservoirs include (a) nutrient cycling, (b) the provision of habitat for plants and animals, (c) flood

control, and (d) water purification and supply. However, water from lakes and reservoirs is also

critical for industrial society. Lakes and reservoirs provide humans with services that include water

for irrigation, drinking, industry, and dilution of pollutants, hydroelectric power, transportation,

recreation, fish, and esthetic enjoyment [Postel and Carpenter, 1997]. As a result, many water bodies

are already degraded from past uses and continue to be threatened by pollution, shoreline

development, over-fishing, the invasion of exotic species, recreational impacts and the effects of ever

increasing withdrawals. The extents of these activities have led many to believe that the natural

ecosystem services themselves may be in jeopardy [Naiman et al., 1995].

2.4.1.4 Rivers

Rivers are natural watercourses, flowing over the surface in extended hollow formations (i.e.,

channels), which drain discrete areas of mainland with a natural gradient [Herbert, 2002]. Although

rivers contain only about 0.0001% of the total amount of water in the world at any given time, they

carry water and nutrients to areas all around the globe, and in doing so, drain nearly 75% of the

earth's land surface. Rivers are thus one of the most important and dynamic components of the

hydrolological cycle. Canadian rivers flow in five major ocean drainage basins: the Pacific, Arctic

and Atlantic oceans, Hudson Bay and the Gulf of Mexico [Figure X]. The largest of these is the

Hudson Bay basin (4 million km2), while the smallest is the Gulf of Mexico basin (29,500 km2). The

Atlantic basin has the greatest average discharge at over 1000 km3 a-1 [Herbert, 2002].

Rivers have many important ecological and economical values and functions. Healthy and

well-vegetated river systems (including their riparian zones) provide a variety of ecological

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functions. These include (a) the reduction of erosion by stabilizing riverbanks in times of elevated

flow, (b) the improvement of water quality by reducing the amount of sediment and nutrients

transported to the river, (c) the retention and cycling nutrients that might otherwise be swept away,

(d) the lowering of water tables, thereby helping to stabilize river banks, (e) the reduction of the

amount of nitrate moving waterways through subsurface flow, thereby protecting downstream

aquatic ecosystems, and (f) the maintenance of animal and plant diversity by providing nourishment

and means of transport to countless organisms {source}. Rivers are also economically important.

They are used as routes for commercial navigation, as sources of water for agriculture, industry, and

power generation, and have traditionally been places to dispose of municipal and industrial waste.

Water transport remains one of the most economical means of moving raw resources in Canada, and

requires high water levels, which may cause bank erosion and disturb bottom sediments and threaten

beaches. It may also facilitate the introduction of unwanted exotic species and cause pollution. Water

withdrawals for agriculture, industry and power generation lead to a suite of problems, including

water pollution (chemical and thermal), lowered river levels, habitat destruction, and increased

sedimentation. Our reliance on these water demands means that it is important to have reliable and

predictable lake and river levels [Draper, 2002].

2.4.1.5 Wetlands

Wetlands are generally defined as land that has the water table at, near, or above the land

surface [National Wetlands Working Group, 1997]. Typically, wetlands are occupied by water-

loving vegetation such as willows, sedges, cattails, bulrushes and mosses. Where open water occurs

in wetlands, it is usually less than 2m deep [Environment Canada, 2004a]. It is estimated that

wetlands cover 14% of Canada’s land area, mostly in the arctic, sub-arctic, boreal, temperate and

mountain regions. Peatlands, which occur primarily in boreal and sub-arctic regions, are the most

common type of wetland, and comprise 85% (1.1 million km2) of all wetlands in Canada [van der

Kamp and Marsh, 2004]. Canada’s wetlands are usually subdivided into five different classes – bogs,

fens, marshes, swamps and shallow water – each with distinctive hydrologic properties [National

Wetlands Working Group, 1997]. Because wetlands occur wherever the ground surface is wet for

most of the year, they tend to occur in flat and poorly drained depressions in the landscape (although

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they can occur on slopes and higher land if they are continuously fed by water from rain, snowmelt

or groundwater seepage) [Environment Canada, 2004a].

Wetlands are recognized as having many important ecological and economic values and

functions [Government of Canada, 1991]. Through the storage and slow release of water, wetlands

provide a variety of ecological functions. They can (a) reduce peak flows during floods, (b) recharge

or discharge groundwater, (c) provide habitat for fish, shellfish, aquatic birds and other animals, (d)

help purify water by trapping sediments, toxins and heavy metals, (e) provide a buffer for inland

areas from storms and high waves (coastal wetlands only), (f) cycle nutrients, and (g) store massive

amounts of carbon [Botkin and Keller, 1998; Clair et al., 1998b; van der Kamp and Marsh, 2004].

Wetlands may also have a moderating influence on climate by maintaining regional

evapotranspiration, even during dry periods, where they occupy a large portion of the landscape [van

der Kamp and Marsh, 2004]. However, wetlands are also economically important. While wetlands

are generally viewed as valuable and productive lands for fish and wildlife, and hence recreational

use, they are also valued as potential lands for agricultural activity, mineral exploitation and building

sites. In Canada, urbanization, pollution, the construction of roads, and hydroelectric projects

contribute to wetland destruction. Agriculture is the leading cause of wetland loss in the Prairie

pothole region, wetlands in the Atlantic salt marshes, in the Great Lakes region and the St. Lawrence

valley. Urban expansion has been a problem in the St. Lawrence lowlands, the lower Great Lakes,

the Atlantic salt marshes and the Pacific estuaries. Hydroelectric projects have destroyed wetlands in

the Peace-Athabasca Delta, the Saskatchewan River Delta, the Liard River in British Columbia, the

Slave River in Alberta, James Bay in Quebec, Churchill Falls in Labrador, and the Nelson River in

Manitoba [Herbert, 2002]. The short-term gains of intensive wetland development may lead to the

permanent destruction of wetlands and their ecosystem and non-ecosystem services.

2.4.2 Potential impacts of climate change on the hydrological cycle

2.4.2.1 The hydrological cycle and its components

Climate change is expected to have numerous effects on the hydrological cycle and its

components. The components of the hydrological cycle can be subdivided into reservoirs (places

where water is stored) and pathways (mechanisms through which water is transferred from one

91

reservoir to another). The main reservoirs of the hydrological cycle are glaciers, ice caps, ice sheets,

permafrost, groundwater, lakes and reservoirs, rivers and streams, wetlands, the oceans and the

atmosphere. Its main pathways are precipitation (atmosphere → land or ocean), transpiration

(vegetation → atmosphere), evaporation (streams, rivers, lakes, land or sea → atmosphere), surface

flows (land → ocean) and subsurface flows (land → ocean). Because the individual components of

the hydrological cycle are inter-linked, changes in one component will likely also likely bring about

changes in others. However, the exact magnitudes and directions of these changes remain unclear.

This uncertainty is mainly due to our limited ability to model changes in precipitation patterns and

extreme events at the regional scale, and our incomplete understanding of the complex interactions

among the components of the hydroclimatic system [Lemmen et al., 2004]. The following paragraphs

summarize the general effects that climate change is expected to have on these components, as

described by the Canada Country Study [Hofmann et al., 1998] and the Third Assessment Report of

the Intergovernmental Panel on Climate Change (IPCC) [Compagnucci et al., 2001]. Regional-

specific issues are addressed in section 2.4.3 (Canadian water resources and climate change). The

effects of climate on the frozen water components of the hydrological cycle (i.e. glaciers and

permafrost) are discussed in greater detail in previous sections.

2.4.2.2 Precipitation

Precipitation is the main driver of variability in the water balance over space and time, and

changes in precipitation have very important implications for hydrology and water resources

[Compagnucci et al., 2001]. Climate warming may change the form, amount, timing, distribution,

intensity, duration and extremes in precipitation [Hofmann et al., 1998]. Increased temperatures are

predicted to affect global circulation patterns, and in doing so, change the tracks of major storms and

the regions receiving precipitation [Environment Canada, 1995]. While GCMs generally predict an

increase in precipitation at high and mid-latitudes under climate warming, the regional effects of

warming on precipitation patterns are far from clear. Indeed, while models predict that some regions

in Canada will experience higher precipitation under global warming (e.g. northern Canada,

particularly during winter [Rouse et al., 1997]), they also predict that other regions will receive less

(e.g. southern Canada). However, it should be noted (a) that models can generate considerably

different results, even for the same region (e.g. Western Canada [Cohen, 1991; Haas and Marta,

1988; Nisbet, 1989; Zaltsberg, 1990]; Central Canada [Cohen, 1986; Croley II, 1990; Sanderson and

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Smith, 1990; Singh, 1988]; Northern Canada [Ripley, 1987]), and (b) that increases in precipitation

will not necessarily increase the water availability in affected regions. This is because higher

evaporation losses due to warmer temperatures may make many areas drier, particularly in summer

(see next section; Environment Canada [1995]).

Potential changes in intense rainfall frequency are difficult to infer from GCMs, largely

because of their coarse spatial resolution [Compagnucci et al., 2001]. Nonetheless, there are

indications that climate warming will not only affect the amount of annual precipitation a region

might receive, but also the frequency of intense local rainstorms [e.g. Hennessy et al., 1997;

McGuffie et al., 1999; Noda and Tokioka, 1989]. Danard and Murty [1988] showed that in CO2-

doubling model scenarios, precipitation increases and the centres of lows deepen. The predicted

increase in storm intensity may be caused by the higher precipitable water content of a warmer

atmosphere [Hofmann et al., 1998]. Increased precipitation may affect the hydrologic cycle in

different ways: It may fall as rain instead of snow during winter, causing more immediate direct

runoff, and the disappearance of snow in regions where snowfall is currently marginal; it may bring

about decreases is snowcover where it falls on snow; and it may bring about changes in the intensity,

timing and magnitude of spring runoff [Wittrock and Wheaton, 1992].

The accumulation and storage of winter precipitation as snowcover has an important role in

the hydrology of many Canadian regions [Hofmann et al., 1998]. Particularly, it serves to reduce

winter runoff by storing more precipitation within the snowpack. In such regions, spring melt

produces large peak flows that have important hydrological, water quality and ecological effects

[Hofmann et al., 1998]. Warmer winter temperatures will likely decrease snowpack accumulation

during winter, leading to an earlier disappearance of the snowpack in spring and, by doing so,

shortening the length of the snowcover season [Boer et al., 1992; Brown et al., 1994]. Warmer

temperatures are also expected to lengthen the open-water season on large lakes. This may increase

the lake-effect storm season in some parts of the country [Hofmann et al., 1998].

2.4.2.3 Evaporation / Evapotranspiration

Evaporation and evapotranspiration are predicted to increase under a warming climate for

most regions in Canada [Hofmann et al., 1998]. This is because a warmer climate provides (a) a

more efficient removal of water from the earth’s surface, and (b) longer ice-free seasons, longer

93

freeze-free seasons and longer growing seasons, which contribute to an extended period of

evaporation and transpiration. As a result, climate warming may lead to lower total water

availabilities, and hence, lower lake levels, stream flows, wetland levels, soil moisture and

groundwater levels [Schindler et al., 1996].

Various attempts have been made to model the effects of climate change on

evapotranspiration. These studies suggest that evapotranspiration will increase significantly with

climate warming [Byrne et al., 1989; Cohen, 1986; Croley II, 1990; Nkendrim and Purves, 1994;

Sanderson and Smith, 1993; Soulis et al., 1994; Zaltsberg, 1990]. Despite the large limitations and

uncertainties in GCM predictions of hydrological responses to climate changes at the watershed

scale, it is generally believed that the higher evapotranspiration expected in many regions will offset

the increases in precipitation that are also predicted under warming scenarios [Croley II, 1990; Soulis

et al., 1994].

2.4.2.4 Groundwater

Groundwater may also be subject to a range of impacts resulting from climate change and

increased climate variability [Mimikou et al., 1991]. Thus, groundwater must be considered as a

principal component of the assessment of the impact of climate change on the hydrologic cycle

[Thomson, 1990]. The most important impact of climate change on groundwater resources is likely to

be associated with declining groundwater levels that may occur from increased evapotranspiration

and decreased precipitation [Sandstrom, 1995].

Decreased groundwater levels will have a number of significant implications. First, because

wells are usually only excavated to the minimum depth required to obtain an adequate supply of

groundwater, declining groundwater levels will cause many wells to become dry and unusable, and

others to become less productive [Compagnucci et al., 2001; Soveri and Ahlberg, 1989]. Second, a

decrease in groundwater levels may also lead to a decreased discharge to surface water bodies,

lowering base flows between flood events [Freeman et al., 1993]. However, because groundwater

flow systems operate over a wide range of geographical and temporal scales, and have differing

abilities to retain and transport water, the impacts of climate change on this resource is likely to be

delayed and dispersed [Wilkinson and Cooper, 1993]. Thus, a reduction in groundwater recharge

may not necessarily translate to an immediate and equivalent reduction in river base flow [Hofmann

94

et al., 1998]. Rather, base flow may decline more slowly than recharge, and this decline may persist

for some time following any subsequent increase in recharge. The impacts of climate change on

groundwater resources may also not be immediately detectable, and thus, the effectiveness of

adaptation strategies targeted toward maintaining groundwater resources may be difficult to monitor

[Hofmann et al., 1998].

Climate change is also expected to have a significant impact on groundwater quality

[Hofmann et al., 1998]. The possible water quality impacts of climate change include changes in

groundwater geochemistry, resulting from an increased mineral solubility brought about by CO2-

enriched precipitation, as well as broader water quality issues brought about by changes in water

temperature [Kukkonen et al., 1994], chemistry [Webster et al., 1990] and the assimilative capacity

of lakes and rivers and lakes [Crowe, 1993]. Temperature fluctuations may have an adverse affect on

aquatic organisms that are sensitive to temperature and dissolved oxygen content [Meisner et al.,

1987]. Reduced base flow will also decrease the dilution of natural and anthropogenic contaminants

within surface bodies [Jury and Gruber, 1989].

While the above impacts are considered direct impacts – that is, impacts that are

approximately proportional to the extent of climate change – groundwater resources may also be

impacted indirectly by climate change, and these impacts may not be proportional to the actual extent

of climate change [Hofmann et al., 1998]. For example, a warmer climate will result in longer

growing seasons, and possibly a shift in agricultural practices towards crops that require irrigation in

regions where irrigation is not presently required. Irrigation consumes a large amount of water, and

may affect groundwater resources to a much greater degree than that predicted solely on the basis of

direct climate change impacts such as reduced discharge [Howitt and M'Marette, 1990].

2.4.2.5 Lakes and reservoirs

Lakes are predicted to be extremely vulnerable to changes in climatic parameters. Variations

in air temperature, precipitation and other meteorological variables directly cause changes in

evaporation, water balance, lake level, ice events, hydrochemical and hydrobiological regimes and

the entire lake ecosystem [Compagnucci et al., 2001]. It is predicted that under some climatic

conditions, lakes may disappear completely. Here, we briefly consider the potential effects of climate

change on large lakes, small lakes and reservoirs in Canada.

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It is generally expected that the net water supplies and lake levels in the large lakes will be

affected by climate change [Hofmann et al., 1998]. Changes in water levels are expected to follow

annual, seasonal and short-term patterns. Seasonal cycles are superimposed on annual lake levels.

Lakes levels are generally at a minimum in January or February. They then rise due to snowmelt and

spring precipitation, reaching a maximum in summer, after which they begin their seasonal decline

[Hofmann et al., 1998; Magnuson et al., 1997]. Short-term variations in lake levels are caused by

storm surge and set-up. Small lake systems are also expected to be significantly affected by a

warming climate, especially in the north [Schertzer et al., 2004]. Small lakes are particularly

vulnerable to a warming climate since because they are more likely to dry out completely during

serious. When lakes are saline, changes in water levels can significantly affect the salinity of the

lakes, as well as their composition of flora and fauna [Hammer, 1990].

Canada’s reservoirs will also be affected by a changing climate. Case studies of various

reservoirs in Canada (Greater Vancouver Reservoir, BC; Grand River, Ontario; Lake Diefenbaker,

Saskatchewan; Churchill Falls, Labrador) suggest that climate warming may lower reservoir levels

by reducing mountain snowpacks and increasing evaporative losses and consumptive water use (e.g.

domestic, industrial and agricultural withdrawals)[Schertzer et al., 2004]. However, in many regions,

the increased precipitation brought about by warming may offset these losses to some degree. Such

changes will likely have large implications where reservoirs are used for power generation. Long-

term changes in the amount and timing of precipitation may affect overall generation capability by

causing problems with turbine capacity during early spring peak flows and less hydroelectric

generating capacity in summer to cope with greatest market demand [Schertzer et al., 2004].

2.4.2.6 Streams and runoff

The impacts of climate change on runoff and stream flow in temperate regions are expected

to be significant. Climate change will likely have an effect on the magnitude of the mean, minimum

and extreme flows as well as their temporal distributions and variation. The most important climate

change effect in temperate regions is predicted to be the timing of streamflow throughout the year

[Compagnucci et al., 2001]. This may include (a) changes in winter runoff due to more winter

rainfall caused by warmer winter temperatures; (b) decreases in the volume of spring runoff due to

reduction in winter snow cover; (c) an earlier onset of spring melt because of earlier spring warming;

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(d) a decline in summer and fall flow rates due to higher evapotranspiration and reductions in

groundwater base flow contributions; (e) longer summer and fall low flow periods, and (f) increased

variability in annual flow [Hofmann et al., 1998; Leavesley, 1994].

However, predictions of the actual nature vary widely [see the Canadian Impact Assessments

and Hofmann et al., 1998]. Simulated annual stream flow for small catchments is often barely

affected by temperature changes, instead leading to larger and earlier winter runoff, and increased

winter/spring stream flow peaks [Ng and Marsalek, 1992]. Model results also suggested that the

effects of precipitation fluctuations are more direct, with annual and seasonal stream flow changes

becoming more directly proportional to precipitation changes. Lower lake levels and flow rates

produced by simulations reflect the increased evapotranspiration that occurs under climate warming

conditions [Hofmann et al., 1998].

The predicted increased variability in annual flow is difficult to predict since it depends not

on the temperature increase but on the change in precipitation [Compagnucci et al., 2001]. This,

combined with the expected increases in precipitation intensity, will likely bring about larger and/or

more frequent and severe floods [Hofmann et al., 1998]. These patterns will alter the physical

characteristics, water quality and biological communities of aquatic ecosystems. Physical changes to

stream and river environments include the increased erosion of banks and an increased sediment

loading and sedimentation of channel bottoms. It is thought that these impacts will be the most

severe for streams and rivers in arid environments where vegetation is sparse [Grimm et al., 1997]

and in agricultural areas where soil is most exposed [Compagnucci et al., 2001]. Water quality will

likely be altered in several ways. Increases in the severity of summer droughts will possibly bring

about lower dissolved oxygen levels and higher concentrations of plant nutrients and contaminants

[Schindler et al., 1996]. Increased flood sizes will bring about larger loadings of sediments, nutrients

and pollution from agricultural and urban regions, which will have detrimental effects on stream

quality and river organisms [Hofmann et al., 1998]. These predicted changes will impact on the way

Canadians manage their water resources: more frequent and larger floods will likely lead to increased

expenditures for flood management and place additional pressure on public finances and the

insurance industry, while flood control structures will need to be modified to accommodate the larger

and more frequent extreme flood events [Hofmann et al., 1998].

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2.4.2.7 Wetlands

Wetlands are expected to be extremely vulnerable to the changes in the hydrologic cycle that

may accompany climate warming [Clair et al., 1998b]. The previously described changes in water

balance components, such as changes in precipitation patterns, decreases in surface runoff, lowered

snowfall, reduced groundwater storage and increased evapotranspiration are all issues of concern

[Hofmann et al., 1998].

A high water table or frequent inundation is required to maintain wetland ecosystems. The

maintenance of wetlands requires that the water supply to a wetland must exceed losses from runoff

and evapotranspiration. Any variation in climate that increases the relative importance of evaporation

compared to precipitation is likely to result in the drying out of wetlands [van der Kamp and Marsh,

2004]. Climate change is expected to alter regional hydrologic processes by modifying the amount

and quality of water supplied to wetland ecosystems, and in doing so, disrupting this balance

[Hofmann et al., 1998]. Thus, the shorter warmer winters and longer summers predicted by most

climate change scenarios indicate that Canada’s wetlands will be under increasing stress due to water

storage, unless decreases in evapotranspiration are offset by increases in precipitation [van der Kamp

and Marsh, 2004]. Climate warming may bring about changes in the frequency of extreme events as

well as a decline in the total water storage associated with climate change. Such changes may disrupt

the functioning of wetland ecosystems and impair their multifunctional values [Poiani and Johnson,

1993a; Poiani and Johnson, 1993b]. For example, wetland productivity is depressed by prolonged

flooding and drying out, but increased by periodic water level changes [van der Valk and Davis,

1978].

The various types of wetland are predicted to react and adapt differently to climate change

and variability [Hofmann et al., 1998]. Wetlands fed by deep groundwater systems – such as fens fed

mainly by groundwater – are less likely to be affected by climate change because they tend to

maintain a steady flow even under large climatic variations [van der Kamp and Marsh, 2004; Winter,

2000]. Other wetlands that are less likely to be affected by climate change include through-flow

wetlands that are maintained in a balance between large surface water inflows, and marshes along the

margins of lakes and rivers that have stable water levels [van der Kamp and Marsh, 2004]. In

comparison, bogs are particularly vulnerable to changes in climate because of their reliance on

precipitation. Precipitation is the major supply into bogs since they are isolated from the local

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groundwater regime by peat accumulation. A decline in precipitation, and a resulting drawdown of

water levels, will alter vegetation and expose the peat and sediments to aerobic conditions, increasing

oxidation and changing physical the properties and the flux of nutrients, gases and sediments [Woo,

1992]. Similarly, many ephemeral Prairie sloughs are fed only by spring snowmelt and precipitation.

During drought they dry out, and persist throughout a season with high precipitation. Semi-

permanent sloughs are fed by groundwater in addition to precipitation and spring snowmelt and only

dry out in serious drought when the groundwater storage is depleted [Hofmann et al., 1998].

2.4.3 Predicting freshwater responses to changing climate: Regional perspectives

2.4.3.1 A National Perspective

Bruce et al., [2000] provide a national- and regional-scale assessment of the impacts of

climate change on Canada’s water resources. At a national scale, it is generally agreed that higher

temperatures and small changes in precipitation will occur, leading to lower total stream flows, lower

minimum flows and lower average peak annual flows. Projections suggest that winters will be drier

from southern BC to Lake Superior and Lake Michigan-Huron basin. They also suggest that the

water availability in regions receiving higher precipitation in summer would be more than offset by

the increases in evapotranspiration brought about by a warming climate. A warming climate is also

expected to bring about the melting of glaciers in the Rocky Mountains, and as a result, a gradual

decline in the flows in the eastward flowing rivers on the Great Plains. The general impacts of

climate change on the water resources of North America are shown in Figure Xand X.

