Download field investigations of permafrost and climatic

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PERMAFROST - Seventh International Conference (Proceedings),
Yellowknife (Canada), Collection Nordicana No 55, 1998
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FIELD INVESTIGATIONS OF PERMAFROST AND CLIMATIC CHANGE
IN NORTHWEST NORTH AMERICA
C.R. Burn
Department of Geography, Carleton University, 1125 Colonel By Drive,
Ottawa, ON K1S 5B6 Canada
e-mail: [email protected]
Abstract
Yukon Territory, adjacent portions of Northwest Territories, and Alaska contain a continental range of permafrost conditions. The response of permafrost to climatic change is recorded in the cryostratigraphy of late
Pleistocene and Holocene sediments, with an early Holocene thaw unconformity being a widespread and
prominent feature. More recently, temperature profiles from deep boreholes show an inflection associated with
near-surface warming of 2¡ to 4¡C since the Little Ice Age. Simultaneously, the southern limit of permafrost has
moved northwards. In order to understand the present climate:ground temperature system, an analytical solution has been verified to relate the annual mean ground surface temperature to the annual mean permafrost
surface temperature under equilibrium conditions. Ground surface temperatures have been obtained from air
temperatures using n-factors. The solution assumes that heat transfer in the active layer is only by conduction.
The relations show that the impact on permafrost temperatures of changes in snow cover and soil moisture
conditions may surpass the effect of changes in air temperature per se. Observations from the sporadic permafrost zone indicate the persistence of permafrost despite recent warming. This is due to minimal snow cover
on residual peat landforms, and to latent heat in ice-rich ground. The persistence further complicates interpretation of the response of permafrost to climate change.
Introduction
Permafrost is a geologic manifestation of climate, so
permafrost conditions should change over time.
Instrumental and paleoenvironmental records indicate
that the climate is warming faster in the Arctic than at
lower latitudes of the Northern Hemisphere, and that
the warming has been greater in the 20th century than
in the previous 400 years (Overpeck et al., 1997). The
response of permafrost to climate change, a theme of
this conference, is the focus of several research projects
supported by the International Permafrost Association
(e.g., Brown, 1997; Harris, 1997). In Canada the
Mackenzie Basin Impact Study, a multi-disciplinary
project, recently produced its final report on the potential impact of climate change on the region, and concluded that a principal threat to the landscape was
"accelerated erosion and landslides caused by permafrost thaw .... especially in sloping terrain and the
Beaufort Sea coastal zone" (Cohen, 1997, 297). Air temperature in parts of Mackenzie Basin has risen by 1.5¼C
over the last century (Maxwell, 1997), and, in the popular press, increased geomorphological activity has been
attributed to such warming (e.g., Grescoe, 1997). In contrast, the scientific literature has ascribed past mass
wasting in Mackenzie Valley to site-specific disturbances (Mackay and Matthews, 1973; Harry and
MacInnes, 1988), although the potential for these events
to follow climate warming is acknowledged (Mackay,
1975a).
On the 1967 Permafrost Map of Canada, R.J.E. Brown
(1967) implicitly recognized the importance of climatic
change. Although the primary purpose of the map was
to indicate the spatial extent of permafrost, Brown
chose the Ð5¼C mean annual air temperature isotherm
to separate the continuous and discontinuous zones. He
recognized that, over the long term, a climatic warming
of over 5¼C would be required to degrade permafrost in
the continuous zone. A climatic shift of such magnitude
is not common during an interglacial period, although
smaller fluctuations occur. Therefore, in the continuous
permafrost zone, permafrost is continuous in time as
well as space, and is discontinuous in these dimensions
to the south.
The response of permafrost to climate change is
shown by changes in the ground temperature profile, or
in the depth of the active layer, or both. In this paper,
research on permafrost and climate change will be considered under four themes: (1) historical climate: permafrost relations; (2) cryostratigraphic relations; (3)
relations between near-surface ground temperatures
and present climate; and (4) the impact of climate
change on permafrost distribution. The purpose of the
paper is to review recent progress in these fields, with
emphasis on evidence from northwest Canada and
C.R. Burn
107
The region is on the western, climatically-leading
edge of the continent, but the Wrangell-St. Elias and
Coast Mountains block maritime air masses from entering the region, causing a subarctic, continental climate
conducive to permafrost (Wahl et al., 1987; Burn, 1994).
Enhancement of temperature inversions by cold-air
drainage in the dissected topography of the region
results in the coldest temperatures of the North
American winter being recorded here (Kalkstein et al.,
1990; Burn, 1993), and the presence of discontinuous
permafrost (Heginbottom, 1995). Taylor et al. (1998)
describe the influence of the inversion on permafrost
temperatures in central Mackenzie Valley: valley-bottoms are underlain by permafrost as a result of frigid
winters, while, at high elevations there is little thawing
under cool summer conditions; in between there may
be a permafrost-free zone. In central Yukon, alpine permafrost is found above 1500 m a.s.l., with welldeveloped cryoplanation terraces and patterned
ground (Hughes, 1983).