2.4.3.2 A Regional Perspective - Atlantic region

2.4.3.2.1 Location and Climate

Atlantic Canada includes the Maritime Provinces of New Brunswick, Prince Edward Island

and Nova Scotia as well as the province of Newfoundland and Labrador. These provinces contain

almost 300,000 surface water bodies and over 40,000 km of coastline [Shaw, 1997]. The climate of

the region is diverse, and is regionally influenced by latitude and proximity to major water bodies,

such as the St. Lawrence River and the Atlantic Ocean [Bruce et al., 2000].

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Figure X. Potential water resources impacts in North America [after Cohen and Miller, 2001]. Figure X. Notes: 1. [Loukas and Quick, 1999]; 2. [Taylor and Taylor, 1997]; 3. [Brugman et al., 1997]; 4. [Hofmann et al., 1998]; 5. [BESIS, 1999]; 6. [Melack et al., 1997]; 7. [Hamlet and Lettenmaier, 1999]; 8. [Cohen et al., 2000]; 9. [Wilby and Dettinger, 2000]; 10. [Leung and Wigmosta, 1999]; 11. [Wolock and McCabe, 2000]; 12. [Felzer and Heard, 1999]; 13. [Gleick and Chalecki, 1999]; 14. [Thompson et al., 1998]; 15. [Fyfe, 1999]; 16. [McCabe and Wolock, 1999]; 17. [Leith and Whitfield, 1998]; 18. [Williams et al., 1996]; 19. [Hauer et al., 1997]; 20. [Wilby et al., 1999]; 21[USEPA, 1998]; 22. [Hurd, 1998]; 23. [USEPA, 1998]; 24. [Marsh and Lesack, 1996]; 25. [Maxwell, 1997]; 26. [Rouse et al., 1997]; 27. [MacDonald et al., 1996]; 28. [Herrington et al., 1997 Canada Country Study: Impacts and Adaptations, Sectoral Volume. #1841]; 29. [Strzepek et al., 1999]; 30. [Clair et al., 1998a]; 31. [Yulianti and Burn, 1999]; 32. [Lettenmaier, 1999]; 33. [Woodhouse and Overpeck, 1998]; 34. [Evans and Prepas, 1996]; 35. [Eheart et al., 1999]; 36. [Hurd, 1998]; 37. [Mortsch and Quinn, 1996]; 38. [Chao et al., 1999]; 39. [Magnuson et al., 1997]; 40. [Moore et al., 1997]; 41. [Abraham et al., 1997]; 42. [Frederick, 1999]; 43. [Hare et al., 1997]; 44. [Mulholland et al., 1997]; 45. [Justic et al., 1996]; 46. [Arnell, 1999]; 47[Cruise, 1999]; 48. [Porter et al., 1996].

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Figure X. Potential water resources impacts in Canada (after Lemmon et al., 2004).

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2.4.3.2.2 GCM Projections

GCMs predict that climate change in Atlantic Canada will take various forms. Simulated

stream flow for small Newfoundland catchments [Ng and Marsalek, 1992] suggested that

temperature increases barely affected annual stream flow, instead leading to larger and earlier winter

runoff when precipitation was stored in the snowpack, and increased winter/spring stream flow peaks

[Bruce et al., 2000; Hofmann et al., 1998]. These simulations indicated that the effects of

precipitation fluctuations were more direct, with annual and seasonal stream flow changes becoming

more directly proportional to precipitation changes [Hofmann et al., 1998].

2.4.3.2.3 Frequency of extreme Events

It is predicted that Atlantic Canada will experience changes in the frequency and intensity of

extreme weather events such as hurricanes, snowstorms and rainstorms [Bruce et al., 2000;

Canadian Climate Impacts and Adaptation Research Network, 2004; Hare et al., 1997]. This will

likely cause an increase in flood and drought frequency. Indeed, floods and droughts have already

become commonplace in Atlantic Canada in recent years. While flooding has occurrence in

Newfoundland, the Saint John River basin (New Brunswick), the Bay of Fundy coast and

Northumberland Straits, many parts of this region have also experienced significant drought periods

[Bruce et al., 2000; Hofmann et al., 1998]. The increased flooding in these regions has been strongly

linked to warming-induced breakups in winter ice cover occurring earlier in the year, and drought

has been strongly linked to lower summer water flows in summer [see Hare et al., 1997; Ng and

Marsalek, 1992; Whitfield and Cannon, 2000; Zhang et al., 1999]. These effects may only be

exacerbated if climate warming also leads to higher winter snowfalls [Canadian Climate Impacts

and Adaptation Research Network, 2004].

2.4.3.2.4 Other Implications

Changes in the frequency and severity of flooding, heavy rain and drought stress are potential

concerns to agriculture, municipal and domestic water use patterns, and hydro-electric power

generation. Climate change may affect agriculture in a number of ways. Because the primary limiting

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factor for agricultural production in Atlantic Canada is heat, climate change may actually provide

some benefits to farmers as long as there is sufficient moisture during the growing season [Bruce et

al., 2000]. However, where such conditions are not met, farmers will likely increasingly need to rely

on irrigation, leading to local stresses on groundwater and streams [Bootsma, 1997]. More frequent

riverine flooding and severe under climate change may place municipal water infrastructure at

greater risk of failure and contamination [Bruce et al., 2000]. Atlantic Canada municipalities are

among the highest consumers of water in Canada which may lead to greater impacts should supplies

become seasonally unreliable. Climate warming will likely also have large implications for hydro-

electric power generation. Long-term changes in the amount and timing of precipitation may affect

overall generation capability by causing problems with turbine capacity during early spring peak

flows and less hydroelectric generating capacity in summer to cope with greatest market demand

[Schertzer et al., 2004].

2.4.3.3 A Regional Perspective - Quebec

2.4.3.3.1 Location and Climate

Over 10% of Quebec’s 1.5 million km2 surface area is covered by freshwater (Government of

Quebec 1999). The St. Lawrence River is the most significant component of Quebec’s water

resources because over 95% of the province’s population lives within the watershed, while 70%

reside in a 10-km strip on either side of the river’s shore (Bergeron et al. 1997). The climate of the

region is diverse, and is regionally influenced by latitude, proximity to water bodies (Gulf of St.

Lawrence, Hudson Strait, Ungava Bay), topography (Laurentians, Appalachians) and storm tracks of

extra-tropical cyclones and remnants of tropical storms [Bruce et al., 2000].

2.4.3.3.2 GCM Projections

GCMs predict that climate change in Quebec will take various forms. Modeling projections

by the Canadian Centre for Climate Modeling and Analysis (CCCma 1999) suggest that warming

will occur in most parts of the province [Bruce et al., 2000], leading to slight increases in total

annual precipitation in most regions. These climatic changes are predicted to have considerable

affects on the hydrology of the province. GCMs predict small increases in total annual runoff for

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various watersheds in Quebec, including the Moisie River basin on the north shore of the St.

Lawrence River (Slivitzky and Morin (1996, 1992)) and the La Grande River, Caniapiscau River and

Opinaca-Eastmain River watersheds in James Bay (Singh (1988)). These increases are thought to be

due to greater winter precipitation and earlier snowmelt. However, GCMs also predicted summer and

fall decreases in runoff in the Mosie River basin, which are likely due to reduced summer

precipitation and greater evapotranspiration (Hofmann et al. 1998). In comparison, models predict

much lower mean annual flows for the St. Lawrence River, the largest of the Quebec watercourses.

2.4.3.3.3 Frequency of Extreme Events

It is predicted that Quebec will experience changes in the frequency and intensity of extreme

flood events [Bruce et al., 2000]. This is because rivers in the northern part of the province that

usually remain covered by ice during the winter may begin to experience winter break-ups and

associated flooding (Clair et al. 1997), although this may be less of a problem for the rivers in the

extreme south of the province [Bruce et al., 2000]. Flood events may also be triggered by the more

previously described more frequent and intense rainfall events that are predicted to accompany

climate change.

2.4.3.3.4 Other Implications

Changes in the frequency and severity of flooding, heavy rain and drought stress are potential

concerns to agriculture, municipal and domestic water use patterns, and hydro-electric power

generation. Increases in precipitation are likely to reduce the need for crop irrigation in most regions

Quebec, with the possible exception being the south of the province where some GCM simulations

predict drier conditions overall [Bruce et al., 2000]. The pressure on Quebec’s water resources by

municipalities and rural domestic water use is less than that of the Atlantic provinces. Indeed, Bruce

et al., [2000] suggest that while reduced water flows are generally not a problem in such a water-rich

landscape, they have been identified as concerns in a few systems, such as the Saint-Charles River

that provides water to the City of Quebec. While the predicted increases in precipitation may benefit

power generation in northern Quebec through the availability of additional water supplies [Mercier,

1998], production from power stations fed from the St. Lawrence River may fall significantly if flow

levels in this river decrease [Bruce et al., 2000].

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2.4.3.4 A Regional Perspective - Ontario

2.4.3.4.1 Location and Climate

The province of Ontario is highly urbanized in the south, where it is heavily industrialized

and has a significant agricultural base. Northern Ontario is sparsely settled, and has a resource-based

economy. Snow accumulation and melt play an important role in the hydrology of the region [Bruce

et al., 2000].

2.4.3.4.2 GCM Projections

GCMs generally predict that increases in total annual precipitation and evapotranspiration

will accompany climatic warming in Ontario, and that the higher evapotranspiration will more than

offset increases in precipitation [Croley II, 1990; Sanderson and Smith, 1993]. These climatic

changes are expected to have considerable impacts on the hydrological characteristics of the Great

Lakes – St. Lawrence Basin, and especially, runoff, lake levels, lake ice, groundwater and wetlands.

2.4.3.4.3 Runoff

Mortsch and Quinn [1996] assessed the impacts of climate warming on the Great Lakes using

four GCM climate scenarios. Each of their scenarios predicted a decrease in annual runoff, a

decrease in mean annual outflow, a decrease in water levels, and an increase in lake temperature for

each of the lakes. Model projections for the Grand River Basin and the Bay of Quinte watershed both

show decreases in annual runoff with climate warming (-11 to -21% and -12%, respectively). The

timing of runoff is also generally expected to occur with this decrease. For example, in the Bay of

Quinte, snowfall was partly replaced by more winter rain, bringing about more frequent runoff

events and a lower spring snowmelt. Lower stream flows will have considerable implications on

water use and water quality issues in Ontario. Low-flow conditions are particularly difficult to

manage because of the many competing interests of waste assimilation, municipal drinking water

supply, recreation, water taking for irrigation and industrial needs as well as instream environmental

needs that must be accommodated (Brown et al. 1996; Southam et al., 1997). Furthermore, lower

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stream flows reduce assimilative capacities of streams and while pollutant loadings may remain

constant their concentrations in water will increase.

2.4.3.4.4 Water levels

Models predict decreases in water levels in Ontario with climate warming (Figure X). The

projected declines in lake levels will reduce the surface area of each lake. For example, a water level

decline of 1.6m for Lake St. Clair will decrease lake surface area by 15% [Lee et al., 1994]. The

lowering of the Great Lakes water levels will have dramatic effects on wetlands, fish spawning,

recreational boating, commercial navigation and municipal water supplies. There is a concern that

there will be less available water to meet projected increases in future demand for navigation,

consumption and water export [Schertzer et al., 2004]. Furthermore, lake level lowering will have

costly implications. Changnon [1993] estimated that the costs for dredging, changing slips and

docks, relocating beach facilities, extending and modifying water intake and sewage outfalls for a

110km stretch of Lake Michigan shoreline could reach as high as $US 827 million for a 2.5m

decrease in water levels [Bruce et al., 2000; Hofmann et al., 1998]. It is important to note that some

Figure X. Projected changes in Great Lakes water levels under a CO2-doubling CO2 using four models (Mortsch and

Quinn, 1996)

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Figure X. General Circulation Model (GCM) Scenario Impacts on the Great Lakes [in Bruce et al., 2000]. [Sources: Croley II, 1990; Hartmann, 1990; Mortsch and Quinn, 1996; Quinn, 2000]. CCC- Canadian Climate Centre - Equilibrium 2xCO2 run [Boer et al., 1992; McFarlane et al., 1992]. CCCma – Canadian Centre for Climate Modelling and Analysis – Transient Run [Boer et al., 1998]. GFDL –- Geophysical Fluid Dynamics Lab - Equilibrium 2xCO2 run [Manabe and Wetherald, 1987]. GISS – Goddard Institute for Space Studies - Equilibrium 2xCO2 run [Hansen et al., 1984; Hansen et al., 1983]. HadCM2 – Hadley Centre – Transient run. OSU – Oregon State University - Equilibrium 2xCO2 run [Schlesinger and Zhao, 1988].

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GCMs project lake level increases, these are no greater than increases resulting from natural

variability [Schertzer et al., 2004].

2.4.3.4.5 Lake Ice Cover

Climate warming is predicted to have a considerable impact on the amount and timing of Great

Lakes ice cover. Assel [1991] modeled lake ice cover for Lake Erie and Lake Superior under a CO2-

doubling warming scenario, and projected a decrease in mid-lake ice formation and a reduction in ice

cover duration of 5-13 weeks on Lake Superior and 8-13 weeks on Lake Erie [Bruce et al., 2000].

These projections are consistent with many observations. Earlier spring break ups attributed to

warmer spring air temperatures and below average snow covers have been observed for Wisconsin

lakes (1968-1988; Anderson, [1996]), the Experimental Lakes Area in northern Ontario (1970-;

Schindler, [1996] and the Great Lakes (1870-1940, Williams [1971]; 1965-1990, Hanson et al.,

[1992]). However, they are not consistent with others (1940-1971, Williams [1971]; 1965-1990 for

Lake Ontario, Hanson et al., [1992]).

2.4.3.4.6 Groundwater

Groundwater resources in Ontario are also expected to be affected by climate change. Modeling

studies project significant reductions in the rate of recharge under a warming climate. For example,

McLaren and Sudicky [1993] assessed the effects of CO2-doubling warming scenarios on

groundwater in the Grand River basin. Projections showed that climate warming will reduce recharge

by 15-35%, exacerbating groundwater supply problems. Decreases in groundwater recharge could

seriously affect rural dwellers that are usually reliant upon shallow dug wells. In some cases, wells

will become dry and unusable, while in others, wells will become less productive due to the loss of

available drawdown [Soveri and Ahlberg, 1989]. Other effects of declining groundwater resources

are less intuitive than those described above. These include changes in groundwater biogeochemistry

brought about by CO2-enriched precipitation affects on mineral solubility and temperature changes.

These temperature changes may have a detrimental impact on aquatic organisms which are sensitive

to remperature and dissolved oxygen content [Meisner, 1990].

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2.4.3.4.7 Wetlands

Great Lakes wetlands also expected to be vulnerable a decline in mean lake levels and changes in

the seasonal progression of high and low periods. Climate change is expected to alter regional

hydrologic processes by disrupting these trends. In the Great Lakes region, marshes are expected to

be able to adapt more readily to lower water levels than swamps because their dominant vegetation

could colonize quickly. Furthermore, enclosed and barrier shoreline wetlands and inland wetlands

would be vulnerable to drying [Bruce et al., 2000].

2.4.3.4.8 Other Implications

In summary, water management issues are very complex in Ontario, and especially, in the south.

Issues include flooding and erosion protection, water apportionment, protection and securing of

surface and ground water supply, and water quality protection and remediation (urban runoff,

agricultural non-point source pollution, nutrient enrichment and toxic chemicals). There are also

major hydro-power developments and significant consumptive uses including irrigation, public water

supply, industrial processes, fossil fuel and nuclear thermoelectric generation and livestock watering.

Many competing interests including human and ecosystem health and economic development must

be balanced. In the north, ecosystem effects of resource development especially forest harvesting and

hydro-power dominate water resources management concerns [Bruce et al., 2000]. All of these

sectors and issues must be addressed if the impact of climate warming on Ontario’s water resources

is to be minimized.

2.4.3.5 A Regional Perspective - Prairie Provinces

2.4.3.5.1 Location and Climate

The Prairie Provinces include Manitoba, Saskatchewan and Alberta. The climate in this region

varies from north to south and from east to west. In the southern portion of the prairies, annual

evaporation normally exceeds precipitation, making the region vulnerable to droughts and soil

moisture deficits. Here, annual streamflow is variable from year-to-year and except for the

Red/Assiniboine River system, much of the source is glacier melt and melting of snow in the

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foothills and mountains [Bruce et al., 2000]. The central and northern portion of the Prairie

Provinces receives more precipitation and generally has adequate soil moisture.

2.4.3.5.2 GCM Projections

Most GCM projections for the Canadian Prairies show considerable increases in temperature

under global warming [Hengeveld, 2000]. These increases are expected to be much greater in the

north (5-10ºC), with winters, in particular, becoming much warmer [Bruce et al., 2000]. However,

the overall change in precipitation in the prairies under climate warming is unclear. Shepherd and

McGinn [2003] and McGinn and Shepherd [2003] modeled the impact of various climate change

scenarios on the climate of the prairie region. While their projections showed an overall increase in

total annual precipitation ranging from 4 – 32% (above 1960–1989 historic values), these increases

were not temporally or geographically uniform. Rather, their projections showed that (a) increases in

precipitation were more likely to occur in the winter and summer, and (b) while model projections

suggest that increases in precipitation may occur in central Alberta, only a slight increase in

precipitation – or even a slight decrease – was projected to occur in the southern and eastern prairies.

These projections suggest that while no major change in drought frequency will occur in parts of

Alberta, this frequency could increase dramatically in the southern and eastern prairie regions. Here,

higher summer temperatures will likely increase evaporation and intensify drought conditions.

However, an increase in precipitation also means that there may also be wetter periods when

temperatures are cool. These climatic changes are expected to have considerable impacts upon the

hydrology of the Prairie Provinces, and particularly, runoff, groundwater regimes and wetlands.

2.4.3.5.3 Runoff

The impacts of climate change on runoff patterns in the Canadian prairies are intimately linked to

snowfall and snowmelt regimes in the mountains. Most prairie communities rely on meltwater

discharged from the glacierized basins of Western Canada for domestic, agricultural and industrial

use, especially during the summer months [Bjonback, 1991]. As a result, the impacts of climate

change on these basins are crucial for understanding the effects of climate change on the fresh water

resources of the Canadian prairies.

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Climate warming is expected to have a significant impact on the amount and timing of water

discharged in glacierised basins. Increases in air temperature, radiation flux and rainfall bring about

increased surface melting, leading to greater rates of runoff than would occur on a snow-free

catchment [Benn and Evans, 1998]. Consequently, meltwater discharge is generally high during

warmer periods when storage decreases and glaciers retreat, and highest when deglaciation is at its

most rapid. Discharge then declines as glaciers retreat and the supply of stored water is exhausted

[Benn and Evans, 1998]. When the glacier has completely disappeared, seasonal discharge patterns

reflect annual variations in precipitation, with a minimum occurring in the summer [Oerlemans,

2001]. The reduction in glacier size reduces the potential for meltwater yield. As a result, continued

climate warming may have far-reaching consequences for communities depending on such sources

of water during the summer months.

Nkendrim and Purves [1994] estimated the potential impact of climatic change on stream flow in

the Oldham River basin in Alberta using an analogue approach with historical records. They found

that an increase of 1°C over a five-year period combined with normal precipitation reduced stream

flow by 15% annually. Demuth and Pietronomo [2003], assessed the impact of climate change on the

glaciers of the eastern slopes of the Canadian Rocky Mountains and the subsequent implications for

water-related adaptation in the Canadian prairies. This study used data describing the marked

seasonality of the energy and moisture fluxes associated with glaciers (e.g., glacier mass balance),

and a memory function represented by changes in the areal extent of glaciation (e.g., advance/retreat

of the glacier margins), to develop both short-term and historical climate perspectives. The specific

findings of this study were: (1) a decrease in transition-to-Base-Flow (TBF) yields from the

glacierised catchments of the upper North Saskatchewan River Basin (NSRB) were observed since

the 1900s, despite generally warmer summer conditions and increased precipitation; (2) a decrease in

the minimum and mean flows from glacierised catchments commensurate with reductions in the

area-wise extent of glacier cover after accounting for variations in regional precipitation amounts and

air temperature; (3) an increased variability of stream flow from the glacierised basins of the upper

NSRB since the mid-1900s in association with decreasing glacier cover; (4) a decrease in glacier

cover since the mid 1800s, and a marked contraction of glacier cover in the later half of the 20th

century; (5) that analyses suggest that glacier cover in NSRB headwaters (eastern slopes of the

Canadian Rockies) is rapidly approaching (ca. 50 a) the state that may have existed during the early

Holocene warm interval (i.e., the warm limit of Holocene variability), and certainly a state that

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available morpho-stratigraphic and botanical evidence suggests has not existed for several millennia;

(6) that such changes can be correlated with shifts in synoptic climate. Such knowledge should assist

in the development of better forecasting tools for use in water resource management under scenarios

of climate variability.

2.4.3.5.4 Frequency of extreme Events

While climatic change in the Prairie Provinces is expected to bring about decreases in annual

peak flow because of more frequent winter melt and rains, spring rains are predicted to become more

intense [Bruce et al., 2000]. When coupled with snowmelt, these rains may bring large flood events.

Examples of such events are the May 1997 floods along the Red River.

2.4.3.5.5 Groundwater

As previously mentioned, the Canadian Prairies’ water supply is mainly surface (river) water

controlled by snow and ice melt in the Rockies. In areas where there is more reliance on

groundwater, watersheds are less well defined and the interrelationship between the land surface

characteristics and the water resource are not well documented or monitored [Bruce et al., 2000]. In

such regions, responding to the effects of climate change will be difficult because ground water is

not legally apportioned. Climate change scenarios where runoff decreases, evaporation increases, and

consumptive use of water (e.g. irrigation, hydroelectric power generation; municipal, industrial use)

increases, have the potential to lead to disagreements over water apportionment.

2.4.3.5.6 Wetlands

Ephemeral prairie sloughs are usually only fed by spring snowmelt and precipitation. These

sloughs dry our during droughts and persist in seasons where precipitation is high. Semi-permanent

sloughs are also fed by groundwater and only dry out during extreme droughts when groundwater

storage is depleted [Poiani and Johnson, 1991; Poiani and Johnson, 1993a; Poiani and Johnson,

1993b]. When these sloughs are saline, changes in water levels can significantly affect the salinity of

the lakes, as well as their composition of flora and fauna [Hammer, 1990]. Furthermore, these

112

sloughs provide critical habitat for waterfowl, whose populations correlate strongly with the number

of wetlands [Bruce et al., 2000].

2.4.3.5.7 Other Implications

Other implications of climate change on the freshwater resources of the Prairie Provinces include

decreased water quality and threats to recreation, agriculture and hydroelectric power generation.

Water quality issues in the Prairie Provinces under a changing climate are strongly linked to the

projected changes in the hydrological characteristics of the region. Many of the surface water bodies

in the prairies are shallow, saline and eutrophic, and these problems are exacerbated under drought

conditions. Recreational activities are also predicted to be affected by a changing climate. Warmer

temperatures and the previously noted decreases in runoff will lower lake and reservoir levels,

decrease water quality and increase algal problems [Bruce et al., 2000]. Agriculture is highly

sensitive to reduced water supplies, and will suffer from lowered surface water and groundwater

levels. Hydroelectric power generation will also be affected by more frequent droughts, reductions in

river flow and reservoir levels.

2.4.3.6 A Regional Perspective - The Arctic and the North

2.4.3.6.1 Location and Climate

The Arctic and the north includes the Yukon, Northwest Territories and Nunavut. This region,

which covers nearly 40% of the Canadian landmass, is characterized by dispersed communities and

low population densities. Water is important to the region for drinking, food, and transportation.

Most communities are located near or on water [Bruce et al., 2000]. Many of the hydrological

responses to climate warming in Canada’s north are related to the presence and characteristics of

permafrost.

2.4.3.6.2 GCM Projections

Most climate change projections suggest that the most significant increase in annual temperature

will occur in the Canadian north (Figure X(a)). Indeed, the coupled general circulation model of the

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Canadian Centre for Climate Modelling and Analysis (CCCma) of Environment Canada predicts an

increase in mean annual air temperature ranging from 2 - 6°C over the Canadian permafrost region

under a CO2-doubling scenario [Smith and Burgess, 1999; Smith et al., 2005]. Significant warming in

this region has already been detected in some northern regions. For example, the Mackenzie region

has experienced a 2.3ºC warming in winter and a 2.4ºC warming in spring since 1895 [Bruce et al.,

2000]. In the eastern arctic, however, recent cooling and aggradation of permafrost has occurred.

Records for northern Quebec show a decrease in air temperature between 1947 and 1992 ranging

from 0.02 - 0.03ºC per year [Allard et al., 1995].

2.4.3.6.3 Permafrost

A warming climate is expected to have significant affects on the permafrost distribution in

Canada. Indeed, there is ample evidence that a warming climate has already begun to have

considerable impacts on Canada’s permafrost. An observed warming of 1°C in the western arctic has

caused the eradication of thin permafrost and an apparent northward displacement in the southern

boundary of the discontinuous permafrost zone [Kwong and Gan, 1994], and an increase in

permafrost temperatures in Yukon Territory and western Northwest Territories over the last century

[Halsey et al., 1995]. More recent observations have shown permafrost degradation in Manitoba and

Quebec in the southern margin of the permafrost region, especially where there is no peat layer

[French and Egerov, 1998; Laberge and Payette, 1995], and the Mackenzie valley, Alert, Baker

Lake and Iqaluit [Romanovsky et al., 2002; Smith et al., 2005; Tarnocai et al., 2004]. The results of

these studies generally show that (a) the observed warming of Canadian permafrost is consistent with

regional changes in air temperature since the 1970s; (b) this observed warming is spatially highly

variable, and is highly dependence on local surface conditions that influence the response of

permafrost to changes in air temperature; (c) while permafrost warming has occurred in the western

Canadian arctic since the mid-to-late 1980s, the greatest warming (0.3 – 0.6°C per decade) has

occurred in the central and northern Mackenzie valley; and (d) the warming of permafrost in the high

and eastern Canadian arctic appears to have occurred later than the western arctic, with the greatest

warming occurring since the mid-1990s. In the eastern arctic, however, recent cooling and

aggradation of permafrost has occurred. Records for northern Quebec show a decrease in air

temperature between 1947 and 1992 ranging from 0.02 - 0.03ºC per year. An analysis of ground

114

temperatures in the upper 20m for the period 1988-1993 indicates that permafrost also cooled over

this time period [Allard et al., 1995]. The effects of climate change on permafrost are considered in

greater detail in section 4.1 (Permafrost).

2.4.3.6.4 Runoff

Climate impact assessment studies on hydrology and water resources have focused on the

Mackenzie River Basin (Cohen, 1993; 1994; 1997). There have been no formal impact assessments

in the high arctic or the eastern arctic. Soulis et al., [1994] show that projected runoff either increases

or decreases, depending on the climate model utilized. Bruce et al., [2000] note that it should be

noted that the observed trend (1967-96) in most Arctic rivers has been towards increased annual

mean flow except for the Mackenzie River where the headwaters extend well into the drier Prairies.

There are approximately 25,000 lakes in the Mackenzie Delta that are expected to be

significantly affected by a warming climate, particularly by changes in river ice growth, river

discharge, ice break- up and jamming, and changes in flooding magnitude and frequency [Marsh and

Hey, 1989; Marsh and Schmidt, 1993]. Modeling attempts by Marsh and Lesack [1997] predicted

(b)

Figure X. (a) Selected Regional Temperature Trends in the Arctic Region (change in temperature over period of analysis); Environment Canada (1995). (b) Precipitation Trends in

the Arctic Region [Mekis and Hogg, 1999]. Tables taken from Bruce et al.,[2000].

(a)

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that water levels in high perched lakes would decline more rapidly between episodes of flooding, and

that a typical high-closure lake would disappear within 10 years [Hofmann et al., 1998]. These

perched lakes are particularly vulnerable to a warming climate since they show a slightly negative

water balance if not flooded by the Mackenzie River during spring peak flow [Marsh and Schmidt,

1993].

Permafrost is one of the greatest influences in the response of Arctic landscape hydrology to

climate change. A warming climate will likely affect the way rivers in permafrost environments

respond to snowmelt and rainfall events. Rivers usually exhibit a quick response to these events

where permafrost is present. In these regions, the active layer is easily saturated and most of the

water reaches streams as overland flow [Woo, 1976]. Drainage basins in the permafrost region

therefore have high runoff-to-rainfall ratios [Kane et al., 1998; Lilly et al., 1998]. Once the

precipitation event is over, however, stream flow quickly decreases because permafrost restricts

groundwater flow to the stream [Dingman, 1973]. As permafrost degrades and active layers thicken,

subsurface flow will become a more important contribution to baseflow, and streamflow will become

more uniform throughout the year [Ashmore and Church, 2001; Michel and van Everdingen, 1994;

Woo, 1976]. An unfrozen zone may develop between the base of the active layer and the permafrost

table allowing for streamflow to be sustained in winter [Hinzman and Kane, 1992]. However,

enhanced winter streamflow may result in more extensive ice-river formation and the possibility of

more serious flooding during break-up in the arctic river ice [Michel and van Everdingen, 1994]. A

deeper active layer will be associated with a greater variation in the amount of water stored in the

soil as well as an increase in the movement of subsurface water downslope [Hinzman and Kane,

1992].

2.4.3.6.5 Groundwater

Under climate warming, groundwater will play a more important role in hydrological and

landscape processes in northern regions, especially in areas currently underlain by continuous

permafrost [Michel and van Everdingen, 1994]. Frost heave in the active layer may increase due to a

greater availability of unfrozen water and this has important engineering implications. Frost blisters,

which are formed when frost heave occurs, may become more numerous in the arctic region. Icing

activity, which can present a serious road hazard, may also increase in the continuous permafrost

116

zone. Greater exchanges between surface water and groundwater may lead to a greater dissolved

solids content in rivers which may have an adverse affect on fish and other aquatic life. As

permafrost thaws, increased regional groundwater flow may promote further warming and thawing

of permafrost. Groundwater may also be discharged offshore (through the sea floor) where it may

influence near shore circulation and sea-ice cover.

2.4.3.6.6 Snow Cover

Bruce et al., [2000] report that despite greater total snowfalls, decreases in winter and early

spring snow depths have been observed in the north since 1946. These decreases were also paralleled

by decrease in spring and summer snow cover duration for most of Western Canada and the Arctic.

The effects of climate change on non-permanent snow cover are discussed in greater detail in section

2.3 (Non-Permanent (Seasonal) Snow and Ice Coverage).

2.4.3.6.7 Other Implications

Climate warming in Canada’s north and the Arctic has implications for water quality, land and

water transportation, and mining. Water quality issues may occur when drinking water and sewage

lines or storage tanks buried in or above the permafrost buckle as a result of permafrost degradation,

increasing delivery problems and contamination problems. Land transportation problems are also

related to permafrost degradation where roads became impassable because of differential settlement

as underlying permafrost thaw. Water transport is vital in the north because this region has so few

roads. However, rivers are also vulnerable to climate-change related impacts. For example,

landslides caused by permafrost degradation and low flow rates combine to cause increased siltation

problems, limiting navigation [Bruce et al., 2000]. However, not all impacts of climate warming are

detrimental to water transport. Indeed, longer ice-free periods may lengthen the shipping season on

the Mackenzie River and in the Beaufort Sea. Resource extraction activities are also an important

economic activity in the north. Climate change may affect these activities by causing permafrost

deterioration, which impacts upon road and reservoir stability, and by a declining quality of water

supply, brought about by increased siltation and turbidity [Bruce et al., 2000].

117

2.4.3.7 A Regional Perspective - British Columbia

2.4.3.7.1 Location and Climate

British Columbia contains four main climatic regions – the Pacific Coast, South B.C. Mountains,

Yukon/North B.C. Mountains, and Northwestern Forest. These regions include some of the wettest

(coastal mountains) and driest (southern interior) places in southern Canada [Bruce et al., 2000].

This diversity of climate results in a great diversity of ecosystems. BC contains more than 24,000

streams and lakes that yield water for domestic use, agriculture, industry, hydro-electric power

generation, wastewater assimilation, recreation and transport functions. Much of the human use of

water occurs in the southwest and southern interior portions of the province. Almost 60% of BC’s

population is concentrated in the Vancouver and Victoria areas [Bruce et al., 2000].

2.4.3.7.2 GCM Projections

GCM projections suggest that mean annual temperatures will increase from 1ºC in the extreme

southwest to 3ºC in northern areas by 2050 (relative to 1961-1990 average; Bruce [2000]). This

warming is not predicted to be temporally or geographically uniform. Winter warming along the

coast is predicted to be in the order of 1-2ºC by 2050, but 2-5ºC in the interior. Summer temperatures

are expected to increase by 2-3ºC over the entire province over the same time period. Slight changes

in annual precipitation are also projected to accompany changes in temperature (although a small

reduction may occur in northern coastal areas).

2.4.3.7.3 Sea Level Rise

Sea-level is expected to rise to varying degrees along the B.C. coast. Allowing for thermal

expansion of oceans, melting of land-based ice, oceanic and coastal winds, isostatic rebound and

tectonic processes, Thomson and Crawford [1997] estimate that relative sea-level will change from -

1 to 2mm/year along the south coast and from -1 to 6mm/year along the north coast [Bruce et al.,

2000]. The affects of sea level rise on Canadian coastal zones is considered in greater detail in

section 2.6 (Coastal Zones).

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2.4.3.7.4 Runoff

Climate warming in BC is expected to bring about various changes in the hydrology of the

Province. This includes (a) a retreat of glaciers in lower elevations, but an expansion of glaciers in

the north because of increased winter snowfall [Brugman et al., 1997]; (b) earlier snowmelt,

especially in southern BC, where it may occur up to one full month earlier [Coulson, 1997], (c)

increased annual, winter and spring runoff, which would likely be offset by greater

evapotranspiration associated with rising temperatures and longer growing seasons [Coulson, 1997],

(d) lower summer streamflows, (e) increased peak streamflows, and (f) water table decline [Hii,

1997].

2.4.3.7.5 Extreme Events

Flooding and coastal erosion is predicted to accompany climate change. Warmer winter

temperatures will lead to a greater proportion of total annual precipitation falling as rain, leading to

higher winter flows and more frequent flooding, especially along the coast [Bruce et al., 2000]. The

frequency of landslides may also be expected to increase as precipitation rises [Evans and Clague,

1997]. Rock avalanches and outburst floods may also become common as glacial ice melts,

debutressing mountain slopes under rising temperatures [Evans and Clague, 1997].

2.4.3.7.6 Other Implications

Climate warming in BC has implications for water quality, agriculture, municipal and rural

domestic water use, fisheries and hydro-electric power generation. Water quality issues may be

brought about by salt water intrusion into aquifers, water-borne health effects and increased water

turbidity caused by a greater frequency of landslides and surface erosion. Agriculture will also be

affected by a warming climate, particularly in terms of water sources for irrigation purposes. Ninety-

nine percent of the water used for irrigation in BC is from surface sources, many of which are fed by

snowmelt systems. The projected warmer and drier summers, in combination with an increased

demand from non-agricultural users [Zebarth et al., 1997], makes the reliability of this resource

questionable in the long term [Bruce et al., 2000]. Agricultural practices will thus need to adapt to

face this challenge. The primary concerns for municipal and rural water use are the depletion of

119

groundwater reserved and the ability of existing reservoirs to capture enough water to satisfy a

longer demand season associated with climate warming [Bruce et al., 2000]. A greater frequency of

extreme precipitation events will lead to a decrease in water infiltration and groundwater recharge, a

particular concern in regions such as the Gulf Islands, where groundwater is the primary domestic

supply [Bruce et al., 2000]. Fisheries will also be affected by changes in the hydrological regime

brought about by warming climate. Of concern to fisheries are decreased seasonal stream flows and

changing water quality (temperature and increased sediment loads), which may have severely

detrimental impacts on spawning habitat. Less clear, however, is the affect of climate warming on

ocean upwelling, productivity and temperature. It is generally accepted that a warming climate will

have only a small affect on hydro-electric power generation in BC, with the exception of

southeastern BC where decreased runoff may become a problem [Ross and Wellisch, 1997].

2.4.4 References

Abraham, J., T. Canavan, and R. Shaw, Climate Change and Climate Variability in Atlantic Canada. Volume IV of the Canada Country Study: Climate Impacts and Adaptation, Environment Canada, Atmospheric Science Division, Atlantic Region, Toronto, ON, Canada, 1997.

Allard, M., B. Wang, and J.A. Pilon, Recent cooling along the southern shore of Hudson Strait Quebec, Canada, documented from permafrost temperature measurements, Arctic and Alpine Research, 27, 157-166, 1995.

Anderson, W.L., Roberson, D. and Magnuson, J., Evidence of recent warming and El Nino related variations in ice breakup of Wisconsin Lakes, Limnology and Oceanography, 41 (5), 815-821, 1996.

Arnell, N.W., Climate change and global water resources, Global Environmental Change, 9, S31-S49, 1999.

Ashmore, P., and M. Church, The impact of climate change on rivers and river processes in Canada, pp. 58, Natural Resources Canada, Government of Canada, Ottawa, Ontario, Canada, 2001.

Assel, R., Implications of Carbon Dioxide Global Warming on Great Lakes Ice Cover, Climatic Change, 18 (4), 377-396, 1991.

Benn, D.I., and D.J.A. Evans, Glaciers and Glaciation, 734 pp., John Wiley and Sons, London, UK, 1998.

BESIS, Assessing the consequences of climate change in Alaska and he Bering Sea region, in Proceedings of a workshop at the University of Alaska Fairbanks, edited by G. Weller, and P.A. Anderson, pp. 94, Bering Sea Impacts Study, Center for Global Change and Arctic ystem Research, Fairbanks, AK, USA, 1999.

120

Bjonback, D., Symposium on the impacts of climatic change and variability on the Great Plains, in Symposium on the impacts of climatic change and variability on the Great Plains, edited by G. Wall, pp. 167-169, University of Waterloo, Waterloo, Ontario, Canada, 1991.

Boer, G.J., G. Flato, and D. Ramsden, A transient climate change simulation with greenhouse gas and aerosol forcing: projected climate to the 21st century, Atmospheric Environment Service, Canadian Centre for Climate Modelling and Analysis, Victoria, BC, Canada, 1998.

Boer, G.J., N. McFarlane, and M. Lazare, Greenhouse gas-induced climate change simulated with the CCC second-generation general circulation model, Journal of Climatology, 5, 1045-1077, 1992.

Bootsma, A., Regional Climate Sensitivities – Agriculture. Climate Change and Climate Variability in Atlantic Canada. Volume VI of the Canada Country Study: Climate Impacts and Adaptation, edited by J. Abraham, T. Canavan, and R. Shaw, Environment Canada, Bedford, NS, Canada, 1997.

Botkin, D.B., and E.A. Keller, Environmental Science: Earth as a living planet, 649 pp., John Wiley and Sons, New York, NY, USA, 1998.

Brown, R., M. Hughes, and D. Robinson, Characterizing long term variability of snow cover extent over the interior of North America, Annals of Glaciology, 21, 45-50, 1994.

Bruce, J.P., I. Burton, H. Martin, B. Mills, and L. Mortsch, Water sector: vulnerability and adaptation to climate change, pp. 144, Global Change Strategies International Inc. and Meterological Service of Canada, Ottawa, Canada, 2000.

Brugman, M.M., P. Raistrick, and A. Pietroniro, Glacier related impacts of doubling atmospheric carbon dioxide concentrations on British Columbia and Yukon. Responding to global climate change in British Columbia and Yukon, Volume I of the Canada Country Study: Climate impacts and adaptation, edited by E. Taylor, and B. Taylor, pp. 6-1 to 6-9, British Columbia Ministry of Environment, Lands and Parks and Environment Canada, Vancouver, BC, Canada, 1997.

Byrne, J., R. Barendregt, and S. D., Assessing potential climate change impacts on water supply and demand in southern Alberta, Canadian Water Resources Journal, 14 (5-15), 1989.

Canadian Climate Impacts and Adaptation Research Network, Climate Change Impacts on Atlantic Canada, Online document (http://www.dal.ca/~cciarn//workshops/4/background.html), 2004.

Changnon, S.A., changes in climate and levelsof Lake Michigan: shoreline impacts at Chicago, Bulletin of the American Meteorological Society, 70, 1092-1104, 1993.

Chao, P.T., B.F. Hobbs, and B.N. Venkatesh, How climate uncertainty should be included in Great Lakes management: Modeling workshop results, Journal of the American Water Resources Association, 35 (6), 1485-1497, 1999.

Clair, T.A., J. Ehrman, and K. Higuchi, Changes to the runoff of Canadian ecozones under a doubled CO2 atmosphere, Canadian Journal of Fisheries and Aquatic Science, 55, 2464-2477, 1998a.

Clair, T.A., B.G. Warner, R. Robarts, H. Murkin, J. Lilley, L. Mortsch, and C. Rubec, Canadian Inland Wetlands and Climate Change. Climate Change and Climate Variability in Atlantic

121

Canada. Volume VI of the Canada Country Study: Climate Impacts and Adaptation., pp. 189-218, Government of Canada, Ottawa, Canada, 1998b.

Cohen, S.J., Impacts of CO2-induced climatic change on water resources in the Great Lakes Basin, Climatic Change, 8, 135-153, 1986.

Cohen, S.J., Possible impacts of climate warming scenarios on water resources in the Saskatchewan River sub-basin, Climate Change, 19, 291-317, 1991.

Cohen, S.J., and K. Miller, North America, in Climate Change 2001: Impacts, Adaptation and Vulnerability. Contribution of Working Group II to the Third Assessment Report of the Intergovernmental Panel on Climate Change, edited by J.J. McCarthy, O.F. Canziani, N.A. Leary, D.J. Dokken, and K.S. White, pp. 735-800, Cambridge University Press, Cambridge, UK, 2001.

Cohen, S.J., K.A. Miller, A.F. Hamlet, and W. Avis, Climate change and resource management in the Columbia River Basin, Water International, 25 (2), 253-272, 2000.

Compagnucci, R., L. da Cunha, K. Hanaki, C. Howe, G. Mailu, I. Shiklomanov, E. Stakhiv, P. Döll, A. Becker, and J. Zhang, Hydrology and Water Resources, in Climate Change 2001: Impacts, Adaptation, and Vulnerability. Contribution of Working Group II to the Third Assessment Report of the Intergovernmental Panel on Climate Change (IPCC), edited by J.J. McCarthy, O.F. Canziani, N.A. Leary, D.J. Dokken, and K.S. White, pp. 195-233, Cambridge University Press, Cambridge, UK, 2001.

Coulson, H., The impacts of climate change on river and stream flow in British Columbia and Southern Yukon. Responding to global climate change in British Columbia and Yukon, Volume I of the Canada Country Study: Climate impacts and adaptation, edited by E. Taylor, and B. Taylor, pp. 5-1 to 5-11, British Columbia Ministry of Environment, Lands and Parks and Environment Canada, Vancouver, BC, Canada, 1997.

Croley II, T., Laurentian Great Lakes double-CO2 climate change hydrology impacts, Climatic Change, 17, 27-47, 1990.

Crowe, A., The application of a coupled water-balance-salinity model to evaluate the sensitivity of a lake dominated by groundwater to climatic variability, Journal of Hydrology, 141 (1-4), 33-73, 1993.

Cruise, J.F., A.S. Limaye, and N. Al-Abed,, Assessment of impacts of climate change on water quality in the Southeastern United States, Journal of the American Water Resources Association,, 35 (6), 1539-1550, 1999.

Danard, M., and T. Murty, Some recent trends in precipitation in Western Canada and their possible link to rising carbon dioxide, Atmosphere-Ocean, 24, 52-72, 1988.

Demuth, M.N., and A. Pietronomo, The impact of climate change on the glaciers of the Canadian Rocky Mountain eastern slopes and implications for water resource-related adaptation in the Canadian prairies. Phase I - Headwaters of the North Saskatchewan River Basin, pp. 96, Climate Change Action Fund - Prairie Adaptation Research Collaborative, Ottawa, Canada, 2003.

122

Dingman, S.L., Effects of permafrost on stream flow characteristics in the discontinuous permafrost zone of central Alaska, in Proceedings of the 2nd International Conference on Permafrost, North American contribution,, pp. 447-453, National Academy of Sciences, Washington, D.C., USA, 1973.

Draper, D., Our Environment: A Canadian Perspective, 540 pp., Nelson, Scarborough, Canada, 2002.

Eheart, J.W., A.J. Wildermuth, and E.E. Herricks, The effects of climate change and irrigation on criterion low streamflows used for determining total maximum daily loads, Journal of the American Water Resources Association, 35 (6), 1365-1372, 1999.

Environment Canada, The state of Canada's climate: Monitoring variability and change, pp. 52, Minister of Public Works and Government Services Canada, Ottawa, ON, Canada, 1995.

Environment Canada, Threats to water availability in Canada, NWRI Scientific Assessment Report Series No. 3 and ACSD Science Assessment Series No. 1, pp. 128, National Water Research Institute, Burlington, ON, Canada, 2004a.

Environment Canada, Water – Vulnerable to Climate Change, (http://www.ec.gc.ca/water/en/info /pubs/FS/e_FSA9.htm), Last accessed March 2005, 2004b.

Evans, J.C., and E.E. Prepas, Potential effects of climate change on ion chemistry and phytoplankton communities in prairie saline lakes, Limnology and Oceanography, 41 (5), 1063-1076, 1996.

Evans, S., and J. Clague, The impact of climate change on catastrophic geomorphic processes in the mountains of British Columbia. Responding to global climate change in British Columbia and Yukon, Volume I of the Canada Country Study: Climate impacts and adaptation, edited by E. Taylor, and B. Taylor, pp. 7-1 to 7-16, British Columbia Ministry of Environment, Lands and Parks and Environment Canada, Vancouver, BC, Canada, 1997.

Felzer, B., and P. Heard, Precipitation differences amongst GCMs used for the U.S. National Assessment, Journal of the American Water Resources Association, 35 (6), 1327-1340, 1999.