Unfortunately, these topographic effects cannot be
resolved at the scale of present general circulation models (GCMs), which generate a climate similar to
Figure 1. Permafrost map of Yukon and adjacent Northwest Territories (after
Heginbottom, 1995).
adjacent areas of Alaska. Field data are presented to
illustrate relations discussed in the literature. The
review builds on a summary prepared in 1992 (Burn
and Smith, 1993), and focuses on material published
since then. The review is regionally-restricted, in contrast with the paper presented at the Beijing Permafrost
Conference on this subject (Nelson et al., 1993).
The Yukon and adjacent Northwest
Territories, permafrost, and climate change
A north-south transect across northwest Canada or
adjacent Alaska from the Beaufort Sea to the Pacific
Ocean covers a continental range in permafrost conditions (Figure 1; Smith et al., 1998, this conference). At
Tuktoyaktuk, N.W.T., the mean annual air temperature
(MAAT) is -10.5¼C, while at Whitehorse, Y.T., MAAT is 1.0¡C, and on the coast at Juneau, AK, MAAT is 4.5¼C
(Arctic Environmental and Data Center, 1986;
Environment Canada, 1993). Continuous permafrost
over 600 m thick, and near-surface ground temperatures below Ð8¼C are found near the Beaufort Sea coast
of Alaska, Yukon Territory, and the Mackenzie Delta
area, N.W.T. (Mackay, 1974; Lachenbruch and Marshall,
1986; Judge et al., 1987). In contrast, the mean annual
ground temperature in the sporadic discontinuous permafrost of southern Yukon Territory is above Ð1¼C, and
permafrost is less than 20 m thick (Burn, 1998).
108
Figure 2. Relations between thawing degree-days and distance from the
Beaufort Sea, 1994-96, along a transect from Pelly Island to Inuvik, N.W.T.
(see also Burn, 1997, Figure 12b). Two sites were not occupied in 1994.
The 7th International Permafrost Conference
Scandinavia for the region. In Scandinavia there is relatively little permafrost, and so the GCM results are
unsuitable for investigations of potential climate
change in northwest North America (Stuart and Judge,
1991). Field experiments by SeppŠlŠ (1982) demonstrated the importance of snow cover on permafrost distribution in the discontinuous zone, and this critical variable is poorly represented for the region in GCMs,
because the rainshadow caused by the coastal mountains is not reproduced (Burn, 1994).
In a similar fashion, a steep summer climatic gradient
in the Beaufort Sea coastal zone, due to the presence of
proximal pack ice offshore (Haugen and Brown, 1980;
Zhang et al., 1996a), is not apparent at the scale of present GCMs. However cooler winter temperatures
inland offset this gradient on an annual basis, so that
MAAT changes little with distance from the coast.
Nevertheless, the summer gradient controls the development of vegetation communities, which, in turn,
impact snow accumulation and, hence, near-surface
ground temperature and active-layer development
(Clebsch and Shanks, 1968; Mackay, 1974; Romanovsky
and Osterkamp, 1995; Nelson et al., 1997). Changes in
conditions along the gradient may not be uniform
under future climatic change, as suggested by the varying range in interannual variability of summer climate
along a transect across treeline in the Mackenzie Delta
area (Figure 2; see Burn, 1997), and in the scattered
covariance of air and ground surface temperature series
on the Alaskan coastal plain (Romanovsky and
Osterkamp, 1995). Ecological changes following climate
change, such as northward treeline migration, may
compound ground temperature increases (Gavrilova,
1993; Burn, 1997).
upwards, at a rate less than 2 cm a-1, with heat supplied
by the geothermal flux, so the response of permafrost
thickness is over glacial time scales (Osterkamp and
Gosink, 1991).
By carefully monitoring sites near the north coast of
Alaska between 1983 and 1993, Osterkamp et al. (1994)
detected a cycle of about 10 years in ground temperatures. The derived amplitude at the permafrost surface
decreased inland from 2¼C at the coast. The magnitude
of fluctuation near the coast may represent sensitivity
to maritime effects, particularly sea ice, on air temperature and snow cover, although Osterkamp et al. (1994)
drew attention to the coincidence of the period with
sunspot activity. Subsequently, Osterkamp and
Romanovsky (1996) supplemented these data with
measurements taken between 1986 and 1993 from the
upper 20 m of permafrost, which were consistent with
the original interpretation. However, they were unable
to judge whether the near-surface temperature series
formed part of an overall warming or was the rising
limb of a cyclic fluctuation.
The ground temperature profile in the discontinuous
permafrost zone has also responded to the 20th century
Historical ground temperature:
Permafrost relations
Pioneering work by Lachenbruch and Marshall (1969,
1986) established the use of ground temperature profiles from arctic Alaska to infer past changes in the
annual mean permafrost surface temperature (AMPST).
Warming of permafrost occurs as a result of either
warmer summers and/or winters, or, in particular, winters with thicker or more persistent snow cover (Smith,
1975; Zhang et al., 1997). This means that increases in
permafrost temperatures are not necessarily associated
with increases in MAAT or in the depth of the active
layer, and vice-versa. Osterkamp et al. (1994) report that
in northern Alaska near-surface temperatures have varied by 4¼C over a decade, on the same order as the last
century, but variation in active-layer thickness has
shown little correlation with the ground temperature
fluctuation. In thick permafrost, where AMPST remains
below 0¼C, degradation occurs from the bottom
Figure 3. Temperature profiles in permafrost, summer, 1990 and, 1997, at a
site near Mayo, Y.T. (Site A1 of Burn, 1992). The change in gradient at 18
m depth is associated with a change in lithology from ice-rich glaciolacustrine sediments (above) to sand (below). The annual temperature range at 5
m is 0.25¼C.
C.R. Burn
109
within 5 km of Mayo in July 1997 all indicated recent
cooling of the ground, and none showed the near-surface isothermal zone evident in 1990.
Warming during the last three decades has led to
eradication of some thin permafrost and an apparent
northward displacement of the southern boundary of
the discontinuous zone in northwest Canada. Thawing
of permafrost in peat bogs at the southern margins of
the permafrost zone in Manitoba began at the end of
the Little Ice Age (Thie, 1974), and continues today
(French and Egorov, 1998). Similar permafrost degradation has been reported for northern Alberta (Vitt et al.,
1994), where in 1988, a field survey of permafrost conditions following the route taken in 1962 by Brown
(1964), concluded that the southerly limit of permafrost
had moved 120 km northwards (Kwong and Gan,
1994). Less than 20% of the ground in northern Alberta
is underlain by permafrost (Heginbottom, 1995), and
under such marginal conditions, we must expect thin
permafrost to aggrade and degrade periodically. Zoltai
(1993) demonstrated such cyclic behaviour in the marginal permafrost of peatlands in northwestern Alberta,
due to wild fires.
Figure 4. Snow accumulation at Mayo Airport, Y.T. (a) in two winters of the
1990Õs; (b) mean accumulation in March 1955 - 1997. Data in (b) for 1955 1988 supplied by Atmospheric Environment Service, and for 1989 - 1997
from station weather record at Mayo Airport.
climate warming, but this has received less attention.
For the same increase in heat flux, the temperature of
ÒwarmÓ permafrost may change less than that of ÒcolderÓ permafrost due to latent heat effects (Riseborough,
1990). Several temperature profiles measured at
Norman Wells by Imperial Oil Ltd in the late 1940Õs and
early 1950Õs indicated curvature to depths of 50 m and
near-surface warming of 3¼C (Hemstock, 1953; Mackay,
1975a), following the warmest air temperatures of this
century. Curvature was still apparent in the upper portions of temperature profiles recorded in central
Mackenzie Valley during the early 1970Õs (Kurfurst et
al., 1974), even though air temperatures had declined.
Climate variation in central and southern Yukon is
dominated by fluctuations in winter conditions, with
snow accumulation being positively associated with
winter air temperature (Burn, 1990). In 1990, the temperature profile in permafrost at Mayo (Figure 1) was
consistent with a warming, since the early 1970s, of
1.25¼C in AMPST (Burn, 1992). Subsequently, the
ground has cooled by 0.2¼C or more in the upper 15 m
of permafrost (Figure 3), in association with reduced
snow accumulation (Figure 4). Five temperature profiles through permafrost measured at different sites
110
In all cases discussed above, changes in the temperature profile have been driven by variations in AMPST,
which integrates both summer and winter climate.
AMPST is a thermal variable which only summarizes
variations in hydroclimate implicitly. For explicit evidence of changes in seasonal and moisture regimes, we
must turn to the cryostratigraphy.
Cryostratigraphic relations
Two recent sets of cryostratigraphic observations from
the region relate to climatic conditions during the last
glacial maximum (c. 20,000 yr. BP). The first are observations of sand and ice wedges on and around Richards
Island, which indicate that the region experienced cold,
dry conditions conducive to eolian activity (Murton
and French, 1993; Murton et al., 1997). The second concern data from the Klondike area, where Fraser and
Burn (1997) determined that the ice-rich, unconsolidated silt deposits, which mantle auriferous gravel in valley bottoms, are of Late Wisconsinan age. The absence
of massive icy bodies from the lowermost parts of most
silt sections suggests this area, too, was arid during the
glacial maximum. Data on the cryostratigraphy of these
unconsolidated materials are presented by Kotler and
Burn (1998), who have determined that most of the
ground ice in the silts formed at the very end of the
Wisconsinan, after 12,000 14C years BP.
The principal cryostratigraphic data relevant to our
topic have been observations of truncated ice wedges
associated with an early Holocene thaw unconformity
The 7th International Permafrost Conference
occur due to progressive deepening of the active layer,
and are not tied to any particular climate scenario.