Frederick, K.D.a.P.H.G., Water and Global Climate Change: Potential Impacts on U.S. Water Resources, 55 pp., The Pew Center on Global Climate Change, Arlington, VA, USA, 1999.

Freeman, C., M. Lock, and B. Reynolds, Fluxes of CO2, CH4, and N2O from a Welsh peatland following simulation of water table draw-down; potential feedback to climate change, Biogeochemistry, 19 (1), 51-60, 1993.

French, H.M., and I.E. Egerov, 20th century variations in the southern limit of permafrost near Thompson, northern Manitoba, in International Conference on Permafrost, pp. 297-304, Université Laval, Québec, Canada, Yellowknife, Canada, 1998.

Fyfe, J.C.a.G.M.F., Enhanced climate change and its detection over the Rocky Mountains, Journal of Climate, 12 (1), 230-243, 1999.

Gleick, P.H., and E.L. Chalecki, The impacts of climatic changes for water resources of the Colorado and Sacramento-San Joaquin river basins, Journal of the American Water Resources Association, 35 (6), 1429-1442., 1999.

123

Government of Canada, The Federal policy on wetland conservation, Supply and Services Canada, Ottawa, Canada, 1991.

Grimm, N., A. Chacon, C. Dahm, S. Hostetler, O. Lind, P. Starkweather, and W. Wurtsburgh, Sensitivity of aquatic ecosystems to climate and anthropogenic changes: The basin and range, American Southwest, and Mexico, Hydrological Processes, 11, 1023-1042, 1997.

Haas, J., and T.J. Marta, Assessing the effect of climate change on the operation of a water-supply reservoir, in Unpublished manuscript,, pp. 6, Inland Waters Directorate, Environment Canada, Ottawa, Ontario, Canada, 1988.

Halsey, L.A., D.H. Vitt, and S.C. Zoltai, Disequilibrium response of permafrost in boreal continental western Canada to climate change, Climate Change, 30, 57-73, 1995.

Hamlet, A.F., and D.P. Lettenmaier, Effects of climate change on hydrology and water resources in the Columbia River Basin, Journal of the American Water Resources Association, 35 (6), 1597-1624., 1999.

Hammer, U.T., The effects of climate change on the salinity, water levels and biota of Canadian Prairie saline lakes, International Association of Theoretical and Applied Limnology, 24, 321-326, 1990.

Hansen, J., A. Lacis, D. Rind, G. Russell, P. Ston, I. Fung, R. Ruedy, and J. Lerner, Climate sensitivity: Analysis of feedback mechanism, in Climate Processes and Climate Sensitivity, edited by J.E. Hansen, and T. Takahashi pp. 103-163, American Geophysical Union, Washington, D.C, USA, 1984.

Hansen, J., G. Russell, D. Rind, P. Stone, A. Lacis, S. Lebedeff, R. Rudey, and L. Travis, Efficient three-dimensional global models for climate studies: Models I and II, Monthly Weather Review, 111 (609-662), 1983.

Hanson, H., C. Hanson, and B. Yoo, Recent Great Lakes ice trends, Bulletin of the American Meteorological Society, 73 (5), 577-584, 1992.

Hare, K.H., R.B.B. Dickison, and S. Ismail, Climatic Variation over the Saint John Basin : An Examination of Regional Behaviour. Climate Change Digest Series, CCD 97-02, Environment Canada, Downsview, ON, Canada, 1997.

Hartmann, H.C., Climate change impacts on Laurentian Great Lakes levels, Climatic Change, 17, 49-67, 1990.

Hauer, F.R., J.S. Baron, D.H. Campbell, K.D. Fausch, S.W. Hostetler, G.H. Leavesley, P.R. Leavitt, D.M. McKnight, and J.A. Stanford, Assessment of climate change and freshwater ecosystems of the Rocky Mountains, USA and Canada, Hydrological Processes, 11, 903-924, 1997.

Hengeveld, H.G., Projections for Canada's climate future: a discussion of recent simulations with the Canadian Global Climate Model, Climate Change Digest Series, CCD 00-01, pp. 27, Environment Canada, Ottawa, ON, Canada, 2000.

Hennessy, R.J., J.M. Gregory, and J.F.B. Mitchell, Changes in daily precipitation under enhanced greenhouse conditions, Climate Dynamics, 13, 667-680, 1997.

124

Herbert, P.D.N., Wetlands, in Canada's Aquatic Environments (Online Article), University of Guelph, Guelph, Canada, 2002.

Herrington, R., B. Johnson, and F. Hunter, Responding to Global Climate Change in the Prairies. The Canada Country Study: Climate Impacts and Adaptation, pp. 75, Environment Canada, Toronto, ON, Canada, 1997.

Hii, B., The impacts of climate change on the Abbotsford aquifer. Responding to global climate change in British Columbia and Yukon, Volume I of the Canada Country Study: Climate impacts and adaptation, edited by E. Taylor, and B. Taylor, pp. 18-1 to 18-6, British Columbia Ministry of Environment, Lands and Parks and Environment Canada, Vancouver, BC, Canada, 1997.

Hinzman, L.D., and D.L. Kane, Potential response of an arctic watershed during a period of global warming, Journal of Geophysical Research, 97 (D3), 2811-2820, 1992.

Hofmann, N., L. Mortsch, S. Donner, K. Duncan, R. Kreutzwiser, S. Kulshrestha, A. Pigott, S. Schellenberg, B. Schertzer, and M. Slivitzky, Climate Change and Variability: Impacts on Canadian Water. Volume VIII of the Canada Country Study, pp. 1-120, Government of Canada, Ottawa, Canada, 1998.

Howitt, R., and m. M'Marette, The value of groundwater in adapting to drought: lessons from 1976-1977, in Coping with water scarcity: The role of groundwater, edited by J. De Vries, pp. 45-54, California Water Resources Center, University of California, Riverside, CA, USA, 1990.

Hurd, B.H., J.M. Callaway, J.B. Smith, and P. Kirshen,, Economic effects of climate change on U.S. water resources, in The Economic Impacts of Climate Change on the U.S. Economy, edited by R. Mendelsohn, and J.E. Neumann, pp. 133-177, Cambridge University Press, Cambridge, United Kingdom and New York, NY, USA, 1998.

Jansson, A., C. Folke, J. Rockström, and L. Gordon, Linking Freshwater Flows and Ecosystem Services Appropriated by People: The Case of the Baltic Sea Drainage Basin, Ecosystems, 2, 351-356, 1999.

Jury, W., and J. Gruber, Stochastic analysis of the indluence of soil and climatic variability on the estimate of pesticide groundwater pollution potential, Water Resources Research, 25 (12), 2465-2474, 1989.

Justic, D., N.N. Rabalais, and R.E. Turner, Effects of climate change on hypoxia in coastal waters: a doubled CO2 scenario for the northern Gulf of Mexico, Limnology and Oceanography, 41 (5), 992-1003, 1996.

Kane, D.L., D.J. Soden, L.D. Hinzman, and R.E. Gieck, Rainfall runoff of a nested watershed in the Alaskan arctic, in Proceedings of the 7th International Conference on Permafrost, North American contribution,, pp. 539-543, Université Laval, Québec, Canada, Yellowknife, NWT, Canada, 1998.

Kukkonen, I., V. Cermak, and J. Safanda, Subsurface temperature-depth profiles, anomalies due to climatic ground surface temperature changes or groundwater flow effects, Global and Planetary Change, 9 (3/4), 221-232, 1994.

125

Kwong, Y.T., and T.Y. Gan, Northward migration of permafrost along the Mackenzie Highway and climate warming, Climate Change, 26, 399-419, 1994.

Laberge, M.-J., and S. Payette, Long-term monitoring of permafrost change in a palsa peatland in northern Quebec, Canada, Arctic and Alpine Research, 27, 167-171, 1995.

Leavesley, G.H., Modeling the effects of climate change on water resources - a review, Climatic Change, 28, 159-177, 1994.

Lee, D.H., F.H. Quinn, D. Sparks, and J.C. Rassam, Modification of Great Lakes regulation plans for simulation of maximum Lake Ontario outflows, Journal of Great Lakes Research, 20 (3), 569-582, 1994.

Leith, R.M.M., and P.H. Whitfield, Evidence of climate change effects on the hydrology of streams in south-central B.C., Canadian Water Resources Journal, 23 (3), 219-230, 1998.

Lemmen, D.S., F.J. Warren, E. Barrow, R. Schwartz, J. Andrey, B. Mills, and D. Riedel, Climate change impacts and adaptation: A Canadian perspective, pp. 174, Climate Change Impacts and Adaptation Directorate, Natural Resources Canada, Ottawa, Ontario, Canada, 2004.

Lettenmaier, D.P., A.W. Wood, R.N. Palmer, E.F. Wood, and E.Z. Stakhiv,, Water resources implications of global warming: a U.S. regional perspective, Climatic Change, 43 (3), 537-579, 1999.

Leung, L.R., and M.S. Wigmosta, Potential climate change impacts on mountain watersheds in the Pacific Northwest, Journal of the American Water Resources Association, 35 (6), 1463-1472, 1999.

Lilly, E.K., D.L. Kane, L.D. Hinzman, and R.E. Gieck, Annual water balance for three nested watersheds on the north slope of Alaska, in Proceedings of the 7th International Conference on Permafrost, North American contribution,, pp. 669-674, Université Laval, Québec, Canada, Yellowknife, NWT, Canada, 1998.

Loukas, A., and M.C. Quick, The effect of climate change on floods in British Columbia, Nordic Hydrology, 30, 231-256, 1999.

MacDonald, M.E., A.E. Hershey, and M.C. Miller, Global warming impacts on lake trout in arctic lakes, Limnology and Oceanography, 41 (5), 1102-1108, 1996.

Magnuson, J., R. Assel, C. Bowser, P. Dillon, J. Eaton, H. Evans, E. Fee, R. Hall, L. Mortsch, F. Quinn, D. Schindler, and K. Webster, Potential effects of climate changes on aquatic systems: Laurentian Great Lakes and Precambrian Shield region, Hydrological Processes, 11, 825-872, 1997.

Manabe, S., and R.T. Wetherald, Large-scale changes in soil wetness induced by an increase in carbon dioxide, Journal of Atmospheric Sience, 44, 1211-1235, 1987.

Marsh, P., and M. Hey, The flooding hydrology of the Mackenzie Delta near Inuvik, NWT, Canada, Arctic, 42, 41-49, 1989.

Marsh, P., and L.F.W. Lesack, The hydrologic response of perched lakes in the Mackenzie delta: potential responses to climate change, Limnology and Oceanography, 41 (5), 849-856, 1996.

126

Marsh, P., and L.F.W. Lesack, Climate change and the hydrologic regime of lakes in the Mackenzie Delta, Limnology and Oceanography, 41, 849-856, 1997.

Marsh, P., and T. Schmidt, Influence of a Beaufort sea storm surge on channel levels in the Mackenzie Delta, Arctic, 46, 35-41, 1993.

Maxwell, B., Responding to global climate change in Canada's arctic. Volume II of the Canada Country study:Climate impacts and adaptation, pp. 82, Environemntal Adaptation Research Group, Atmospheric Environment Service, Environment Canada, 1997.

McCabe, G.J., and D.M. Wolock, General circulation model simulations of future snowpack in the western United States, Journal of the American Water Resources Association, 35 (6), 1473-1484, 1999.

McFarlane, N.A., G.J. Boer, J.-P. Blanchet, and M. Lazare, The Canadian Climate Centre General Circulation Model and its equilibrium climate, Journal of Climate, 5, 1013–1044, 1992.

McGinn, S.M., and A. Shepherd, Impact of climate change scenarios on the agroclimate of the Canadian prairies, Canadian Journal of Soil Science, 83, 623-630, 2003.

McGuffie, K., A. Henderson-Sellers, N. Holbrook, Z. Kothavala, O. Balachova, and J. Hoekstra, Assessing simulations of daily temperature and precipitation variability with global climate models for present and enhanced greenhouse climates, International Journal of Climatology, 19, 1-26, 1999.

McLaren, R.G., and E.A. Sudicky, The Impact of Climate Change on Groundwater, in The impact of climate change on water in the Grand River Basin, Ontario, edited by M. Sanderson, pp. 53-67, University of Waterloo, Waterloo, Ontario, Canada, 1993.

Meisner, J., The role of groundwater in the effect of climatic warming on stream habitat of Brook Trout, in Climate change: Implications for water and ecological resources, edited by G. Wall, and M. Sanderson, pp. 209-215, University of Waterloo, Waterloo, Ontario, Canada, 1990.

Meisner, J., J. Goodier, H. Regier, B. Shuter, and W. Christie, An assessment of the effects of climate warming in Great Lakes Basin fishes, Journal of Great Lakes Research, 37, 340-352, 1987.

Mekis, É., and W.D. Hogg, Rehabilitation and analysis of Canadian daily precipitation time series, Atmosphere-Ocean, 37 (1), 53-85, 1999.

Melack, J.M., J. Dozier, C.R. Goldman, D. Greenland, A.M. Milner, and R.J. Naiman, Effects of climate change on inland waters of the Pacific coastal mountains and western Great basin of North America, Hydrological Processes, 11, 971-992, 1997.

Mercier, G., Climate Change and Variability: Energy Sector. Canada Country Study: Impacts and Adaptations, Sectoral Volume., edited by G. Koshida, and W. Avis, pp. 17-1 to 17-5, British Columbia Ministry of Environment, Lands and Parks and Environment Canada, Vancouver, BC, Canada, 1998.

Michel, F.A., and R.O. van Everdingen, Changes in hydrogeologic regimes in permafrost regions due to climate change, Permafrost and Periglacial Processes, 5, 191-195, 1994.

127

Mimikou, M., Y. Kouvopoulos, G. Cadavias, and N. Vayianos, Regional hydroecological effects of climate change, Journal of Hydrology, 123 (1/2), 119-146, 1991.

Moore, M.V., J.R. Mather, P.S. Murdoch, R.W. Howarth, C.L. Folt, C.Y. Chen, H.F. Hemond, P.A. Flebbe, and C.T. Driscoll, Potential effects of climate change on freshwater ecosystems of the New England/Mid-Atlantic region, Hydrologic Processes, 11, 925-947, 1997.

Mortsch, L., and F. Quinn, Climate change scenarios for Great Lakes Basin ecosystem studies, Limnology and Oceanography, 41 (5), 803-911, 1996.

Mulholland, P.J., G.R. Best, C.C. Coutant, G.M. Hornberger, J.L. Meyer, P.J. Robinson, J.R. Stenberg, R.E. Turner, F. Verra-Herrara, and R.G. Wetzel, Effects of climate change on freshwater ecosystems of the south-eastern United States and Gulf of Mexico, Hydrological Processes, 11, 131-152, 1997.

Naiman, R.J., J.J. Magnuson, D.M. McKnight, and J.A. Stanford, The freshwater imperative, Island Press, Washington, D.C., USA, 1995.

National Wetlands Working Group, The Canadian wetland classification system, 68 pp., Wetlands Research Centre, University of Waterloo, Waterloo, Canada, 1997.

Ng, H., and J. Marsalek, Sensitivity of streamflow simulation to changes in climatic inputs, Nordic Hydrology, 23 (4), 257-272, 1992.

Nisbet, E.G., Some northern sources of atmospheric methane: production, history and future implications, Canadian Journal of Earh Sciences, 26, 1603-1611, 1989.

Nkendrim, L., and H. Purves, Comparison of recorded temperature and precipitation in the Oldham Basin with scenarios projected in general circulation models, Canadian Water Resources Journal, 19 (2), 157-164, 1994.

Noda, A., and T. Tokioka, The effect of the doubling of the CO2 concentration on convective and non-convective precipitation in a general circulation model coupled with a simple mixed later ocean, Journal of the Meteorological Society of Japan, 67, 1055-1067, 1989.

Oerlemans, J., Glaciers and climate change, 148 pp., A.A. Balkema, Exton, PA, USA, 2001.

Poiani, K., and W. Johnson, Global warming and prairie wetlands: Potential consequences for waterfowl habitat, BioScience, 41, 611-618, 1991.

Poiani, K., and W. Johnson, Potential effects of climate change on a semi-permanent Prairie wetland, Climatic Change, 24, 213-232, 1993a.

Poiani, K., and W. Johnson, A spatial simulation model of hydrology and vegetation dynamics in semi-permanent prairie wetlands: Simulations of long-term dynamics, Ecological Applications, 3, 279-293, 1993b.

Porter, K.G., P.A. Saunders, K.A. Haberyan, A.E. Macubbin, T.R. Jacobsen, and R.E. Hodson, Annual cycle of autotrophic and heterotrophic production in a small, monomictic Piedmont lake (Lake Oglethorpe): analog for the effects of climatic warming on dimictic lakes, Limnology and Oceanography, 41 (5), 1041-1051, 1996.

Postel, S., and S.R. Carpenter, Freshwater ecosystem services, in Nature's services, edited by G. Daily, pp. 195-214, Island Press, Washington, D.C., USA, 1997.

128

Quinn, F.H., and B. M. Lofgren,, The influence of potential greenhouse warming on Great Lakes hydrology, water levels, and water management, in Proceedings, 15th Conference on Hydrology, American Meteorological Society Annual Meeting, pp. 271-274, American Meteorological Society, Long Beach, CA, USA, 2000.

Ripley, E., Climate change and the hydrological cycle, in Canadian Aquatic Resources, edited by M. Healey, and R. Wallace, Department of Fisheries and Oceans, Ottawa, ON, Canada, 1987.

Romanovsky, V., S.L. Smith, K. Yoshikawa, and J. Brown, Permafrost temperature records: Indicators of climate change, EOS, Transactions of the American Geophysical Union, 83 (50), 593-594, 2002.

Ross, L., and M. Wellisch, Implications of future climate change on energy production in British Columbia and Yukon. Responding to global climate change in British Columbia and Yukon, Volume I of the Canada Country Study: Climate impacts and adaptation, edited by B. Taylor, pp. 17-1 to 17-5, British Columbia Ministry of Environment, Lands and Parks and Environment Canada, Vancouver, BC, Canada, 1997.

Rouse, W., M. Douglas, R. Hecky, A. Hershey, G. Kling, L. Lesack, P. Marsh, P. McDonald, M. Nicholson, N. Roulet, and J. Smol, Effects of climate change on the fresh waters of Arctic and subarctic North America, Hydrological Processes, 11, 873-902, 1997.

Sanderson, M., and J.V. Smith, Climate change and water balance of the Grand River basin, in Proceedings of the 43rd Conference, Canadian Water Resources Association, pp. 234-261, Penticton, BC, Canada, 1990.

Sanderson, M., and J.V. Smith, The present 2xCO2 climate and water balance in the Basin, in The impact of climate change on water in the Grand River Basin, Ontario, edited by M. Sanderson, pp. 3-24, University of Waterloo, Waterloo, Ontario, Canada, 1993.

Sandstrom, K., Modeling the effects of rainfall variability on groundwater recharge in semi-arid Tanzania, Nordic Hydrology, 26 (4/5), 313-330, 1995.

Schertzer, W.M., W.R. Rouse, D.C.L. Lam, D. Bonin, and L.D. Mortsch, Climate variability and change - Lakes and reservoirs, in Threats to water availability in Canada, NWRI Scientific Assessment Report Series No. 3 and ACSD Science Assessment Series No. 1, pp. 91-99, National Water Research Institute, Meteorological Service of Canada, Environment Canada, Burlington, ON, Canada, 2004.

Schindler, D., Bayley, S., Parker, B., Beaty, K., Cruikshank, D., Fee, E., Schindler, E.,, and M. & Stainton, The effects of climatic warming on the properties of boreal lakes and streams at the Experimental Lakes Area, northwestern Ontario, Limnology and Oceanography, 41 (5), 1004-1017, 1996.

Schindler, D.W., S. Bayley, B. Parker, K.G. Beaty, D.R. Cruikshank, E. Fee, E. Schindler, and M.P. Stainton, The effects of climate warming on the properties of boreal lakes and streams in the Experimental Lakes Area, Northwestern Ontario, Limnology and Oceanography, 41 (5), 1004-1017, 1996.

Schlesinger, M., and Z. Zhao, Seasonal Climate changes induced by doubled CO2 as simulated by the OSU Atmospheric GCM/Mixed-Layer Ocean Model, Oregon State University Climate research Institute, Corvallis, OR, USA, 1988.

129

Shaw, R.W., Summary. Climate Change and Climate Variability in Atlantic Canada. Volume VI of the Canada Country Study: Climate Impacts and Adaptation., edited by J. Abraham, T. Canavan, and R. Shaw, Environment Canada, Bedford, NS, Canada, 1997.

Shepherd, A., and S.M. McGinn, Assessment of climate change on the Canadian prairies from downscaled GCM data, Atmosphere-Ocean, 41 (4), 301-316, 2003.

Singh, B., The implications of climate change for natural resources in Quebec. Climate Change Digest Series CCD 88-08, pp. 11, Environment Canada, Ottawa, ON, Canada, 1988.

Smith, S.L., and M.M. Burgess, Mapping the sensitivity of Canadian permafrost to climate warming, in Interactions Between the Cryosphere, Climate and Greenhouse Gases (Proceedings of IUGG 99 Symposium HS2); IAHS Publication no. 256, pp. 71-80, Birmingham, UK, 1999.

Smith, S.L., M.M. Burgess, D. Riseborough, and F.M. Nixon, Recent trends from Canadian permafrost thermal monitoring network sites, Permafrost and Periglacial Processes, 16, 1-13, 2005.

Soulis, E., S. Solomon, N. Lee, and N. Kouwen, Changes to the distribution of monthly and annual runoff in the Mackenzie Basin using a modified square grid approach, in Mackenzie Basin Impact Study, Interim Report #2, edited by J.E. Cohen, pp. 197-209, Proceedings of 6th Biennial AES/DIAND Meeting on Northern Climate and Mid Study Workshop of MBIS, Environment Canada, Downsview, Ontario, Canada, 1994.

Soveri, J., and T. Ahlberg, Multiannual variations of groundwater level in Finland during the years 1962-1989, in Conference on Climate and Water, pp. 501-510, Valtion Painatuskeskus, Helsinki, Finland, 1989.

Strzepek, K.M., D.C. Major, C. Rosenzweig, A. Iglesias, D.N. Yates, A. Holt, and D. Hillel, New methods of modeling water availability for agriculture under climate change: the U.S. cornbelt, Journal of the American Water Resources Association, 35 (6), 1639-1656, 1999.

Tarnocai, C., F.M. Nixon, and L. Kutny, Circumpolar-Active-Layer-Monitoring (CALM) sites in the Mackenzie valley, Northwestern Canada, Permafrost and Periglacial Processes, 15, 141-153, 2004.

Taylor, E., and B. Taylor, Responding to global climate change in British Columbia and Yukon, Volume I of the Canada Country Study: Climate impacts and adaptation, pp. 1-1 to A-38, Environemntal Adaptation Research Group, Atmospheric Environment Service, Environment Canada, Toronto, ON, Canada, 1997.