Figure 5 indicates relatively large subsidence for a small
increase in active-layer depth, due to the ice-rich zone
at the top of permafrost (Cheng, 1983; Burn, 1986).
The evidence from cryostratigraphy and ground temperature profiles indicates how permafrost has
responded to historical climate change. We now turn to
the relations between permafrost and climate derived
from present conditions.
Present ground temperature:
Climate relations
Figure 5. Potential terrain subsidence (cm) at two sites on Pelly Island,
N.W.T. The two profiles represent material between the base of the present
active layer and the early Holocene thaw unconformity; the short profile is
from the crest of a slope. The potential subsidence was determined from the
excess ice content of near-surface permafrost obtained by core drilling, using
the method outlined by Mackay (1970, Figure 3). If the excess ice content of
a 20 cm interval is 50%, then if this interval is incorporated in the active
layer, there will be at least 10 cm of subsidence and at most 10 cm of activelayer thickening. Subsidence is given in (a) for an absolute increase in active
layer thickness; and in (b) for a relative increase in active-layer thickness.
in many parts of the region (Mackay, 1975b, 1978, 1992;
Burn et al., 1986; Burn, 1997). The unconformity records
the maximum development of the active layer around
9000 cal. years BP. Usually, segregated ice is abundant
between the unconformity and the base of the present
active layer (e.g., Pollard and French, 1980), complicating estimates of paleoactive-layer thickness.
Burn (1997, Figure 9) presented two profiles from
Richards Island in which the depth of the paleoactive
layer was estimated by removing the excess ice which
postdates the unconformity. The method requires
analysis as outlined by Mackay (1970, Figure 3). The
reconstructed paleoactive-layer depth for the coastal
region of the Canadian western Arctic was about twice
the present thickness. A considerable portion of the
increase in temperature relative to the present regime
generating the deeper summer thaw is from the paleogeography of the region rather than climatic change per
se. The sites examined were further from the coast 8000
to 9000 years ago than they are today, and therefore
benefited from the regional summer temperature gradient (Burn, 1997; see Figure 2).
Similar data on the paleoactive layer from Pelly Island
have been rearranged to estimate the subsidence that
may occur if the present active layer deepens following
climate warming (Figure 5). These data provide an indication of the extent of terrain disturbance that may
Air and ground temperatures are being monitored
simultaneously in the region, in order to determine the
response of ground temperatures to climatic variation,
and to detect the impact of climatic change on permafrost (e.g., Nixon et al., 1995; Zhang et al., 1997). In
Takhini River valley, near Whitehorse, Y.T. (Figure 1),
ground temperatures have been monitored monthly or
more frequently since 1991 (Burn, 1998). At the Takhini
site, which is in a spruce forest, the AMPST is -0.8¼C,
the active layer is 1.4 m thick, the base of permafrost is
at 18.5 m, and the temperature gradient in the permafrost is linear. The site is 2.3¼C cooler than
Whitehorse Airport, but air temperature series from
these locations have a coefficient of determination (r2)
of 0.98, indicating that fluctuation is synchronous and
of proportional magnitude. As with other sites in southern and central Yukon, monthly mean air temperatures at Whitehorse exhibit little variation during summer, in contrast with the winter. As a result, the active
layer depth has remained almost constant during the
period of observation.
Figure 6 indicates monthly mean air temperatures
measured at Whitehorse Airport and ground temperatures at 1.5 m depth from the Takhini valley site. There
is relatively little snow accumulation at the site (maximum 35 cm), because of arid conditions in the valley,
and interception by the forest canopy. Overall, the
ground temperature series has decreased by 0.07¼C a-1,
in response to recent relatively cool winters. A similar
trend (cooling at 0.02¼C a-1) summarizes temperature
variation at 50 cm depth in the active layer. Data collected at irregular intervals, mostly in winter, for 1982 1990 show similarly little temporal trend (Burn, 1998).
Instead, a great variation in ground temperatures
occurs in this area between vegetation units, and permafrost degrades once the spruce forest is cleared.
At the Fifth International Conference on Permafrost,
A.H. Lachenbruch pointed out that assessment of the
response of permafrost to climate change involves moC.R. Burn
111
Figure 6. Ground temperatures at the surface of permafrost, 1.5 m depth, in Takhini River valley, Y.T., and monthly mean air temperatures at Whitehorse
Airport, 50 km to the east, 1991 - 1997. Ground temperature data collected manually from thermistor cables.
delling three environmental systems: the atmospheric
climate; the ground surface and active layer; and permafrost (Lachenbruch et al., 1988). Critical relations are
those between air temperature and ground surface temperature, and between the ground surface temperature
and permafrost surface temperature. Permafrost is a
medium within which energy is transferred dominantly
by conduction, and there are well-known methods for
determining the response of the temperature profile in
permafrost to fluctuations in AMPST (Lachenbruch et
al., 1988). Outstanding problems are the coupling of
atmospheric temperature changes to changes in ground
surface temperature, and transmission of changes in the
ground surface temperature through the active layer to
permafrost. Currently there are several approaches contributing to progress on these aspects.
CONVECTIVE EFFECTS
Of fundamental significance to the energy transfer
through the active layer is the mode of such exchange.
Many formulations of these relations assume that the
heat flow is exclusively by conduction (e.g., Goodrich,
1982; Kane et al., 1991). However, Hinkel and Outcalt
(1994) have demonstrated that convective transfer
occurs within the active layer during snowmelt and
during infiltration of precipitation in summer (Hinkel
et al., 1997). Sudden increases in soil temperature that
cannot be generated by conduction have been recorded,
and similar conclusions have been drawn from spectral
analyses of soil temperature time series (Hinkel and
Outcalt, 1993).
In addition, minimal conduction may also occur during the zero-curtain period of freeze-up. Soil tempera-
112
tures measured at intervals of one hour during this
period exhibit variations which have been attributed to
evaporation, vapour flow, and condensation, driven by
osmotic gradients (Outcalt et al., 1990; Outcalt and
Hinkel, 1996). In some areas of arctic Alaska, the zero
curtain period may last two months or more, so the
potential total transfer may be a considerable portion of
the annual budget. This portion is not estimated precisely, and may differ between soil types. However,
Osterkamp and Romanovsky (1997) found little evidence of non-convective effects in seven yearsÕ data collected at three sites on the North Slope of Alaska, and
concluded that the thermal regime could be modelled
adequately by conductive processes alone.
Convective transport is clearly important in gravel
and coarse sand, because iron-stained gravel, well
below the present active layer, and associated with the
early Holocene thaw unconformity, has been observed
by J.R. Mackay (personal communication, 1996) at several sites in the Mackenzie delta area. Baker and
Osterkamp (1988) have reported direct observations of
convective heat transport in coarse-grained subsea permafrost. Similarly, in organic soils the potential for
vapour flow is significant and demonstrated (Outcalt
and Hinkel, 1996). In fine-grained mineral soil the influence of convective transfer should be much less.
Figure 7 presents ground temperatures measured over
a year, beginning on 1 August 1996, at 1 m depth in the
floor of a thaw slump near Mayo, Y.T. (Figure 1), where
permafrost is degrading. A similar series is also presented from an undisturbed site nearby. The soil at both
The 7th International Permafrost Conference
an equation for the offset (Romanovsky and
Osterkamp, 1995, eq. 12):
Offset =
Figure 7. Daily mean ground temperature series at a disturbed site in a retrogressive thaw slump, and an undisturbed site in permafrost near Mayo,
Y.T., for one year beginning on 1 August 1996. The data are daily means of
observations collected every 4.8 hours from 1 m depth by data logger.
sites is a glaciolacustrine silty clay, and the permafrost
table is now 4.5 m below the surface at the site in the
slump. The active layer at the undisturbed site is about
80 cm thick. There is scant evidence of convective
effects at the undisturbed site, where the sensor is in
permafrost. At the site in the slump, convective transfer
may be interpreted in early May, during snowmelt, and
in late July as a result of rainfall. In the rest of the year,
the smooth temperature transitions suggest that the
heat flow follows a conductive regime. Although the
time interval of data collection (4.8 hours, presented in
Figure 7 as a daily mean temperature) may be insufficient to detect convective activity, the net proportion of
the annual thermal regime attributed to convection
appears small. The result is that the thermal regime
may be modelled adequately as a conductive system.
Similarly, data on long-term (1958-1997) permafrost
degradation collected at a site in Takhini valley burned
by forest fire are consistent with the Stefan solution,
and indicate that, in similar fine-grained sediments, the
thermal regime may be considered conductive (Burn,
1998). Any convective effects are masked by conduction
because the sediments are fine-grained.
THERMAL OFFSET
Analysis of soil temperatures, collected at several sites
on a transect southward from the Alaskan coast at
Prudhoe Bay between 1986 and 1992, has provided the
basis for estimating the "thermal offset" (Romanovsky
and Osterkamp, 1995). The offset is the relation
between annual mean ground surface temperature
(AMGST) and AMPST (Goodrich, 1982; Burn and
Smith, 1988). The soil temperatures were recorded at
the ground surface and at various depths within the
active layer and near-surface permafrost every four
hours. The sites were in fine-grained sediments, with
overlying organic horizons and a high degree of saturation. From these data, Romanovsky and Osterkamp
(1995) were able to determine the range in thermal offset registered at various sites and under different annual climatic regimes. Assuming only conductive transfers
of heat within the active layer, they derived analytically
ö
DDTs æ Kt
- 1÷
ç
P è Kf
ø
[1]
where kf and kt are the thermal conductivities of the
active layer in frozen and thawed states, DDTs is the
annual total thawing degree-days at the ground surface, and P is the period of the temperature cycle,
1 year. The equation was compared with determinations of the thermal offset from a numerical model
(Goodrich, 1982), and with the field data. In both cases
the agreement was excellent, indicating the applicability of a conductive model at this scale in saturated, finegrained soils. The greatest difference between field
observations and equation [1] occurred in the warmest
year at the warmest site. This suggests that while the
equation may be effective for cold permafrost, its application may be more limited in the discontinuous zone,
or under transient conditions. It may be inappropriate
at warmer sites with fine-grained soil to characterize
thermal conductivity as a bimodal function over the
annual cycle. To date a validation of the model against
field data from a region with warm permafrost, e.g., an
AMPST close to -1¼C, has not been published.