Thompson, R.S., S.W. Hostetler, P.J. Bartlein, and K.H. Anderson, A Strategy for Assessing Potential Future Change in Climate, Hydrology, and Vegetation in the Western United States, pp. 20, Circular 1153, U.S. Geological Survey, Washington, DC, USA, 1998.

Thomson, R., Effect of climate variability and change in groundwater in Europe, Nordic Hydrology, 21 (3), 185-194, 1990.

Thomson, R., and W. Crawford, Processes affecting sea-level change along the coasts of British Columbia and Yukon. Responding to global climate change in British Columbia and Yukon, Volume I of the Canada Country Study: Climate impacts and adaptation, edited by E. Taylor,

130

and B. Taylor, pp. 1-19, Environemntal Adaptation Research Group, Atmospheric Environment Service, Environment Canada, Bedford, NS, Canada, 1997.

USEPA, National Air Quality and Emission Trends Report, 1997. EPA Trends 1998, Office of Air Quality Planning and Standards, U.S. Environmental Protection Agency, Washington, DC, USA, 1998.

van der Kamp, G., and P. Marsh, Climate variability and change - Wetlands, in Threats to water availability in Canada, NWRI Scientific Assessment Report Series No. 3 and ACSD Science Assessment Series No. 1, pp. 101-106, National Water Research Institute, Meteorological Service of Canada, Environment Canada, Burlington, ON, Canada, 2004.

van der Valk, A.G., and C.B. Davis, Primary production of prairie glacial marshes, in Freshwater wetlands: Ecological processes and management potential, edited by R.E. Good, D.F. Whigham, and R.L. Simpson, pp. 21-37, Academic Press, New York, NY, USA, 1978.

Webster, K., A. Newell, L. Baker, and P. Brezonik, Climatically induced rapid acidification of a softwater seepage lake, Nature, 347 (6291), 374, 1990.

Whitfield, P.H., and A.J. Cannon, Recent Variations in Climate and Hydrology in Canada, Canadian Water Resources Journal, 25 (1), 2000.

Wilby, R.L., and M.D. Dettinger, Streamflow changes in the Sierra Nevada, California, simulated using a statistically downscaled general circulation model scenario of climate change, in Linking Climate Change to Land Surface Change, edited by S. McLaren, and D. Kniveton, pp. 276, J. Kluwer Academic Publishers, Dordrecht, The Netherlands, 2000.

Wilby, R.L., L.E. Hay, and G.H. Leavesley, A comparison of downscaled and raw GCM output: implications for climate change scenarios in the San Juan River basin, Colorado, Journal of Hydrology, 225 (1-2), 67-91, 1999.

Wilkinson, W., and D. Cooper, The response of idealized aquifer/river systems to climate change, Hydrological Sciences Journal, 38 (5), 379-390, 1993.

Williams, G., Predicting the Date of Lake Ice Breakup, Water Resources Research, 7 (2), 323-333, 1971.

Williams, M.W., M. Losleban, N. Caine, and D. Greenland, Changes in climate and hydrochemical responses in a high-elevation catchment in the Rocky Mountains, U.S.A., Limnology and Oceanography,, 41, 936-946, 1996.

Winter, T.A., The vulnerability of wetlands to climate change: a hydrologic landscape perspective, Journal of the American Water Resources Association, 36, 305-311, 2000.

Wittrock, V., and E. Wheaton, Saskatchewan water resources and global warming in Saskatchewan in a warmer world, pp. 73-92, Saskatchewan Research Council, Saskatoon, Saskatchewan, Canada, 1992.

Wolock, D.M., and G.J. McCabe, Estimates of runoff using water-balance and atmospheric general circulation models, Journal of the American Water Resources Association, 35 (6), 1341-1442, 2000.

131

Woo, M., Impacts of climate variability and change on Canadian wetlands, Canadian Water Resources Journal, 17, 63-69, 1992.

Woo, M.-K., Hydrology of a small Canadian high arctic basin during the snowmelt period, Catena, 3, 155-168, 1976.

Woodhouse, C.A., and J.T. Overpeck, 2000 years of drought variability in the central United States, American Meteorological Society Bulletin, 79, 2693-2714, 1998.

Yulianti, J.S., and D.H. Burn, Investigating links between climatic warming and low streamflow in the Prairies region of Canada, Canadian Water Resources Journal, 23 (1), 45-60, 1999.

Zaltsberg, E., Potential changes in mean annual runoff from a small watershed in Manitoba due to possible climatic changes, Canadian Water Resources Journal, 15 (4), 333-334, 1990.

Zebarth, B., J. Caprio, K. Broersma, P. Mills, and S. Smith, Effect of climate change on agriculture in British Columbia and Yukon. Responding to global climate change in British Columbia and Yukon, Volume I of the Canada Country Study: Climate impacts and adaptation, edited by E. Taylor, and B. Taylor, pp. 15-1 to 15-12, British Columbia Ministry of Environment, Lands and Parks and Environment Canada, Vancouver, BC, Canada, 1997.

Zhang, X., K.D. Harvey, W.D. Hogg, and T.R. Yuzuk, Trends in Canadian Streamflow, pp. 38, Meteorological Service of Canada, Environment Canada, Ottawa, Canada, 1999.

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2.5 Coastal Zones

2.5.1 The importance of Coastal Zone Ecosystems

With just over 200,000 km, Canada has the longest coastline in the world [CCRS, 2000]. The

coastline is also unique, because it not only borders the Atlantic, Arctic and Pacific oceans, but the

Great Lakes as well. The large geographic extent of Canada’s coastline means it is exposed to a wide

variety of climatic and oceanographic conditions. This has resulted in the development of a multitude

of coastal landforms including frozen tundra, fjords, rock shores, bluffs, sandy beaches, dunes, tidal

flats, salt marshes, estuaries and deltas. Canadians also have a strong link with their coastlines. A

large portion of Canada’s population lives near the coast, and a significant portion of Canada’s gross

national product (GNP) is generated by activities associated with the coast or with use of our coastal

resources. Canada’s coastal regions also provide important habitat for many different types of

wildlife including both terrestrial and aquatic plants, birds, fish and other animals [Parlee, 2004].

2.5.2 Predicted changes in Canada’s coastlines under a changing climate

For the range of scenarios developed in the IPCC Special Report on Emission Scenarios [IPCC,

2001], models project that climate change will result in the average global surface temperature

increasing between 1.4°C to 5.8°C by 2100 relative to 1990. As the earth’s atmosphere and oceans

become warmer, sea levels are expected to rise. This expected rise in sea level is mainly predicted to

be the result of the thermal expansion of ocean waters, although the melting of glaciers and changes

in volume of the polar ice sheets may also play significant (but secondary) roles [Hengeveld, 2000].

The CGCM1 projections of sea level change suggest (a) a rise in average global sea level of about

5cm during the twentieth century and an additional rise of 40cm by 2080, and (b) considerable

regional differences in sea level rise that are smaller in the Arctic (10cm by 2090), but more severe

in the northwest Atlantic east of the Maritime Provinces (40cm), and the eastern Pacific off the coast

of British Columbia (65cm) [Figure X, Hengeveld, 2000]. It is thought that this sea level rise, in

combination with more extreme weather and a loss of sea ice, will contribute towards more erosion

and flooding along vulnerable parts of Canada’s coastline, such as arctic shorelines [Natural

Resources Canada, 2002].

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Figure X. (a) Sea level rise due to thermal expansion of sea water as projected by CGCM1 and three other model experiments. Rates of change for three of the models are very similar, while the fourth is lower; (b) Because of different rates of ocean warming and local changes in atmospheric pressure patterns, the rates of change in sea levels due to thermal expansion vary significantly from region to

region [after Hengeveld, 2000].

(a)

(b)

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2.5.3 Coastal Zone Impact Issues in Canada: A Regional Perspective

2.5.3.1 General Impacts

Sea level rise is expected to have considerable impacts on the Canada’s coastlines. These impacts

include those brought about by (i) changing oceanographic processes (e.g. potential changes in ocean

temperature, salinity, density and circulation patterns), (ii) changing sea levels (e.g. storm surge,

erosion and flooding), (iii) changing weather conditions (e.g. changing wind patterns, coastal

circulation, storm activities), (iv) changing sea ice conditions (e.g. extent, thickness, location and

season), and (v) changing human use [C-CIARN, 2005]. However, the extent of these impacts is not

expected to be geographically uniform.

2.5.3.2 Impacts on Arctic Coastlines

The potential impacts of sea level rise in the Arctic include [after C-CIARN, 2005]:

• Changing Oceanographic Processes o Impact of potential changes in ocean temperature, salinity, density and circulation

patterns;

o Impact of changes in oceanographic processes on the distribution and productivity of polynyas;

o Impact of changes in oceanographic processes on marine life and traditional lifestyles (e.g. traditional knowledge, hunting, country food).

• Changing Sea Levels (including storm surge, erosion and flooding) o Impact of sea-level rise on tidal spectrum, wave climate, coastal circulation, sediment

redistribution and other physical processes;

o Impact of sea-level rise on coastal stability, particularly along vulnerable low-lying coasts, soft sediment shores, ice-rich coasts and tidewater glacier shores;

o Impact of sea-level rise on coastal ecosystems;

o Impact of sea-level rise on coastal communities and existing coastal infrastructure;

o Impact of sea-level rise on archaeological and cultural resources;

o Impact of sea-level rise on traditional lifestyles and human activities such as fisheries, hunting and trapping, transportation or tourism and recreation;

o Impact of sea-level rise on human health and safety, emergency preparedness, property loss, insurance or construction, maintenance and repair of coastal

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infrastructure;

o Impact of sea-level rise on shoreline development, real estate and legal issues;

o Impact of sea-level rise on the quantum of land allotment to land claim groups, such as the Nunavut Land Claims agreement;

o Impact of sea-level rise on property ownership, legal boundaries and jurisdictional issues.

• Changing Weather Conditions o Impact of changing wind regime on wave climate, coastal circulation, sediment

redistribution and other physical processes;

o Impact of changes wind regime on the distribution and productivity of polynyas;

o Impact of changing wind regime and storm activity on wave climate, coastal circulation, sediment redistribution and other physical processes;

o Impact of changing storm activity on storm surge events;

o Impact of changing storm activity on coastal stability, particularly along vulnerable lowlying coasts, soft sediment shores, ice-rich coasts and tidewater glacier shores;

o Impact of changing storm activity on coastal ecosystems;

o Impact of changing storm activity on coastal communities and coastal infrastructure;

Figure X. Vulnerability of Canada’s Arctic to seal level change [after Shaw et al., 1998].

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o Impact of changing storm activity on traditional lifestyles and human activities such as fisheries, hunting and trapping, transportation or tourism and recreation;

o Impact of changing storm activity on human health and safety, emergency preparedness, search and rescue, property loss, insurance or construction, maintenance and repair of shore protection structures.

• Changing Sea Ice Conditions (including extent, thickness, location and season) o Impact of changes in sea ice on waves and currents (e.g. more open water means

larger fetch and more significant wave development);

o Impact of changes in sea ice on coastal circulation, sediment redistribution and other physical processes;

o Impact of changes in sea ice on coastal stability (e.g. less shore protection from ice, increased ice pile-up and ice ride-up);

o Impact of changes in sea ice on coastal and marine ecosystems including species life cycles, population sizes, migration patterns, species distribution, species diversity and productivity;

o Impact of changes in sea ice on the distribution and productivity of polynyas;

o Impact of changes in sea ice on traditional lifestyles and human activities such as fishing, hunting and trapping, transportation, shipping and navigation, oil/gas exploration and extraction, search and rescue or tourism and recreation;

o Impact of changes in iceberg calving rates on human activities such as transportation, shipping, oil/gas activities and tourism;

o Impact of changes in multi-year sea ice extent on Canadian defence and sovereignty issues.

• Changing Human Use o Impact of the lack of understanding on how the coastal system naturally responds to

erosion, flooding and sea-level rise;

o Impact of lack of planning to consider coastal system response to climate change;

o Impact of increased demand for use, access to, development and protection of the coast.

• Miscellaneous o Impacts of permafrost degradation on shoreline stability, coastal ecosystems, coastal

communities, infrastructure, traditional lifestyles and human activities such as fishing, hunting and trapping, transportation, shipping and navigation, oil/gas exploration and extraction, search and rescue or tourism and recreation;

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o Impact of increased rates of organic carbon input to the nearshore due to increased erosion;

o Impact of increase methane release by permafrost degradation and coastal erosion.

2.5.3.3 Impacts on AtlanticCoastlines

The potential impacts of sea level rise on the Atlantic coastlines include [after C-CIARN, 2005]:

• Changing Oceanographic Processes o Impact of potential changes in ocean temperature, salinity, density and circulation

patterns.

• Changing Sea Levels (including storm surge, erosion and flooding) o Impact of sea-level rise (sea-level decline in Labrador) on tidal spectrum, wave

climate,coastal circulation, sediment redistribution and other physical processes;

o Impact of sea-level rise on coastal stability, particularly in vulnerable low-lying coastal regions and along soft sediment shores;

o Impact of sea-level rise (sea-level decline in Labrador) on coastal ecosystems;

o Impact of sea-level rise on salinization of estuaries and freshwater supplies;

o Impact of sea-level rise (sea-level decline in Labrador) on coastal communities and existing coastal infrastructure, including both unprotected and those currently protected by dykes or other structures;

o Impact of sea-level rise on archaeological or cultural resources;

o Impact of sea-level rise (sea-level decline in Labrador) on human activities such as agriculture, fisheries, hunting, transportation or tourism and recreation;

o Impact of sea-level rise on human health and safety, emergency preparedness, property loss, insurance or construction, maintenance and repair of coastal infrastructure;

o Impact of sea-level rise on shoreline development, real estate and legal issues;

o Impact of sea-level rise (sea-level decline in Labrador) on property ownership, legal boundaries and jurisdictional issues.

• Changing Weather Conditions o Impact of changes in precipitation on runoff, drainage, slope stability, sediment

redistribution and flooding;

o Impact of changing wind regime and storm activity on wave climate, coastal circulation, sediment redistribution and other physical processes;

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o Impact of changing storm activity on storm surge events;

o Impact of changing storm activity on shoreline stability, particularly in vulnerable lowlying coastal regions and along soft sediment shores;

o Impact of changing storm activity on coastal ecosystems;

o Impact of changing storm activity on coastal communities and coastal infrastructure;

o Impact of changing storm activity on human activities such as agriculture, fisheries, transportation or tourism and recreation;

o Impact of changing storm activity on human health and safety, emergency preparedness, property loss, insurance or construction, maintenance and repair of shore protection structures.

• Changing Sea Ice Conditions (including extent, thickness, location and season) o Impact of changes in sea ice on waves and currents (e.g. more open water results in

larger fetch for winter storm wave development);

o Impact of changes in sea ice on coastal circulation, sediment redistribution and other physical processes;

o Impact of changes in sea ice on coastal stability and flooding (e.g. less shore protection from ice, increased ice pile-up and ice ride-up, flooding due to ice jamming in bays, river mouths and estuaries);

o Impact of changes in sea ice on coastal and marine ecosystems;

o Impact of changes in sea ice on coastal communities and existing infrastructure;

o Impact of changes in sea ice on human activities such as fishing, commercial shipping, hunting, transportation, navigation, oil/gas activities, search and rescue or tourism and recreation.

• Changing Human Use o Impact of lack of understanding on how the coastal system naturally responds to

erosion, flooding and sea-level rise;

o Impact of lack of planning to consider coastal system response to climate change;

o Impact of increase demand for use, access to, development and protection of the coast.

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2.5.3.4 Impacts on PacificCoastlines

The potential impacts of sea level rise on the Pacific coastlines include [after C-CIARN, 2005]:

• Changing Oceanographic Processes o Impact of potential changes in ocean temperature, salinity, density and circulation

patterns;

o Impact of increases in ocean surface temperature on marine chemical and biological composition, species physiology, species abundance, distribution, diversity or ecosystem productivity;

o Impact of increases in sea surface temperature on ecosystem health and human health (e.g. algal blooms, red tide).

• Changing Sea Levels (including storm surge, erosion and flooding) o Impact of sea-level rise on tidal spectrum, wave climate, coastal circulation, sediment

redistribution and other physical processes;

o Impact of sea-level rise on coastal stability, particularly in vulnerable low-lying coastal regions and along soft sediment shores;

o Impact of sea-level rise on coastal ecosystems;

o Impact of sea-level rise on the physical, chemical and biological composition of wetlands and the effects on population sizes, species distribution, species diversity and ecosystem productivity;

o Impact of sea-level rise on salinization of estuaries and freshwater supplies;

o Impact of sea-level rise on low-lying coastal communities and existing coastal infrastructure, including both unprotected and those currently protected by dykes or other structures;

o Impact of sea-level rise on archaeological and cultural resources;

o Impact of sea-level rise on human activities such as agriculture, fisheries, hunting, transportation or tourism and recreation;

o Impact of sea-level rise on human health and safety, emergency preparedness, property loss, insurance or construction, maintenance and repair of coastal infrastructure;

o Impact of sea-level rise on shoreline development, real estate and legal issues;

o Impact of sea-level rise on property ownership, legal boundaries and jurisdictional issues.

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• Changing Weather Conditions o Impact of changes in precipitation on runoff, drainage, slope stability, sediment

redistribution and flooding;

o Impact of changing wind regime and storm activity on wave climate, coastal circulation, coastal upwelling, sediment redistribution and other physical processes;

o Impact of changing storm activity on storm surge events;

o Impact of changing storm activity on shoreline stability, particularly in vulnerable low-lying coastal regions and along soft sediment shores;

o Impact of changing storm activity on coastal ecosystems;

o Impact of changing storm activity on coastal communities and coastal infrastructure;

o Impact of changing storm activity on human activities such as agriculture, fisheries, transportation, or tourism and recreation;

o Impact of changing storm activity on human health and safety, emergency preparedness, property loss, insurance or construction, maintenance and repair or shore protection structures.

• Changing Human Use o Impact of lack of understanding on how the coastal system naturally responds to

erosion, flooding and water-level fluctuation;

o Impact of lack of planning to consider coastal system response to climate change;

o Impact of increased demand for use, access to, development and protection of the coast.

2.5.4 C-CIARN and the Identification of Canada’s Vulnerability to Water Level Change

The C-CIARN Coastal Zone Sector began operation in January 2002. The office is hosted by

Natural Resources Canada, as part of the Geological Survey of Canada (Atlantic). The mandate of C-

CIARN Coastal Zone is to develop a network of researchers, stakeholders and decision-makers

working together to improve Canada’s knowledge to the vulnerability of Canada’s coasts to climate

change and to seek solutions to projected impacts.

One of the first tasks of C-CIARN Coastal Zone was to draft a list of potential climate change

impacts on Canada’s coastal zone. Using The Canada Country Study: Climate Impacts and

Adaptation [Environment Canada, 1997a; Environment Canada, 1997b] and other reference

documents, as well as discussions with researchers and stakeholders from across Canada, preliminary

lists of the most important coastal climate change issues for each of the 4 coastal regions were

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compiled. From these regional lists, a national summary of the most important climate change

impacts for Canada was also compiled. Building on these lists, C-CIARN Coastal Zone then drafted

a list which identified important coastal climate change issues from a stakeholder perspective.

It was from these lists that the workshop theme of water level changes for coastal communities

was identified. To address the national scope of water level changes, both sea-level rise on ocean

coasts and decreasing lake levels on the Great Lakes needed to be considered. Specific issues to be

incorporate into the workshop included the affects of water level changes to coastal infrastructure,

utilities, property and community development, as well as the implications of water level changes to

human safety, disaster mitigation, cultural resources, tourism and recreation, property values,

insurance, legal and jurisdictional issues.

The objective of the C-CIARN Coastal Zone workshop was to bring together researchers and

stakeholders to exchange knowledge, ideas, experiences and concerns on coastal climate change

issues across Canada, as well as identify knowledge gaps and define a priority research agenda. It

was also hoped the workshop would facilitate research by enabling researchers and stakeholders to

develop contacts and explore partnership opportunities. From this workshop, five priority

information needs and their associated recommended actions were identified. These are discussed

below [after Parlee, 2004].

2.5.4.1 Water Levels

Improved information and scientific understanding of water level changes is the first and primary

need for coastal communities.

Knowledge gaps within this theme included:

• historical variations and trends in water level changes at a regional and local scale

• improved projections of water level change at a regional and local scale for the next 10, 20, 50 and 100 years;

• improved information on factors that influence water levels such as regional vertical land movement, thermal expansion of water bodies, changes in tidal regime, and climatic variables such as precipitation and/or evaporation rates.

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Recommended actions to meet these requirements included:

• more support and funding for research, measuring, monitoring and modelling of water level changes (particularly at Federal level);

• high resolution GPS measurements of present day ocean surface water levels;

• maintenance of existing coastal tide gauges and installation of additional gauges;

• mapping and predictive modelling of coastal water levels between tide gauges;

• improved network of instruments that measure and monitor the factors that influence water levels at a regional scale (i.e. vertical land movements);

• development of better models for scenarios and projections, particularly for regional and local scale;

• regional climatic and physical inputs for GCMs;

• official, accurate, easily accessible data and maps showing regional and local rates of relative and projected water level changes.

2.5.4.2 Mapping and Surveying

The second priority need identified was the necessity for easily accessible, accurate, high

resolution maps of coastal and inshore topography and integrated maps of physical, biological and

socio-economic factors within the coastal zone.

Knowledge gaps within this theme included:

• baseline mapping at a regional and local level including:

o 3-D high resolution digital topographic and nearshore mapping (i.e. DEMs);

o geomorphic mapping (i.e. landfeature, geology, lithology, water level limits, shore processes);

o ecosystem and biophysical mapping (i.e. habitat, species);

o land-use mapping (i.e. infrastructure, resource, human activity, cultural)

• historical mapping at a regional and local scale (i.e. shoreline positions, erosion rates);

• predictive modelling of potential physical changes at a regional and local scale and over short, moderate and long;

• time frames (i.e. shoreline positions, erosion rates, wave development, storm surge, ice changes);

• integrated mapping (GIS).

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Recommended actions to meet these requirements included:

• more support and funding for mapping and surveying within the coastal zone (particularly at Federal level);

• development of consistent terminology, methodologies or protocols (i.e. define shoreline, data collection standards);

• increased efforts to combine mapping initiatives and activities in an attempt to save time and expense (i.e. aerial photography, LIDAR, CASI, tidal measurements, sea level rise, water temperature);

• improved monitoring of physical, biological and socio-economic processes and changes within the coastal zone;

• integration of traditional and local knowledge with scientific data;

• development of more reliable models for predicting potential physical changes within the coastal zone;

• identification of entity or group(s) responsible for dissemination, management and distribution of data;

• make data and mapping products easily available to practitioners and decision-makers.

2.5.4.3 Vulnerability and Risk Assessment Mapping

Vulnerability mapping was identified as the third step in ensuring communities better understand

the impacts of water level changes and determine how best to adapt. According to the

Intergovernmental Panel on Climate Change (IPCC), vulnerability is the degree to which a system is

susceptible to, or unable to cope with, adverse effects of climate change, including climate variability

and extremes. For the purpose of this report, vulnerability includes physical, social, cultural and/or

economic susceptibilities to water level change.