Romanovsky and Osterkamp (1995, eq. 13) also
derived the equation presented by Kudryavtsev (1981)
for the AMPST in terms of the thermal regime at the
ground surface:
AMPST =
kt DDTs - k f DDFs
kf P
[2]
where DDFs is the annual total freezing degree-days
at the ground surface. The outstanding value of equation [2] is the succinct expression of a fundamental relation. It identifies critical summary values for the temperature regime at the ground surface (DDFs, DDTs)
and lithology (kf, kt) governing the difference between
AMGST and AMPST. The equation is rapidly applied,
in comparison with numerical simulations of the annual active-layer thermal regime (Smith, 1977; Smith and
Riseborough, 1983). As a result, equation [2] is a valuable tool for regional assessments of the impact of
changes in ground surface temperature on ÒcoldÓ permafrost. For studies of specific sites, numerical simulations are required to provide the details often necessary.
THE N-FACTORS
The data analysed by Romanovsky and Osterkamp
(1995) were combined with records of air temperature
and snow cover characteristics collected simultaneously
C.R. Burn
113
to assess the influence of climate variation on AMPST
(Zhang et al., 1997). While summer ground temperatures and active-layer depth responded to variations in
air temperatures and the duration of the thaw season
(Romanovsky and Osterkamp, 1997), subsurface conditions in winter were relatively insensitive to these variables, but not to characteristics of the snow pack. These
effects are due in part to the inverse relation between
air temperature and snow depth on the North Slope of
Alaska over the period examined, in contrast with central Yukon. On a regional basis, however, the variation
in winter ground temperatures is dominated by the systematic increase in snow accumulation (Zhang et al.,
1996b), due to trapping by microrelief or vegetation,
which increases in height with distance from the coast.
The data from northern Alaska clearly separate air
and ground temperature relations into two seasonal
regimes. These relations have been summarized
through determination of n-factors. The n-factor is
defined as the ratio of the ground surface temperature
index for the thaw (or freezing) season to the air temperature index for the same season, usually expressed
as accumulated thawing (or freezing) degree-days
(Lunardini, 1978,, 1981). This method offers potential
for summarizing microclimatic exchanges within generalized vegetation units. In general, the n-factors are
determined empirically, by collecting air and ground
temperatures simultaneously at sites representative of
ecological units. Various n-factors for natural surfaces
in Mackenzie valley are presented by Taylor (1995),
adding to the inventory of Jorgenson and Kreig (1988)
and Shur and Slavin-Borovskiy (1993). The n-factors are
generally greater in the thawing season than the freezing season and at tundra sites than in the boreal forest
(Taylor, 1995; Smith et al., 1998).
For sites with permafrost, the freezing season n-factor
varies mostly with snow cover characteristics, associated with the vegetation community. As a result, Smith
and Riseborough (1996, eq. 5) presented a modified
form of equation [2] to relate the air temperature regime
to AMPST, using n-factors to convert the air temperature index to that of the ground surface:
kt
× n1 × DDTa - n f × DDFa [3]
kf
AMPST =
P
where nt and nf are the thawing and freezing season
n-factors, and DDTa and DDFa are the thawing and
freezing indices for air temperature. Equation [3] allows
exploration of changes in equilibrium AMPST in
response to climatic change, as summarized by either
fluctuations in DDTa, DDFa, and/or nf.
Burn (1998) calculated the n-factors for sites with
equilibrium and degrading permafrost in Takhini valley, southern Y.T., and noted that while the thaw season
n-factors were consistent with data from other areas,
the n-factor in winter varied with subsurface conditions. The freezing season n-factor is lowered at sites
without permafrost, or with a very deep active layer,
because of the continuing contribution of latent heat
released during frost penetration. During the period
with Òzero curtainÓ, nf may be close to zero. In contrast,
once freeze-up has occurred at sites with permafrost,
the ground surface temperature may readily decline.
The difference in winter soil temperatures between sites
in forested and burned areas, where permafrost is
degrading, is shown in Figure 8. Air temperatures are
cooler by 1.1¼C at the forested site, while the average
daily soil temperature difference is 2.45¼C.
Equation [3] is restricted in consideration of permafrost degradation by the dependence of nf on the
thickness of the active layer, and by the delay in
response of AMPST to climate change caused by energy
exchanges within permafrost (e.g., Riseborough, 1990).
Projecting permafrost distribution after
climate change
Figure 8. Daily mean soil temperatures at 20 cm depth, 1 November 1994 31 March 1995, at forested and burned sites, Takhini River valley, Y.T.
Observations were taken every 4.8 hours by data logger. The least-squares
regression line for the data is TB = 0.43TF - 0.88 (r2 = 0.90), for soil temperatures (¼C) in the forest (TF) and the burned area (TB).
114
In North America three principal approaches have
been adopted towards modelling the impact of climate
change on permafrost distribution. The first, the
Nelson-Outcalt frost index (Nelson and Outcalt, 1987),
is based entirely on climatic statistics, and has been
The 7th International Permafrost Conference
used with GCM output to project permafrost distribution under various climate scenarios (Anisimov and
Nelson, 1996, 1997). The second, the TTOP model
(equation [3], Smith and Riseborough, 1996), explicitly
recognizes the impact of surface and soil conditions on
ground temperatures, and therefore has potential application over smaller areas. The third is the use of calibrated numerical models for site-specific applications
(Romanovsky et al., 1997).
The Nelson-Outcalt frost index is a normalized ratio
of frost penetration to thaw penetration on a regional
basis, using the climatic thawing index and a modified
freezing index, to accommodate snow cover, derived
from climatic data. A ratio of 0.50 represents the equatorward boundary of permafrost, and 0.67 represents
the boundary between continuous and discontinuous
permafrost zones. The model has been verified by comparing projections of the index with extant permafrost
maps (Anisimov and Nelson, 1997). The model is suited
to application at hemispherical scale, and has been used
to forecast changes in the extent of permafrost associated with various scenarios for global climatic warming
(Anisimov and Nelson, 1996, 1997).
Nelson and Outcalt (1987) explicitly indicated that the
model applied only to equilibrium permafrost, and did
not accommodate degrading permafrost. The result is
that the model's utility is limited for consideration of
transient conditions, such as may be expected over the
next century. Nevertheless, results from GCMs have
been used with the model to examine the area where
changes in permafrost distribution may occur as a
result of a changed climate, and a scenario of permafrost distribution during the early Holocene climatic
optimum has been presented (Anisimov and Nelson,
1996). This latter case illustrates well the difficulties
implicit in the technique for considering transient conditions.
The model indicates that, relative to present conditions, permafrost was considerably restricted in spatial
extent during the early Holocene, with the discontinuous and continuous permafrost zones smaller by about
25% and 67% respectively. However, in the discontinuous permafrost of central Yukon there is field evidence
of permafrost persistence during this period from the
x18O concentration in ground ice (Burn et al., 1986;
Kotler and Burn, 1998). Permafrost in Takhini Valley,
only 15 m thick, is presently degrading as a result of
forest fire, but will likely require over 1200 years for its
eradication (Burn, 1998). Furthermore, a large portion
of the permafrost extant during the Little Ice Age in the
peatlands of the western Canadian provinces has been
mapped from aerial photographs taken between 1949
and 1952, a century after climate warming began
(Halsey et al., 1995). The persistence of perennially-
frozen ground is due to elevation of permafrost-cored
landforms in peatlands and the associated reduction in
snow cover (SeppŠlŠ, 1982), and to the high ice content
of frozen peat. These observations led Halsey et al.
(1995) to the counter-intuitive result that permafrost is
more extensive where MAAT is presently between 0¼
and -3.5¼C, than it was in regions of similar climatic
regime 150 years ago.
Anisimov et al. (1997) conceded this point in attempting to estimate active-layer thickness rather than permafrost distribution per se, but suggested that active
layer response to climate change may be rapid. They
used the method presented by Kudryavtsev et al. (1974)
to couple GCM output to forecasts of active-layer
development, which, unfortunately, cannot accommodate variations in the ice content of near-surface permafrost. Persistence of active-layer thickness is provided by the characteristically ice-rich zone at the surface
of permafrost. Where the active layer is relatively deep,
further deepening requires evacuation of latent heat
from melting near-surface ground ice along a gentle
temperature gradient. Maximum active-layer thickness
occurred about 1000 calendar years after the period of
maximum solar insolation in western Arctic Canada
(10,000 cal. years BP), but initiation of thermokarst lakes
coincided with maximum insolation (see Burn, 1997).
The persistence of permafrost was illustrated by
Riseborough and Smith (1993) in simulations of climatic
variability over periods of a millennium. The model
coupled a randomly varying climate based on the
record from Fort Simpson, N.W.T., to the TONE ground
thermal simulator written by L.E. Goodrich (1982, modified after Steven, 1982), via explicit consideration of
the thermal regime in snow. A key result is that while
permafrost may form rapidly during several cold winters, numerous warm years are required to thaw the
same thickness of ground. The development and thawing of excess ice was not simulated, and this would
tend to further stretch the periods with permafrost.