Knowledge gaps within this theme included:

• identification, mapping and analysis at a regional and local level of:

o vulnerable shorelines and/or ecosystems;

o natural hazards (i.e. erosion, flooding);

o vulnerable people/communities, infrastructure, sectors, land-use, human activities and/or economic, cultural and heritage resources;

• identification and analysis of community values and priorities, including risk perception and willingness to pay;

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• identification and analysis of economic vulnerability at a regional and local level (i.e. cost/benefit analysis);

• resilience or ability of ecosystems to naturally adapt to water level change;

• capacity of a community, sector or activity to adapt to changing water levels.

Recommended actions to meet these requirements included:

• more financial and technical support for mapping, assessment and analysis of vulnerability;

• develop guidelines, standards and methodologies for vulnerability assessments;

• improved methodology for impact costing and evaluation of non-monetary assets;

• integration of traditional and local knowledge with scientific data (i.e. improve stakeholder involvement);

• more vulnerability case studies (ex: recent CCAF funded project conducted in Charlottetown, PEI);

• dentification of entity or group(s) responsible for dissemination, management and distribution of data;

• make baseline and risk assessment data easily available to practitioners and decision-makers to assist in preparing detailed community scale vulnerability maps.

2.5.4.4 Adaptation Options and Decision-Making

The need for more information on adaptation options and the decision-making process was

identified as the next step in ensuring communities better understand water level changes and how to

adapt.

Knowledge gaps within this theme included:

• risk management practices for community planning and decision-making processes;

• integrated management frameworks for community planning and decision-making processes;

• identification and understanding of the range of adaptation strategies available and appropriate for communities;

• choosing an appropriate decision-making process and choosing among different adaptation options;

• defining or assessing the viability of a particular adaptation option:

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o cost/benefit analysis - economic, social and/or cultural;

o cost of adaptation alternatives (e.g. protection works vs. relocation);

o identification and assessment of direct and indirect impacts from different adaptation options;

o how to address conflict between different users and stakeholders;

o better understanding of existing maladaptation;

• tools and technology to facilitate adaptation.

Recommended actions to meet these requirements included:

• more case studies;

• development of toolkits to help communities better understand and/or develop guidelines for adaptation;

• establishment of an organization that would provide information to communities about adaptation options and on-going adaptation projects across Canada (ex: "Partners for Climate Adaptation" - modelled on the Federation of Canadian Municipalities ''Partners for Climate Protection");

• develop decision-making process or framework for practitioners and/or policy makers to help integrate community, private sector and stakeholders needs into climate change adaptation initiatives;

• involve communities and stakeholders in identifying appropriate adaptation strategies and adaptation strategy evaluation;

• identification or development of policies or strategies to reduce risk (i.e. insurance prevention programs, regulatory planning policies, municipal by-laws, legal options);

• support and financing for community adaptation;

• capacity development to help communities understand, integrate data and make informed decisions.

2.5.4.5 Education and Communication

Workshop participants also noted the necessity for education and outreach as a priority. This

activity needs to occur concurrently with all of the above recommendations.

Knowledge gaps within this theme included:

• awareness of climate change impacts, their implications and adaptation options with the government, researchers, professionals, communities, decision-makers and the general public (i.e. all sectors/all levels);

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• effective communication between researchers and stakeholders (i.e. what to communicate, clear definition of terms, definition of goals, format, timing, avoiding conflicting information).

Recommended actions to meet these requirements included:

• improved flow of information between government, researchers, communities and stakeholders;

• improved two-way consultative process (i.e. stakeholder involvement, incorporation of scientific and traditional/local knowledge, integration of various interests, communications between various interests);

• improved information sharing and more effective mechanisms for information exchange between communities and stakeholders;

• capacity development and training in schools

2.5.5 References

C-CIARN, Coastal Zone Primary Impact Issues, Online document, http://c-ciarn.bio.ns.ca/ impacts_e.php (Last accessed March 2005), 2005.

CCRS, Shoreline by Province and Territory (in kilometres) in The Atlas of Canada: Facts about Canada, Coastline and Shoreline, Online document, http://atlas.gc.ca/site/english/ facts/coastline.html (Last accessed February 2005), Canadian Centre for Remote Sensing, GeoAccess Division, Natural Resources Canada, 2000.

Environment Canada, The Canada Country Study: Climate Impacts and Adaptation, Volumes I to VI, Regional Volumes, Online document, http://www.climatechange.gc.ca/english/issues/ how_will/canada_country.shtml, Ottawa, Canada, 1997a.

Environment Canada, The Canada Country Study: Climate Impacts and Adaptation, Volumes VII and VIII, National Sectoral and Cross-Cutting Issues, in Online document, http://www.climatechange.gc.ca/english/issues/how_will/canada_country.shtml, Environmental Adaptation Research, Ottawa, ON, Canada, 1997b.

Hengeveld, H.G., Projections for Canada's climate future: A discussion of recent simulations with the Canadian Global Climate Model, Climate Change Digest Series, CCD 00-01, pp. 27, Environment Canada, Ottawa, ON, Canada, 2000.

IPCC, Emissions Scenarios, in A Special Report of Working Group III of the Intergovernmental Panel on Climate Change, edited by N. Nakicenovic, and R. Swart, pp. 599, Cambridge University Press, Cambridge, UK, 2001.

Natural Resources Canada, Degrees of variation: Climate change in Nunavut, Online document (http://adaptation.nrcan.gc.ca/posters/articles/nu_05_en.asp?Region=nu&Language=en), CCCIARD, Natural Resources Canada, 2002.

Parlee, K.A., The highs and lows of water level: The vulnerability of coastal communities to water level change, in C-CIARN Coastal Zone Workshop 2003, pp. 16, C-CIARN Coastal Zone Report 04-1, Dartmouth, NS, Canada, 2004.

Shaw, J.H., R.B. Taylor, D.L. Forbes, M.-H. Ruz, and S. Solomon, Sensitivity of the Coasts of Canada to Sea-level Rise, Geological Survey of Canada, Ottawa, ON, Canada, 1998.

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2.6 Grasslands

2.6.1 The Importance of Grassland Ecosystems

Grasslands are plant associations containing few trees and are characterized by a cover of mixed

herbaceous vegetation that is dominated by grasses (family Poaceae) [Risser, 1985]. This mixed

vegetation layer usually includes a significant proportion of species from the legume (Leguminosae),

composite (Compositae) and sedge (Cyperaceae) families, and sometimes includes shrubs and sparse

trees [Kephart et al., 1993]. Water availability (or more specifically, a lack thereof) is the principal

climatic determinant affecting the development of grassland regions. Natural grasslands generally

occur in climates that are too arid to support a fully developed forest, but not so adverse as to inhibit

the development of a closed perennial herbaceous layer.

Permanent grasslands cover approximately 26% of the earth's total land surface [Ikeda et al.,

1999]. These regions are economically and ecologically significant. The economic importance of

grasslands is primarily a result of their strong links to global food production. Grasslands with

suitable climates and fertile soils are almost always quickly converted into croplands [Brown, 1989].

Where conditions are too arid for crop production, grasslands often exist as rangelands – the natural

pastures for grazing animals from which a considerable portion of the world draws its protein foods

[Coupland, 1979]. Grassland ecosystems are also ecologically important, and play a major role in the

global carbon cycle. These systems may contain as much as 30% of the earth’s total carbon stocks

[Ojima et al., 1996; Parton et al., 1996], and their annual contribution to global net primary

production may be as much as 16% of that of the terrestrial biosphere (19 x 109 tons C year-1)

[Whittaker and Likens, 1973].

2.6.2 The Grasslands of North America

2.6.2.1 The North American Grassland Biome

The grassland biome dominates the interior of the North American continent, extending west-east

from the Western Cordillera to the eastern deciduous forest, and north-south from central

Saskatchewan to Mexico [Bragg, 1995]. This biome contains the widest diversity of grassland types

on earth [Kephart et al., 1993], and is home to grasses, grass-like plants (sedges and rushes), forbs

1

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and woody plants [Risser, 1985]. Indeed, far more species of these latter taxa are actually present,

but grasses constitute most of the biomass.

The North American grasslands occur across a large climatic range. Grassland systems are found

in semiarid to subhumid regions, where mean annual precipitation (MAP) ranges from 250mm to

1000mm [Bragg, 1995; Brown, 1989; Risser, 1985]. Drier areas generally support deserts whereas

wetter areas generally support forests [Kephart et al., 1993]. Grassland mean annual temperatures

(MAT) typically range from –7ºC to 30ºC [Whittaker, 1975]. These climatic variations exist as a

west-to-east gradient of increasing MAP and a north-to-south gradient of increasing MAT. While

this precipitation gradient is a product of the rainshadow effect of the entire Western Cordillera, the

temperature gradient is controlled by latitude. The interaction of these two factors, prevailing soil

conditions and management histories largely determine the distribution of various grassland types

across North America [Bragg, 1995].

The largest unbroken grassland formation in North America stretches from the Appalachians to

the Rocky Mountains. This expanse is covered by tallgrass, mixed and shortgrass prairie, which

grade from one type to another, arranging themselves along moisture, temperature, and soil gradients

[Bragg, 1995]. The tallgrass prairie, which occupies the eastern part of the Great Plains, receives

more rainfall (as much as 1000mm per annum) than the mixed and shortgrass prairies to the west

(500mm per annum and 300mm per annum, respectively). The shortgrass prairie occupies the driest

part of the Great Plains. Between these two moisture extremes lies the mixed grass prairie. However,

these three prairie types are not absolutely restricted to separate geographic zones. In an area

generally identified as mixed grass prairie, tallgrass communities may exist in low-lying moist spots,

while shortgrass communities may develop along dry, rocky outcrops. In such instances it is

available soil moisture, rather than rainfall and evaporation, that is the controlling factor on species

distribution [Brown, 1989]. Other grassland communities are separated as much by geography as by

vegetation [Brown, 1989]. Desert grassland is found on the edges of the southwestern deserts,

California grassland is found in the Central Valley of California, and Palouse prairie is found on the

volcanic soil of the Columbia Plateau.

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Figure X. The northern mixed grass prairie (modified from Coupland (1992)).

2.6.2.2 The Canadian Northern Mixed Grass Prairie

The northern mixed grass prairie occupies the

northern portion of the Great Plains grasslands. This

biome extends northwards from the northern boundary

of the short-grass steppe to the southern border of the

fescue prairie, and eastwards from the foothills of the

Rocky Mountains to the western border of the tallgrass

prairie (Figure X) [Lauenroth et al., 1994]. This region

includes large portions of Alberta (34,000 km2) and

Saskatchewan (25,000 km2) [Samson and Knopf,

1994]. The climate of the region is characterized by

great annual extremes in temperature (ranging from –7

°C to 17 °C in the southwest and from –14 °C to 24 °C

in the northeast) and comparatively low mean annual

precipitation (300mm to 450mm) [Bryson and Hare, 1974]. Low precipitation, high summer

temperatures and dry winds (largely prevailing from the west) frequently make soil moisture the

limiting factor for plant growth [Coupland, 1992]. The mean climate, however, is not the shaping

feature in this region. Rather, it is these aforementioned extremes in temperature and precipitation

that create hardy, drought and frost tolerant communities.

2.6.3 Predicting grassland response to future climates

2.6.3.1 Grassland responses to future climates

It is predicted that a doubling of atmospheric CO2 over pre-industrial levels will occur between

2050 and 2100 AD. This change in atmospheric composition is expected to bring about an average

global increase in air temperature ranging from 1 – 3.5ºC, a warming that is expected to be greatest

at mid-to-high latitudes [Flato et al., 2000; Maxwell, 1997]. While most GCM projections for the

Canadian Prairies show considerable increases in temperature under global warming, there is less

agreement as to whether increases in total annual precipitation will also occur [Hengeveld, 2000].

Shepherd and McGinn [2003] and McGinn and Shepherd [2003] modeled the impact of various

climate change scenarios on the agroclimate of the Canadian prairies. Their model results predicted

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an increase in minimum and maximum air temperatures ranging from 3 – 5ºC and increases in total

annual precipitation ranging from 4 – 32% (above 1960–1989 historic values). However, these

precipitation increases were not predicted to occur uniformly across the region. Rather, model results

suggest that while summer and winter increases in precipitation may occur in central Alberta, only a

slight increase in precipitation – or even a slight decrease – will occur in the southern and eastern

prairies.

These predictions will have significant

short- and long-term implications for the

prairie grasslands. While the model

projections described above suggest that no

major change in drought frequency will occur

in parts of Alberta, they do indicate that the

frequency of drought and severe drought

could increase dramatically in the southern

and eastern prairie regions (Figure X). In

these regions, higher summer temperatures

will increase evaporation and intensify

drought conditions. Over longer periods of

time, this may lead to the establishment of a semi-desert zone in the areas most deeply affected by

aridity [Rizzo and Wiken, 1992]. However, an increase in precipitation also means that there may

also be wetter periods when temperatures are cool. Overall, this suggests that soil moisture

conditions could become more variable in many places. These changes are projected to have an

influence on the productivity, plant composition, and ultimately, spatial extent, of the Canadian

grassland ecosystem.

The productivity and plant composition Canadian prairie is likely to be highly responsive to

changes in climate because it is already highly moisture and nutrient limited [Yang et al., 1998].

However, it is unlikely that changes in these parameters will result from climate warming alone.

Other environmental disturbances, such as land use change, nutrient enrichment, irrigation, fire and

grazing also affect grassland productivity and plant composition [Goodin and Henebry, 1997]. While

some disturbances (e.g. drought) modify grasslands directly by reducing the photosynthetic

capabilities of existing species, others (e.g. grazing, fire) often do so indirectly by bringing about

Figure X. Predicted changes in moisture conditions in southern Saskatchewan under climate warming (after Williams (1988)).

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changes in community composition. The impacts of many simultaneous disturbances are complex;

disturbances rarely, if ever, occur independently [Coppedge et al., 1998; Sage et al., 1999] and the

presence of many confounding variables makes the isolation of cause and effect difficult. As a result,

although it is predicted that large-scale changes in some or all of these factors may combine to

dramatically affect grassland biogeochemistry and their carbon stocks, the combined effects of these

disturbances on the overall functioning of grassland systems are complex and are not well

understood.

The spatial extent of the Canadian prairie

is also likely to be significantly affected by

changing climate. Rizzo and Wiken [1992]

used a climate–vegetation classification model

to examine spatial shifts in 10 ecoclimatic

provinces under two doubled-CO2 scenarios

and found that the boreal forest was projected

to be reduced from 28.9% of the land area to

14.9% and to be displaced northward by an

average of 500 km (Figure X). This decrease

is predicted to occur because grassland and

temperate deciduous species invade from the south, while the northern expansion of boreal forest is

limited by poor soils and insufficient sunshine amounts [Figure X, Environment Canada, 2005].

Rizzo and Wiken’s projections suggest that Canada’s grasslands will expand by as much as 7%, and

that a semi-desert zone will appear in southern Saskatchewan and Alberta. While reductions in the

extents of boreal forest and taiga and expansions of grassland and temperate deciduous species are

generally consistent with other modeling attempts [e.g. Lenihan and Neilson, 1995; Scott et al.,

2002], the exact size of these changes seems to be highly dependent on the GCM used.

Figure X. Projected changes in grassland distribution under CO2-doubling (after Rizzo

and Wiken, 1992)

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2.6.3.2 In situ monitoring of grassland dynamics

2.6.3.2.1 The role of in situ monitoring

In situ measurements of grassland conditions are necessary for monitoring grassland

dynamics through both space and time. In recent years, in situ observations have become

increasingly valuable for calibrating satellite data and supporting and validating models.

2.6.3.2.2 Case study: Climate change impacts on Sand Hills in the Canadian Prairies

Aeolian processes affect much of the southern Canadian prairies, as evidenced by sand dunes,

loess, dust storms and deflation of cultivated soils [Catto, 1983; Wheaton, 1992; Wheaton, 1990].

Sand hills are a significant physical feature of the Prairie Provinces, with more than 120 areas

occurring in Alberta, Saskatchewan and Manitoba. These are typically post-glacial sandy deposits

that have been reworked into dunes by wind at various times throughout the Holocene [Muhs and

Wolfe, 1999]. In the drier sub-humid prairie, sand hills reside in a delicate balance between bare

active dunes and vegetation-stabilized hills that are sensitive to changing climate.

The sensitivity of sand hills on the

Canadian prairie and their diversity of

ecosystems and land uses, makes them

particularly relevant to investigations of

the potential impacts of climate change.

A further significance of examining sand

hills is that they represent similar

landscapes occurring in different settings

in the prairies. Therefore, the influences

of potential climate change on sand hills

within different ecoregions can be compared. Changes in grassland composition and productivity in

different prairie ecoregions will affect livestock grazing and wildlife habitats to varying extents.

Wolfe and Thorpe [2005] examined the present conditions of sand hills in the Prairie Provinces and

assessed the potential impacts of climate change on the ecology and land use of these areas along

with the potential adaptive responses that may be needed to manage these impacts.

Figure X. Distribution of sand dunes in the Canadian Prairies (after David, 1977). Bold dashed line depicts limit of Palliser Triangle, fine dashed line depicts limit of brown soils zone.

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The Wolfe and Thorpe [2005] study identified six sand hill areas for analysis. These study

regions were located in the Mixed grassland, Moist Mixed Grassland, and Aspen Parkland

ecoregions of the Prairie ecozone and the boreal transition ecoregion of the Boreal Plains ecozone.

The land use was diverse in these regions, with transportation, recreation and conservation occurring

in all six sand hills and grazing occurring in all but one. Detailed management reports were available

for most of these selected areas. These reports provided a means of assessing the present sand hills

ecosystems and management practices.

Climate change data from seven internationally recognized models were obtained from the

Canadian Climate Impacts Scenarios project of Environment Canada [Canadian Climate Impacts

and Scenarios, 2001]. Each model simulated increasing temperatures compared to present conditions

(1961-1990 normals) in the sand hill study regions. The models predict a rise in the order of 1-3ºC by

the 2020s, a rise of 3-5ºC by the 2050s and a rise of 3-7ºC by the 2080s. These models generally also

simulate increased precipitation over the same time period.

The potential impacts of climate on the vegetation in the six study areas were analyzed in three

ways. These were (a) an analogue approach, (b) a regression model approach, and (c) a modeling

approach. The analogue approach identified areas in the US that had a warmer climate similar to that

projected for the study regions. The main analogue source was the Nebraska Sand Hills, an area of

about 5 million hectares, as well as smaller areas of dunes extending from Nebraska into northern

Colorado and southern South Dakota. The regression model approach used the relations between

climate and productivity of grasslands using the simple regression derived by Sims et al., [1978]and

Epstein et al., [1997). These models were applied to climatic data in the six study areas, and the

percentage changes from the 1961-1990 base period to the 2050s were calculated. The modeling

approach used the models of Muhs and Maat {, 1993 #1791] and Wolfe [1997], that incorporated

wind strengths and the ratio of annual precipitation to potential evapotranspiration (P:PE). These

models assessed potential changes in dune activity based on 2050s changes in P:PE ratios, assuming

no change in wind regime.

The results of the analogue study showed that by the 2050s, the climates of the study areas

approached the current climates of the analogue areas in the US. The climate of the three driest study

areas in the 2050s (Middle, Great and Dundurn/Pike Sand Hills) approximate the current climate of

western Nebraska and northeastern Colorado. The coolest areas in the 2050s (Manito Lake and

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Nisbet/Fort à la Corne) are similar to the present climate of central North Dakota. The simulated

future climate of the Brandon Sand Hills approximates that of the current Nebraska Sand Hills. A

significant finding from this research is that all of the focus areas could experience climatic

conditions that are less suitable for tree growth. This impact could be greatest for the Nisbet and Fort

à la Corne Sand Hills, currently almost entirely forested, and they could shift to a climate like central

North Dakota that supports only grassland [Thorpe et al., 2001].

The results of Wolfe and Thorpe [2005] also showed that the shift to drier climates mean that the

potential for dune activity increases, particularly in the driest areas. These results are similar to those

obtained by Wolfe [1997], who determined the level of dune activity that might occur if the climate

of the prairies were consistently warmer and more arid than at present. Wolfe’s [1997] study suggests

that increased aridity caused by prolonged drought or climatic warming would increase dune activity.

In regions where dune activity is presently greatest, this increase this increase would likely result in

an increase in the number of blowouts, the reactivation of stabilized dunes, and merging of presently

active sand dunes to form nearly continuous bare cover in some areas. Wolfe and Thorpe [2005]

suggest that dune activation is not a uniform process, but may develop in certain locations following

a series of dry years. If significant areas of active dunes develop, this could reduce the amount of

productive vegetation for livestock grazing and wildlife habitat. This increased potential for dune

activity could also increase erosion problems associated with roads, wellsites or military operations.

In grasslands, a more arid climate could favour warm-season (C4) over cool season (C3) species

[Thorpe et al., 2001], though the carbon fertilization effect of increased atmospheric CO2 would

favour the productivity of cool-season species. Thus, grasslands will have a considerable capacity to

adjust to climate change through shifts in the relative proportions of species that are already present.

Areas will likely undergo a gradual northwards migration of species currently absent or uncommon

in Canadian grasslands, especially warm-season grasses. Shifts in vegetation patterns may also be

accompanies by shifts in the populations and distributions of other species, because of habitat

changes and direct climatic effects. The impact on grazing capacity is unclear. A shift to a drier

climate suggests reduced forage production, and this is supported by the simple production models

that show decreased yields in the 2050s [Thorpe et al., 2001]. However, the current grazing

capacities of analogue areas in the US are rated higher than the Canadian focus areas, with the higher

portion of warm-season grasses and the lower woody cover possibly contributing to this difference.

155

Forested areas like Nisbet/Fort à la Corne are likely to increase grazing capacity with loss of tree

cover.

Wolfe and Thorpe [2005] present a variety of adaptive responses of land use activities to climate

change. If climate does change as projected by current GCMs, many physical, biological, economic

and social impacts can be expected. To minimize the harmful effects of these changes, and to

maximize the potential benefits, it is important to consider human adaptation options and to develop

strategies for coping with climatic change. These include the following strategies:

• Live stock grazing: Within most sand hills on the prairies, recommended stocking rates and

range assessment tables should be revised [Wroe et al., 1988] as new information on changes

emerges form monitoring range benchmark sites. The benefits and ecological consequences

of introducing warm season grasses could be reviewed. While reseeding of existing

grasslands is undesireable, use of such species in reclamation may be a viable adaptation to

climate change. It is likely that better fire-fighting capabilities will be required to protect

rangelands from an increased fire hazard.

• Forestry: In forested sand hill areas, calculations of allowable harvest, and harvest allocations

to individual operators, may require adjustment if evidence from monitoring programs

indicates changes in growth rate or incidence of disturbance. Increased spending on

regeneration following harvest or disturbance may be necessary to cope with regeneration

failures. Research on optimal harvest and regeneration methods for dry conditions may be

required. Increased fire protection may be needed as fire frequency increases. The benefits

and ecological consequences of introducing species from nearby North American ranges

should be reviewed.

• Oil, gas and mining development: Policies and regulations are already in place to control the

impact of oil/gas activities on soil erosion and other environmental issues. However, it may

be necessary to apply current policies more stringently as the climatic potential for dune

activation increases. The option for winter-only operations may become less viable in

southern areas as milder winters reduce the period during which the soil is frozen.

• Recreation: In most sand hills areas the existing dispersed recreation patterns will likely need

minimal adaptive adjustment. However, sites sustaining intensive recreation activity may

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require additional trail development to control pedestrian impacts and more reclamation of

disturbed sites. In general, greater regulation may be required for motorized recreation

vehicle use on sensitive sites.

• Conservation: Each of the areas considered in the study are to some extent remnant islands of

natural vegetation, and as such are considered important for conservation of wildlife habitat

and biodiversity. Current concepts of conservation are based on maintaining all the species

and ecosystems that are currently present. However, climate change could result in shifts in

current patterns as some species would be less adaptable to the new conditions. In such a

situation, a rethinking of conservation policy would be required.

Overall, the results of Wolfe and Thorpe [2005] and Wolfe [1997] indicate that there are likely to

be significant influences to sand dune areas in the Canadian prairies under projected changes in

prairie climates, and that these changes may include changes in grassland species composition and

productivity, reduction in forest productivity, and potential increases in the susceptibility of sand

hills to erosion. These physical and biological impacts, in turn, may have social and economic

consequences that will require society to alter its behavior and respond with adaptive strategies that

minimize negative economic and environmental consequences with the climate change.

Management planning practices in most sand hills areas are clearly more protective and

conservation-oriented than in other parts of the prairies. Progress in land-use planning is also

comparatively advanced. This situation should facilitate the adaptation to climate change. However,

detailed biophysical inventories should be completed for all areas to obtain a baseline for monitoring.

Land use plans should be developed for all areas and should consider the impacts of climate change

on future land uses. Land use plans should incorporate mechanisms for adaptive management.

Vegetation monitoring programs should be designed to include indicators of climate change impacts.

Repeat time-series aerial photography or satellite imagery can be used for monitoring or changes in

the woodland/grassland mosaic over large areas. As the Boreal Transition ecozone is susceptible to

change, forest monitoring programs in areas such as the Nisbet and Fort à la Corne Sand Hills should

be designed to include indicators of climate change impacts. Such indicators may include changes in

tree growth rates, drought-related tree mortality, incidence of insects and diseases and changes in

plantation survival.

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2.6.3.3 Remote sensing of grassland dynamics

2.6.3.3.1 The advantage of the remote sensing approach

Because grasslands generally cover large geographical extents and are mostly found in isolated

locations, the use of traditional assessment techniques in these regions is often time-consuming and

costly [Asrar et al., 1986]. Thus, other monitoring approaches must be ultilized. One possible

approach is the use of satellite remote sensing systems. Through the unique combination of extensive

spatial, spectral and frequent temporal data collection, remote sensing has the potential to provide

both scientists and managers with a powerful monitoring tool at regional to local scales [Goodin and

Henebry, 1997]. The remote sensing approach has great potential to provide quantitative information

on the amount, condition, and type of vegetation, provided that the effects of physical and

physiological processes on spectral characteristics of grassland canopies are fully understood [Asrar

et al., 1986].

The increasing availability of remotely-sensed data at various spatial and spectral resolutions

offers the potential to monitor the biophysical characteristics of ecosystems at various landscape

scales [Tieszen et al., 1997]. This has been particularly well demonstrated by those who have focused

their attentions on grasslands. Spectral observations over large tracts of grassland have typically been

acquired using National Oceanic and Atmospheric Administration (NOAA) Advanced Very High

Resolution Radiometer (AVHRR) data [Burke et al., 1991; Davidson and Csillag, 2003; Goetz,

1997; Paruelo and Lauenroth, 1995; Tieszen et al., 1997; Tucker et al., 1991]. While these spatial

resolutions (1-4km) provide valuable information about climatic and aggregate anthropogenic

forcings on vegetation dynamics [Henebry, 1993], they are often at a coarser spatial resolution than

some applications require. Medium spatial resolution satellite observations have been derived from

Landsat Multi-Spectral Scanned (MSS) [e.g. Mino et al., 1998; Pickup et al., 1993] and Landsat

Thematic Mapper (TM) sensors [e.g. Henebry, 1993; Henebry and Su, 1993; Mino et al., 1998; Todd

et al., 1998] at resolutions of 50m and 30m, respectively. At even finer resolutions, information has

been provided by airborne [e.g. Walthall and Middleton, 1992] and ground-based observations

[e.g.Asrar et al., 1986; Davidson and Csillag, 2001; Davidson and Csillag, 2003; Girard, 1982;

Goodin and Henebry, 1997; Goodin and Henebry, 1998; Pickup et al., 1993; Tucker et al., 1991;

Weiser et al., 1986], whose spatial resolutions are determined by sensor height and field of view

(FOV), respectively.

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2.6.3.3.2 Case study: Monitoring grassland vegetation water content using spectral data

The measurement of vegetation water content is important for determining the physiological

status of plants [Datt, 1999], and for evaluating drought and wildfire risk in natural plant

communities [Peñuelas et al., 1997]. These assessments are particularly important in environments

that frequently experience extreme water stress. The northern portion of the mixed grass prairie is an

example of such a region. Here, relatively little summer precipitation (150mm to 225mm), high

summer temperatures (often > 20 °C) and strong and dry prevailing winds frequently combine to

create a severely water-limited landscape during much of the growing season [Bryson and Hare,

1974; Coupland, 1992; Loveridge and Potyondi, 1983]. However, the sheer vastness this ecosystem

means that in situ observations are rarely sufficiently dense to accurately characterize its regional

variation in vegetation water content. Thus, other measurement approaches must instead be utilized.

One possible approach is the use of satellite remote sensing systems.

Davidson et al., [in review] assessed the potential of Landsat-TM data for assessing the

vegetation water content of grassland-shrubland plant communities in the Canadian mixed-grass

prairie. The authors collected vegetation water content and spectral radiometer data over plots of

comparable ground resolution (0.5 m) at seven field sites in Grasslands National Park,

Saskatchewan, in June 2004. Sites were randomly located in four of the major plant communities

found in the region, namely ungrazed and grazed grassland, and ungrazed and ungrazed shrubland.

These observations were then scaled “up” to an observational scale consistent with that of Landsat-

TM satellite imagery (30m). This allowed the potential of remote sensing for predicting vegetation

water content to be assessed at both observational scales.

The authors used a stratified random sampling design to locate 15 square sampling plots of 0.5m

resolution at each of the above sites. Plots were evenly distributed among areas of low, medium and

high productivity. This allowed a wide range of vegetation water contents (6 – 550 g H2O m-2) to be

represented within the sampling framework. An additional 9 plots were randomly located among the

grassland sites, giving a total sample size of n = 114. Each plot was sampled once during the field

campaign. Sampling was spread across the entire month, thereby allowing within-site temporal

variations in surface conditions (e.g. dry vs. wet soil; changes in plant phenology) to be adequately

sampled.

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Spectrally reflected radiation in Landsat TM bands 1 (0.45-0.52µm), 2 (0.52-0.60µm), 3 (0.63-

0.69µm), 4 (0.76-0.90µm) and 5 (1.55-1.75µm) was measured over each plot using a Cropscan

Model MSR5 Multispectral Radiometer (Cropscan Inc., Rochester, MN, USA). The radiometer has a

28° field of view (FOV) and was mounted normal to the ground surface at a height of 1m, giving a

spatial resolution of 0.5m. All spectral measurements were collected within 2h of solar noon on

under cloud-free sky conditions. The mean of three separate reflectance measurements was used as a

representative measure of plot reflectance. Plot reflectances were used to assess the predictive ability

of four spectral approaches. These were: a combination of spectral bands, a combination of spectral

derivatives, a combination of principal components generated from the spectral bands, and a

combination of principal components generated from the spectral derivatives.

Vegetation water content was measured as follows. Plots were clipped of all their standing

vegetation within 20 minutes of spectral sampling. Clipped vegetation was sorted immediately after

harvest then weighed. After fresh (wet) biomass weights were determined, clippings were transferred

to a paper bag and force air-dried in an oven at 60ºC for 48h. Dry biomass weights were determined

at the end of each drying cycle. The fresh and dry biomass weights were used to calculate three

measures of vegetation water content for each of the sorted clippings. These were (a) absolute water

content (AWC), calculated as the difference between the fresh and dry biomass weights, (b) relative

water content as a percentage of fresh biomass weight (RWCF), calculated as the AWC divided by

the fresh biomass weight, and (c) relative water content as a percentage of dry biomass weight

(RWCD), calculated as the AWC divided by the dry biomass weight.

The results of this study showed that (a) that the band-combination approach provides the

most accurate and precise estimates of AWC and RWC at both 0.5m and 30m sampling resolutions

(Figure X(a) and (d)); (b) that the combination of bands providing the greatest predictive abilities are

those that emphasize the contrast in reflectance between the near infrared and shortwave infrared

spectral regions; (c) that the band-combination approach predicts AWC with much greater accuracy

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Figure X. The scale dependence of the correspondence between predicted and observed vegetation water content using the Band-Combination approach [(a), (b), (c) = 0.5m; (d),(e),(f) = 30m]. Vegetation water content values

were predicted using a leave-one-out cross-validation approach (see text for details). Linear regression lines are fit separately to each method. The 1:1 lines in each inset represent perfect validity.

0 100 200 300 400

0

100

200

300

400

Obs

erve

d (g

H2O

/ m

2 )

AWC = 564.4 +

677.8(log(NIR)) - 977.3(log(SWIR))

(a) AWC

0.5m

0 100 200 300 400

0

100

200

300

400

Obs

erve

d (g

H2O

/ m

2 )

AWC = 496.3 +

755.2(log(NIR)) - 999.2(log(SWIR))

(d) AWC

0.3 0.4 0.5 0.6 0.7 0.80.3

0.4

0.5

0.6

0.7

0.8

Obs

erve

d (%

of F

resh

)

RWCF = 0.56 + 0.31(log(B))

+ 0.59(log(NIR) - 0.79(log(SWIR))

(b) RWCF

0.3 0.4 0.5 0.6 0.7 0.80.3

0.4

0.5

0.6

0.7

0.8

Obs

erve

d (%

of F

resh

) (e) RWCF

RWCF = 0.71 + 0.46(log(B))

+ 0.60(log(NIR) - 0.99(log(SWIR))

0.0 0.5 1.0 1.5 2.0 2.50.0

0.5

1.0

1.5

2.0

2.5

Obs

erve

d (%

of D

ry)

(c) RWCD

0.0 0.5 1.0 1.5 2.0 2.50.0

0.5

1.0

1.5

2.0

2.5

Obs

erve

d (%

of D

ry) (f) RWCD

RWCD = 2.52 + 2.27(log(B))

+ 2.87(log(NIR) - 5.09(log(SWIR))

RWCD = 2.49 + 1.96(log(B))

+ 2.79(log(NIR) - 4.82(log(SWIR))

Predicted Predicted

30m

161

and precision than RWC (Figure X(a) and (d)) vs (b),(c),(e) and (f) and, (d) that predictive ability

of the band-combination approach decreases only slightly when plot-level data are aggregated to a

30m sampling resolution.

The better performance of spectral approaches based on the near infrared/shortwave infrared

spectral space and the predictive abilities of each approach are generally consistent with the results

of other studies and with theory. While the results suggest of this study suggest that simple spectral

methods (e.g. simple band combinations or indices) are good estimators of absolute vegetation water

content at both plot and Landsat-TM spatial resolutions, they are less encouraging for the estimation

of relative water content. Furthermore, despite their good predictive abilities, the temporal and

geographical portabilities of the spectral approaches for estimating AWC must be further assessed

before they can be considered reliable and robust predictive tools. Thus, the further testing of these

techniques over larger geographical extents is required.

2.6.3.3.3 Case study: Impact of climate variations on surface albedo of a temperate grassland

Changes in land surface albedo are closely linked to ecosystem dynamics. Therefore, the

effects of climate change on ecosystem processes may lead to changes in surface albedo

characteristics. Since the physical climate system is very sensitive to surface albedo, ecosystems

could significantly feedback to the projected climate scenarios through albedo changes. As such,

impacts of climate change on surface albedo and ecosystem feedbacks have been recommended for

further investigation. This is of particular significance for those ecosystems whose structure is highly

responsive to climate change variations. One example is temperate grasslands in arid and semi-arid

regions, such as the Canadian prairies which dominates the south-central part of the Canadian

landmass. The phenological patterns, biomass production, and species composition of this

ecosystem, are strongly affected by the climatic conditions of the region, especially precipitation

which is highly variable both inter-annually and intra-annually [White et al., 2000]. However, the

effects of these changes on surface albedo have yet to be properly investigated.

Wang and Davidson [in review] used climate station observations and satellite albedo

observations to assess the impact of climate variations on the surface albedo of a temperate

grassland. They then discussed the implications of their results in the broader context of climate

change and ecosystem studies. This study focused on a portion of the Canadian mixed-grass prairie

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ecosystem (Grasslands National Park, Saskatchewan). This region was chosen for study because its

recent disturbance history is well documented.

Meteorological observations, including daily mean air temperature, rain and snow from

nearby weather stations were used in the study. These data were aggregated to create 16-day

averages to match the temporal resolution of the satellite albedo data. Three years of data (2001-

2003) were used in the study because each represented a very different climate regime. The year

2001 was significantly drier than its historical mean. As a result, drought severely affected

agricultural production over the region. In 2002, precipitation was significantly higher than its

historical mean. However, most of this fell during the late growing season, with drought still

occurring in the early growing season. In 2003, annual precipitation was slightly lower than its

historical mean, but more precipitation fell in the early growing season.

Satellite albedo observations were taken from the MODIS terra albedo product (16-day

composite, MOD43B3). Every 16 days, multitemporal, atmospherically corrected, cloud-free data

and a semi-empirical kernel-driven bidirectional reflectance model are used to compute a global set

of parameters describing the Bidirectional Reflectance Distribution Function (BRDF) of the land

surface. The MODIS albedo product (MOD43B3) is generated by integrating the BRDF parameters

over all viewing angles at local solar noon to produce direct and diffuse radiation albedo for each of

the 7 MODIS bands, as well as for the visible (VIS), near infrared (NIR) and total shortwave

broadbands (0.3-0.7µm, 0.7-5.0µm, and 0.3-5.0µm, respectively) [Schaaf et al., 2002]. The MODIS

albedos represent the best quality data possible for each 16-day period. A detailed description of the

MODIS albedo product, its creation, and validation, is provided by Wanner et al. [1997], Lucht et al.

[Lucht et al., 2000], Schaaf et al. [Schaaf et al., 2002], and Jin et al. [Jin et al., 2003a; Jin et al.,

2003b]. Wang and Davidson [in review] acquired MODIS direct radiation albedo data (MOD43B3,

validated version V004) for the entire Canadian landmass for 2001, 2002 and 2003 in the VIS and

NIR wavelengths. Their study was restricted to an analysis of direct radiation albedo because direct

radiation is the dominant component around noon under clear sky conditions [Wang et al., 2002].

They also restricted their study to the analysis of VIS and NIR broadbands because these

wavelengths are the most commonly used in climate and ecosystem models.

The GNP grassland is a natural ecosystem that has seen only little disturbance by human

management practices in recent history (i.e. within last 20 years or so). Thus, changes in ecosystem

163

characteristics in 2001, 2002 and 2003 were mainly induced by climate variations. The results of this

study indicate:

(1) In the winter season, the amount of precipitation (snow) greatly affects the surface albedo

of this ecosystem. The Canadian prairie has a semi-arid climate, and thus, the total amount of

snowfall in winter is generally low. Melting rarely happens in winter due to the low temperature, and

so snow amount directly determines the snow coverage fraction of the landscape. Since snow has

much higher albedo than the bare soil and dead grass biomass that also composes the ground surface

of the ecosystem, snow coverage fraction plays the most important role in determining winter season

albedo. Another reason for the strong impact of snow coverage fraction on surface albedo is that this

ecosystem is mainly composed of short grass species, and thus, the productivity and biomass

produced by this ecosystem is typically lower than other less arid grassland landscapes. The lower

productivity of this system means that the influence of snow on surface albedo is different to that of

other high latitude ecosystems (e.g. boreal forests). Since the solar zenith angle in winter is very

large for high latitude ecosystems, the probability of direct radiation reaching the snow surface under

forest canopies is very small [Wang, in press]. As such, even a full coverage of snow on the ground

of a forest can only increase the surface albedo by about 0.05-0.1 [Betts and Ball, 1997; Wang, in

press]. It is also worth mentioning that the high wind speed, which is another characteristic of the

climate over the Canadian prairies, can also influence the surface albedo of the grassland by (a)

redistributing the snow amount through blowing snow and (b) changing the snow coverage fraction

through uneven snow accumulation (drifting). Although the study region has a rolling landscape,

significant changes in topography can occur over relatively short distances. Strong wind over the

region can thus cause uneven distribution of snow over the surface in winter.

(2) During the winter-to-summer and summer-to-winter transitional periods, air temperature

plays an important role in determining surface albedo. Climate records show that temperatures over

the region are highly variable from year to year. The date that air temperature rises above 0°C in

spring or autumn can differ by more than one month among years. Since snow melting in spring and

the start of snow fall and accumulation in autumn are strongly related to air temperature, surface

albedo during these time periods is highly variable, and depends on snow absence or presence. In our

results, 2002 showed a much higher albedo in March and early April than the other two years. This

was likely was caused by the delay of snow melt due to the lower air temperatures during that

164

season. The relatively high albedo in early November of 2003 demonstrates another example of how

air temperature influences surface albedo in autumn.

(3) In the growing season, ecosystem water conditions can significantly alter the surface

albedo of the semiarid grassland (see Figure X ). Because precipitation is the only source of surface

water for the ecosystem, the highly variable annual precipitation and its seasonal distribution can

result in albedo variations at different temporal scales. Our results clearly indicate that climate

drought causes higher surface albedo, and improved ecosystem water conditions cause lower surface

albedo. This is probably a result of two processes. First, soil albedo changes with the surface soil

water content. Under most circumstances, soil albedo increases when soil water content decreases. In

GNP, a significant amount of radiation can be reflected by the soil surface due to the low vegetation

cover in places (e.g. patchy shrubland, eroded communities). Second, vegetation growth is strongly

controlled by the water conditions in this semi-arid ecosystem. Because of the high temperature, high

solar radiation, and low precipitation in summer, vegetation growth is frequently constrained by

water deficit. As a result, improved water conditions by precipitation can significantly increase the

vegetation growth and biomass accumulation, and vice versa. Vegetation growth can affect surface

albedo when: (a) the plant canopy has a much lower albedo in the visible broadband than that of the

soil surface; (b) changes in plant growth rates significantly alter the green coverage fraction of

Figure X. Difference in albedo between dry (2001) and wet (2003) Junes for the Canadian prairies (after Wang et al., (2004).

Visible albedo June 2001 Visible albedo June 2003

NIR albedo June 2001 NIR albedo June 2003

Difference in visible band

Difference in NIR band

165

canopy in this grassland ecosystem; and (c) differences in plant growth result in the different ratios

of green-to-dead biomass and different amount of litter falls and dead biomass coverage in the later

growing season. Grass litter and dead biomass have high reflectance in both VIS and NIR bands

[Asner, 1998], which can affect the grassland albedo, particularly in late summer. This likely

contributed to the higher NIR albedo in the late summer of 2003, which had higher precipitation in

the early growing season and drier conditions in the late growing season compared to of 2002, which

had a drier early growing season. It is worth mentioning that many of the plant species found within

this ecosystem are highly dynamic and can quickly respond to the precipitation regimes.

These results have a number of implications in weather forecasting, climate change, and

ecosystem studies. First, in regions where snow cover tends to exist throughout a long winter season,

but where total snow amount is limited, accuracies in simulating snow coverage fraction will have a

large impact on calculating surface albedo, and consequently, the energy balance that is crucial in the

modelled climate. In addition, this snow coverage fraction is a dynamic process, particularly in

regions with strong wind and rolling topography. In current land surface schemes coupled with

climate or weather models, the snow coverage and blowing snow calculations are either very simple

or ignored. This may cause large errors in radiation and energy calculations and biases of modelled

climate. The accuracy in snow coverage calculation is also important in ecosystem studies. With its

low thermal conductivity, snow cover largely controls soil microclimate and further influences the

plant physiological processes. For example, it was found that snow removal in a high latitude maple

tree stand caused more than 10ºC decrease in soil temperature. This resulted in sever physiological

responses, including increased canopy dieback and earlier leaf senescence in the following growing

season [Pilon et al., 1994]. Even for a temperate forest in a mild winter, snow removal was found to

cause significant fine root mortality and nutrient loss, although it lowered the root zone temperature

only by 2-4ºC [Groffman, 2001].

Second, climate predictions – and especially temperature predictions – during the seasonal

transitional period (winter-summer of summer- winter) are extremely important because they

determine the dates that snow covers the land surface which, in turn, strongly impact on simulations

of surface albedo. In other words, ecosystems can strongly feedback to the physical climate systems

through albedo due to the changes in climate during the seasonal transitional period. For ecosystem

studies, the timing of snow melt controls the soil warm up and influence the growing season length.

For many ecosystems, the changes in growing season length have a much more profound impact on

166

ecosystem carbon cycles than the changes in plant photosynthesis or respiration brought about by the

predicted climate changes. In addition, the timing and rate of snow melt also influence the

hydrological cycles of the watershed.

Third, surface albedo changes with vegetation dynamics, including the response of vegetation

growth to climate change and its variations. This albedo change is highly dynamic. Because it

happens in summer when the solar radiation can be very high, even a small change in albedo could

significantly feedback to the physical climate system and influence the climate and weather. Chapin

et al. [Chapin et al., 2000] indicated that the net climate-forcing due to about 0.05 difference in

albedo between forest tundra and shrub tundra of northern Alaska is in the order of 5.5 W m-2, which

is comparable to the effect of doubling global atmospheric CO2 concentration [4.4 W m-2, Wuebbles,

1995]. Results from our study show that even for the same grassland ecosystem, climate variations

among different years could cause albedo differences of 0.02 or more due to the impact on

ecosystem function (Figure X. The higher summer albedo in dry years could have significant

negative effect on moisture flux convergence and rainfall, which will positively feedback and cause

drier climates. In most climate models, the albedo of vegetation is parameterized as vegetation-

specific constants. The decoupling of albedo change with climate conditions could thus result in

biases of modeled climate.

In general, surface albedo is controlled by the dynamic processes of ecosystems and could be

significantly influenced by the climate conditions. The impact of climate change and variations on

surface albedo and the feedbacks of the albedo response to the physical climate system need to be

included in the climate model projections.

2.6.4 Implications of climate changes on Canada’s grasslands

2.6.4.1 Potential impacts on agriculture

Climate warming in Canada’s grassland regions is expected to have considerable impacts on

Canadian agriculture. Most regions of the country are expected to experience warmer conditions,

longer frost-free seasons and increased evapotranspiration [Natural Resources Canada, 2002].

However, the actual impacts of these changes on agricultural operations will vary depending on

factors such as precipitation changes (i.e. increases vs decreases in total annual precipitation;

changes in timing of precipitation; changes in fraction of total precipitation falling as snow), soil

167

conditions and land use [Lemmen et al.,

2004]. In many cases, the positive and

negative impacts of climate change would

tend to offset each other. For instance, the

positive impacts of warmer temperatures

and enhanced CO2 on crop growth are

expected to largely offset the negative

impacts of increased moisture stress and

accelerated crop maturation time (Figure X).

In general, northern agricultural regions are

expected to benefit most from longer and

warmer frost-free seasons. Some northern

locations (e.g., Peace River region of

Alberta and British Columbia, and parts of

northern Ontario and Quebec) may also

experience new opportunities for cultivation, although the benefits will likely be restricted to areas

south of latitude 60°N for the next several decades. Although precipitation is expected to increase

over the prairies, this increase is generally not projected to be sufficient to offset increased moisture

losses from warmer temperatures and increased evapotranspiration rates [Lemmen et al., 2004].

Modeling projections by McGinn and Shepherd [2003] suggest that southeastern Saskatchewan and

southern Manitoba may be the most affected by decreases in precipitation.

2.6.4.2 Potential feedbacks to the global carbon cycle

Various modeling studies [e.g. Burke et al., 1991] have attempted to assess the impacts that

warming-induced changes in grassland landscapes will have on the global carbon cycle. Burke et al.,

[1991] used the CENTURY model [Parton et al., 1988] to evaluate the effects of climate warming

on the carbon stocks of North American Great Plains grasslands. Their study indicated that (a)

increased atmospheric temperatures may bring about losses in organic carbon throughout the entire

Great Plains region because of the increased decomposition rates brought about by higher

temperatures, especially in regions where increased temperatures coincided with high precipitation,

Figure X. Potential impacts of climate change on agricultural crops in Canada [after Lemmen et al., 2004].

168

and (b) that these losses of carbon from grassland ecosystems may be partially offset by increases in

ANPP, which were caused by plant responses to increased decomposition and nitrogen

mineralization that enhanced nutrient availability [Schimel et al., 1990]. While Burke et al., [1991]

estimate that these changes could lead to a potential net loss of carbon from grassland soils to the

atmosphere of 0.0014 x 1015 g C yr-1, this is flux is thought to be small relative to historical

cultivation effects 0.018 x 1015 g C yr-1.

The study of Burke et al., [1991] suggests that direct human influences on grassland regions, via

decisions about land management, exert the greatest controls on the regional carbon balance of

grasslands across human time scales. As a result, it is the management response to climate change

that will determine the size of the feedbacks of grasslands to the global carbon cycle. The carbon

stocks of grassland landscapes can be altered by various processes, including conversion to

agriculture, urbanization, desertification, fire, livestock grazing, and the introduction of non-native

species [White et al., 2000]. When grasslands are converted to croplands, the removal of native

vegetation and the subsequent cultivation of the landscape reduces surface cover, destabilizes the

soil, and leads to the loss of soil organic carbon through enhanced decomposition [Sala and Paruelo,

1997; Samson et al., 1998]. The paving of grasslands for urban development, as well as the

desertification or degradation of grasslands, initiates the loss of vegetation cover and enhances soil

erosion, causing a carbon to be lost from the ecosystem [White et al., 2000]. Fires and livestock

grazing in grassland landscapes can release a considerable amount of carbon. During fires, carbon is

directly lost to the atmosphere. During grazing, carbon is lost as food to the livestock, which is then

released as methane gas from the livestock itself. In addition, grazing enhances carbon loss from

grassland landscapes through the trampling and compaction of the soil surface, leading to increases

in soil runoff, increased erosion, and eventually, losses of soil carbon [Sala and Paruelo, 1997]. The

introduction of non-native species may also allow more organic carbon to be released into the

atmosphere. Studies in the Canadian Prairie have shown that the soil beneath certain non-native

species (e.g. Crested Wheatgrass) contains less organic carbon that the soil beneath native species

[Christian and Wilson, 1999].

While the overall feedback affect of climate change on the carbon stocks of grassland ecosystems

depends on the degree to which the above influences accompany climate change, the actual

combined effects of these influences remain far from clear. Understanding the response of grasslands

to these changes requires the knowledge of how regional precipitation patterns will change with

169

climate warming (i.e. increase vs decrease; timing of precipitation), and the synergisms among

climate change and the other human impacts discussed above [Gelbard, 2003]

2.6.5 References

Asner, G.P., Biophysical and biochemical sources of variability in canopy reflectance, Remote Sensing of Environment, 64, 234-253, 1998.

Asrar, G., R.L. Weiser, D.E. Johnson, E.T. Kanemasu, and J.M. Milleen, Distinguishing among tallgrass prairie cover types from measurements of multispectral reflectance, Remote Sensing of Environment, 19, 159-169, 1986.

Betts, A.K., and J.H. Ball, Albedo over the Boreal forest, Journal of Geophysical Research, 102 (D24), 28901-28909, 1997.

Bragg, T.B., The Physical Environment of Great Plains Grasslands, edited by A. Joern, and K.H. Keeler, pp. 49-81, Oxford University Press, New York, USA, 1995.

Brown, L., The Audubon Society Nature Guides. Grasslands, 595 pp., Chanticleer Press, Inc, New York, USA, 1989.

Bryson, R.A., and F.K. Hare, The Climates of North America, 420 pp., Elsevier Scientific Publishing Company, Amsterdam, Holland, 1974.

Burke, I.C., T.G.F. Kittel, W.K. Lauenroth, P. Snook, C.M. Yonker, and W.J. Parton, Regional analysis of the central Great Plains, BioScience, 41, 685-692, 1991.

Canadian Climate Impacts and Scenarios, Online document (http://www.cics.uvic.ca/scenarios/index.cgi/Introduction), (last updated Janurary 2005), 2001.

Catto, N.R., Loess in the Cypress Hills, Alberta, Canada, Canadian Journal of Earth Sciences, 20, 1159–1167, 1983.

Chapin, F.S.I., W. Eugster, J.P. McFadden, A.H. Lynch, and D.A. Walker, Summer differences among arctic ecosystems in regional climate forcing, Journal of Climate, 13, 2002-2010, 2000.

Christian, J.M., and S.D. Wilson, Long-term ecosystem impacts on an introduced grass on the northern Great Plains, Ecology, 80 (7), 2397-2407, 1999.

Coppedge, B.R., D.M. Engle, C.S. Toepfler, and J.H. Shaw, Effects of seasonal fire, bison grazing and climatic variation on tallgrass prairie vegetation, Plant Ecology, 139, 235-246, 1998.

Coupland, R.T., Grassland Ecosystems of the World, Cambridge University Press, Cambridge, UK, 1979.

Coupland, R.T., Mixed Prairie, pp. 151-181, Elsevier, Amsterdam, Holland., 1992.

Datt, B., Remote Sensing of Water Content in Eucalyptus Leaves, Australian Journal of Botany, 47, 909-923, 1999.

170

Davidson, A., and F. Csillag, The influence of vegetation index and spatial resolution on a two-date remote sensing derived relation to C4 species coverage, Remote Sensing of Environment, 75 (1), 138-151, 2001.

Davidson, A., and F. Csillag, A comparison of three approaches for predicting C4 species cover of northern mixed grass prairie, Remote Sensing of Environment, 86 (1), 70-82, 2003.

Davidson, A., S. Wang, and J. Wilmshurst, Remote sensing of grassland-shrubland vegetation water content in the shortwave domain, International Journal of Applied Earth Observation and Geoinformation, in review.

Environment Canada, Climate Change: Overview, Online document (http://www.ec.gc.ca/climate/overview_science-e.html), Last accessed March 2005, 2005.

Epstein, H.E., W.K. Lauenroth, I.C. Burke, and D.P. Coffin, Productivity patterns of C3 and C4 functional types in the US Great Plains, Ecology, 78 (3), 722-731, 1997.

Flato, G.M., G.J. Boer, W.G. Lee, N.A. McFarlane, D. Ramsden, M.C. Reader, and A.J. Weaver, The Canadian Centre for Modelling and Analysis Global Coupled Model and its climate, Climate Dynamics, 16, 451–467, 2000.

Gelbard, J.L., Grasslands, in Buying time: A user's manual for building rtesistance and resilience to climate change in natural systems., edited by L.J. Hansen, J.L. Biringer, and J.R. Hoffman, pp. 1-42, World Wildlife Fund, Berlin, Germany, 2003.

Girard, C.M., Estimation of phenological stages and physiological stages of grasslands from remote sensing data, Vegetatio, 48, 219-226, 1982.

Goetz, S.J., Multi-sensor analysis of NDVI, surface temperature and biophysical variables at a mixed grassland site, International Journal of Remote Sensing, 18 (1), 71-94, 1997.

Goodin, D.G., and G.M. Henebry, A technique for monitoring ecological disturbance in tallgrass prairie using seasonal NDVI trajectories and a discriminant function mixture model, Remote Sensing of Environment, 61 (2), 270-278, 1997.

Goodin, D.G., and G.M. Henebry, Seasonality of finely-resolved spatial structure of NDVI and its component reflectances in tallgrass prairie, International Journal of Remote Sensing, 19 (16), 3213-3220, 1998.

Groffman, P.M., Driscoll, C.T., Fahey, T.J., Hardy, J.P., Fitzhugh, R.D., Tierney, G.L.,, Effects of mild winter freezing on soil nitrogen and carbon dynamics in a northern hardwood forest, Biogeochemistry, 56, 191-213, 2001.

Henebry, G.M., Detecting change in grasslands using measures of spatial dependence with Landsat TM data, Remote Sensing of Environment, 46, 233-234, 1993.

Henebry, G.M., and H. Su, Using landscape trajectories to assess the effects of radiometric rectification, International Journal of Remote Sensing, 14 (12), 2417-2423, 1993.

Hengeveld, H.G., Projections for Canada's climate future: A discussion of recent simulations with the Canadian Global Climate Model, Climate Change Digest Series, CCD 00-01, pp. 27, Environment Canada, Ottawa, ON, Canada, 2000.

171

Ikeda, H., K. Okamoto, and M. Fukuhara, Estimation of aboveground phytomass with a growth model using Landsat TM and climate data, International Journal of Remote Sensing, 20 (11), 2283-2294, 1999.

Jin, Y., C.B. Schaaf, F. Gao, X. Li, A.H. Strahler, W. Lucht, and S. Liang, Consistency of MODIS surface bidirectional reflectance distribution function and albedo retrievals: 1. Algorithm Performance, Journal of Geophysical Research, 108 (D5), 4158, 2003a.

Jin, Y., C.B. Schaaf, C.E. Woodcock, F. Gao, X. Li, A.H. Strahler, W. Lucht, and S. Liang, Consistency of MODIS surface bidirectional reflectance distribution function and albedo retrievals: 2. Validation, Journal of Geophysical Research, 108 (D5), 4159, 2003b.

Kephart, K.D., C.P. West, and D.A. Wedin, Grassland ecosystems and their improvement, edited by R.F. Barnes, D.A. Miller, and C.J. Nelson, pp. 141-153, Iowa State University Press, Ames, Iowa, USA, 1993.

Lauenroth, W.K., O.E. Sala, D.P. Coffin, and T.B. Kirchner, The importance of soil water in the recruitment of bouteloua gracilis in the shortgrass steppe, Ecological Applications, 4 (4), 741-749, 1994.

Lemmen, D.S., F.J. Warren, E. Barrow, R. Schwartz, J. Andrey, B. Mills, and D. Riedel, Climate change impacts and adaptation: A Canadian perspective, pp. 174, Climate Change Impacts and Adaptation Directorate, Natural Resources Canada, Ottawa, Ontario, Canada, 2004.

Lenihan, J., and R. Neilson, Canadian vegetation sensitivity to projected climatic change at three organizational levels, Climate Change, 30, 27-56, 1995.

Loveridge, D.M., and B. Potyondi, From Wood Mountain to the Whitemud: A historical survey of the Grasslands National Park area, National Historic Parks and Sites Branch, Parks Canada, Environment Canada, Ottawa, 1983.

Lucht, W., C.B. Schaaf, and A.H. Strahler, An algorithm for the retrieval of albedo from space using semiempirical BRDF models, IEEE Transactions on Geoscience and Remote Sensing, 38 (2), 977-998, 2000.

Maxwell, B., Responding to global climate change in Canada's arctic. Volume II of the Canada Country study:Climate impacts and adaptation, pp. 82, Environemntal Adaptation Research Group, Atmospheric Environment Service, Environment Canada, 1997.

McGinn, S.M., and A. Shepherd, Impact of climate change scenarios on the agroclimate of the Canadian prairies, Canadian Journal of Soil Science, 83, 623-630, 2003.

Mino, N., G. Saito, and S. Ogawa, Satellite monitoring of changes in improved grassland management, International Journal of Remote Sensing, 19 (3), 439-452, 1998.

Muhs, D.R., and S.R. Wolfe, Sand dunes of the northern Great Plains of Canada and the United States, in Holocene climate and environmental change in the Palliser Triangle: A geoscientific context for evaluating the impacts of climate change on the southern Canadian Prairies, edited by D.S. Lemmen, and R.E. Vance, pp. 183-197, Geological Survey of Canada, Ottawa, Ontario, Canada, 1999.

Natural Resources Canada, The winds of change: Climate change in Prairie Provinces, Online document (http://adaptation.nrcan.gc.ca/posters/reg_en.asp?Region=pr), Climate Change Impacts and Adaptation Adaptation Directorate, Natural Resources Canada, 2002.

172

Ojima, D.S., W.J. Parton, M.B. Coughenor, J.M.O. Scurlock, T.B. Kirchner, T.G.F. Kittel, D.O. Hall, D.S. Schimel, E.G. Moya, T.G. Gilmarov, T.R. Seastedt, A. Kamnalrut, J.I. Kinyamario, S.P. Long, J.-C. Menaut, O.E. Sala, R.J. Scholes, and J.A. Veen, Impact of climate and atmospheric carbon dioxide changes on grasslands of the world, edited by D.O.H. Breymeyer, J.M. Melillo, and G.I. Ågren, pp. 273-311, John Wiley and Sons, 1996.

Parton, Stewart, and Cole, Dynamics of C, N, P and S in grassland soils: a model, Biogeochemistry, 5, 109-131, 1988.

Parton, W.J., M.B. Coughenor, J.M.O. Scurlock, D.S. Ojima, T.G. Gilmanov, R.J. Scholes, D.S. Schimel, T.B. Kirchner, J.-C. Menaut, T.R. Seastedt, E.G. Moya, A. Kamnalrut, J.I. Kinyamario, and D.O. Hall, Global grassland ecosystem modelling: development and test of ecosystem models for grassland systems, edited by D.O. Breymeyer, J.M. Melillo, and G.I. Ågren, pp. 229-269, John Wiley and Sons, 1996.

Paruelo, J.M., and w.k. Lauenroth, Regional patterns of normalized difference vegetation index in North American shrublands and grasslands, Ecology, 76 (6), 1888-1898, 1995.

Peñuelas, J., J. Piñol, R. Ogaya, and I. Filella, Estimation of plant water concentration by the reflectance Water Index WI (R900/R970), International Journal of Remote Sensing, 18 (13), 2869-2875, 1997.

Pickup, G., V.H. Chewings, and D.J. Nelson, Estimating changes in vegetation cover over time in arid rangelands using landsat MSS data, Remote Sensing of Environment, 43, 243-263, 1993.

Pilon, C.E., B. Côté, and J.W. Fyles, Effect of snow removal on leaf water potential, soil moisture, leaf and soil nutrient status and leaf peroxidase activity of sugar maple, Plant and soil, 162, 81-88, 1994.

Risser, P.G., Grasslands, edited by B.F. Chabot, and H.A. Mooney, pp. 233-256, 1985.

Rizzo, B., and E. Wiken, Assessing the sensitivity of Canada’s ecosystems to climatic change, Climate Change, 21, 37-55, 1992.

Sage, R.F., D.A. Wedin, and M. Li, The biogeography of C4 photosynthesis: Patterns and controlling factors, edited by R.F. Sage, and R.K. Monson, pp. 313-356, Academic Press, San Diego, CA, USA, 1999.

Sala, O.E., and J.M. Paruelo, Ecosystem services in grasslands, in Nature's services, societral dependence and natural ecosystems, edited by G.C. Daily, pp. 237-252, Island Press, Washington, DC, USA, 1997.

Samson, F.B., and F.L. Knopf, Prairie conservation in North America, BioScience, 44, 418-421, 1994.

Samson, F.B., F.L. Knopf, and W.R. Ostlie, Grasslands, in Status and Trends of the Nation's Biological Resources, edited by M.J. Mac, P.A. Opler, C.E. Puckett, and P.D. Doran, pp. 437-472., U.S Department of the Interior, U.S. Geological Survey,, Reston, VA, USA., 1998.

Schaaf, C.B., F. Gao, A.H. Strahler, W. Lucht, X. Li, T. Tsang, N.C. Strugnell, X. Zhang, Y. Jin, J.-P. Muller, P. Lewis, M. Barnsley, P. Hobson, M. Disney, G. Roberts, M. Dunderdale, C. Doll, R. d'Entremont, B. Hu, S. Liang, and J.L. Privette, First operational BRDF, albedo nadir reflectance products from MODIS, Remote Sensing of Environment, 83 (1-2), 135-148, 2002.

173

Schimel, D.S., W.J. Parton, G.F. Kittel, S. Ojima, and C.V. Cole, Grassland biogeochemistry: links to atmospheric processes, Climatic Change, 17, 13-25, 1990.

Scott, D., J.R. Malcolm, and C. Lemieux, Climate change and modelled biome representation in Canada’s national park system: implications for system planning and park mandates, Global Ecology and Biogeography, 11 (6), 475-485, 2002.

Shepherd, A., and S.M. McGinn, Assessment of climate change on the Canadian prairies from downscaled GCM data, Atmosphere-Ocean, 41 (4), 301-316, 2003.

Sims, P.L., J.S. Singh, and W.K. Lauenroth, The structure and function of ten North American grasslands, Journal of Ecology, 66, 251-285, 1978.

Thorpe, J., S.A. Wolfe, J. Campbell, J. LeBlanc, and R. Molder, An ecoregion approach for evaluating land use management and climate change adaptation strategies on sand dune areas in the Prairie Provinces, Saskatchewan Research Council, Saskatoon, SK, Saskatchewan, 2001.

Tieszen, L.L., B.C. Reed, N.B. Bliss, B.K. Wylie, and D.D. DeJong, NDVI, C3 and C4 production, and distributions in Great Plains grassland land cover classes, Ecological Applications, 7 (1), 59-78, 1997.

Todd, S.W., R.M. Hoffer, and D.G. Milchunas, Biomass estimation on grazed and ungrazed rangelands using spectral indices, International Journal of Remote Sensing, 19 (3), 427-438, 1998.

Tucker, C.J., W.W. Newcomb, S.O. Los, and S.D. Prince, Mean and inter-year variation of growing-season normalized difference vegetation index for the Sahel 1981-1989, International Journal of Remote Sensing, 12 (6), 1133-1135, 1991.

Walthall, C.L., and E.M. Middleton, Assessing spatial and seasonal variations in grasslands with spectral reflectances from a helicopter platform, Journal of Geophysical Research, 97, 18905-18912, 1992.

Wang, S., Dynamics of land surface albedo for a boreal forest and its simulation, Ecological Modelling, in press.

Wang, S., W. Chen, and J. Cihlar, New calculation methods of diurnal distributions of solar radiation and its interception by canopy over complex terrain, Ecological Modelling, 155, 191-204, 2002.

Wang, S., and A. Davidson, Impact of climate variations on surface albedo of a temperate grassland, Agricultural and Forest Meteorlogy, in review.

Wang, S., A. Davidson, R. Latifovic, and A. Trishchenko, The impact of drought on land surface albedo, Eos Transactions, AGU, 85(17), Jt. Assem. Suppl., Abstract., 2004.

Wanner, W., A.H. Strahler, B. Hu, P. Lewis, J.-P. Muller, X. Li, C.B. Schaaf, and M.J. Barnsley, Global retrieval of bidirectional reflectance and albedo over land from EOS MODIS and MISR data: Theory and algorithm, Journal of Geophysical Research, 102, 17143-17161, 1997.

Weiser, R.L., G. Asrar, G.P. Miller, and E.T. Kanemasu, Assessing grassland biophysical characteristics from spectral measurements, Remote Sensing of Environment, 20, 141-152, 1986.

174

Wheaton, E.E., Prairie dust storms — a neglected hazard, Natural Hazards, 5, 53-65, 1992.

Wheaton, E.E.C., A.K., Dust storms in the Canadian Prairies, International Journal of Climatology, 10, 829-837, 1990.

White, R., S. Murray, and M. Rohweder, Pilot Analysis of Global Ecosystems: Grassland Ecosystems, World Resources Institute, Washington, DS, USA, 2000.

Whittaker, R.H., Communities and Ecosystems, 385 pp., Macmillan Publishing, New York, USA, 1975.

Whittaker, R.H., and G.E. Likens, Carbon in the biota, edited by G.M. Woodwell, and E.V. Pecan, pp. 281-302, National Technical Information Service, Springfield, 1973.

Williams, G.D.V., R.A. Fautley, K.H. Jones, R.B. Stewart, and E.E. Wheaton, Estimating effects of climate change on agriculture in Saskatchewan, Canada, in The Impact of Climatic Variations on Agriculture Volume 1: Assessments in Cool Temperate and Cold Regions, edited by M.L. Parry, T.R. Carter, and N.J. Konijn, pp. 221-382, Kluwer Academic Publishers, Boston, MA, USA, 1988.

Wolfe, S.A., Impact of increased aridity on sand dune activity in the Canadian prairies, Journal of Arid Environments, 36, 421-432, 1997.

Wolfe, S.A., and J. Thorpe, Shifting sands: Climatic change impacts on sand hills in the Canadian Prairies and implications for land use management, Prairie Forum, 30 (1), 123-142, 2005.

Wroe, R.A., S. Smoliak, B.W. Adams, W.D. Williams, and M.L. Anderson, Guide to range condition and stocking rates for Alberta grasslands, Alberta Agriculture, Food and Rural Development, Public Land Management, Edmonton, Alberta, Canada, 1988.

Wuebbles, D.J., Air Pollution and Climate Change, in Composition, Chemistry, and Climate of the Atmosphere, edited by H.B. Singh, pp. 480-518, Van Nostrand Reinhold, New York, NY, USA, 1995.

Yang, L., B.K. Wylie, L.L. Tieszen, and B.C. Reed, An analysis of relationships among climate forcing and time-integrated NDVI of grasslands over the U.S. Northern and Central Great Plains, Remote Sensing of Environment, 65, 25-37, 1998.

175

3 Conclusions