Smith and Riseborough (1996) and Riseborough and
Smith (1998, this conference) have conducted a series of
sensitivity analyses on the critical variables responsible
for the AMPST, as represented by equation [3]. The
analyses have included consideration of the importance
of changes in snow cover and variations in the thermal
properties of ground materials. These data indicate that
changes in precipitation regime will influence the
response of permafrost to climate change, potentially
swamping the effects of changes in air temperature per
se. Increases in soil moisture content tend to reduce
AMPST, while the effect of snow depth is directly related to AMPST (Riseborough and Smith, 1998). In combination, these relations make more explicit the response
of permafrost to climate change.
C.R. Burn
115
Finally, for specific sites, calibrated numerical simulations have continued to provide effective models of the
ground thermal response to climatic change
(Lachenbruch et al., 1988; Burn, 1992; Zhang and
Osterkamp, 1993). Such analyses of ground temperature profiles has allowed reconstruction of the late
Pleistocene and Holocene environmental history of
Mackenzie Delta area (Taylor et al., 1996). Changes in
surface temperatures due to the glacial/interglacial
transition, submergence during post-glacial sea level
rise, and emergence during delta progradation are
amenable to modelling as step functions, whose magnitudes can be estimated independently. The general
environmental history of the region is known
(Rampton, 1988), so Taylor et al. (1996) were able to use
the temperature profiles to test specific hypotheses
about permafrost evolution in the area, and determine
times of emergence and submergence for various sites.
Conclusion
The paper has attempted to summarize research on
permafrost and climatic change in northwest Canada
and Alaska, emphasizing insights gained from field and
theoretical studies. Within the region, the ground temperature profile in permafrost shows the impact of climatic warming over the last 30 to 150 years, and cryostratigraphy records the effect of the warmest climatic
period of the Holocene. Research on heat transfer has
provided a model for the translation of the thermal
regime at the ground surface into the temperature at
the surface of permafrost, while n-factors are used to
obtain ground surface temperatures from the air temperature. Field evidence from the discontinuous permafrost zone indicates the considerable persistence of
permafrost following climate change, due to latent heat
contained in ground ice. GCMs indicate that future climate change at these latitudes may be most apparent in
winter, and will therefore affect both nf and DDFa. The
active layer will respond to such effects, but its most
rapid response will still be to changes in surface conditions such as those following forest fire. At this point,
then, some suggestions are offered on outstanding
problems and future work.
From a practical perspective, it is essential that various monitoring programs, which have emphasized this
region, continue to collect data on a consistent basis. In
particular, the value of the CALM program (Brown,
1997), which includes a transect of Mackenzie Valley
(Nixon and Taylor, 1998), and key sites in Alaska
(Nelson et al., 1997), increases as time passes, the record
is extended, and trends, cycles and unusual events are
recognized (Burt, 1994). Similarly, the value of near-surface ground temperature monitoring increases with
time, but the significance of such records can only be
evaluated if they are continuous and of quality. Recent
116
changes in emphasis for Federal Departments in
Canada mean that such programs become difficult to
continue. Instead, recognition of the potential for spatially extensive investigations using information technology is growing (Nelson et al., 1997). It would be
dangerous, though attractive, to assume that computer
modelling may be substituted for field investigations.
Within such programs, however, the issue raised by
Smith and Riseborough (1983) of the impact of microclimatic modulation on the response of permafrost to climate change remains outstanding.
The model of air temperature - permafrost relations
developed by Romanovsky and Osterkamp (1995) has
been validated for cold permafrost and may be widely
applied to estimate equilibrium conditions in such terrain, if the n-factors are known. The variation in n-factor with snow conditions is not well described, but
appears to change abruptly across treeline (Smith et al.,
1998). Within vegetation units there is, as yet, little
assessment of the variation in n-factor from site to site,
or from year to year at the same site. The model has yet
to be validated for warm permafrost in fine-grained
soil, and is formulated for equilibrium conditions. The
transient response of permafrost to climate change is
not easily estimated because of the number of compounding variables set in a context of a naturally varying climatic system. At a site scale, calibrated numerical
models may provide precise predictions of the thermal
regime of the active layer and ÒwarmÓ permafrost, even
under transient conditions. Efficient extrapolation of
these results in a regional context is a significant challenge, one which I doubt will be overcome without continuing conscientious efforts from field workers.
Acknowledgments
The research program has been supported by the
National Sciences and Engineering Research Council of
Canada, the Polar Continental Shelf Project (PCSP) and
the Geological Survey of Canada, Natural Resources
Canada, the Inuvik Research Centre of Aurora College,
the Northern Research Institute of Yukon College, the
Atmospheric Environment Service, Environment
Canada, and the Northern Affairs Program of Indian
Affairs and Northern Development Canada. Assistance
from many people in the Yukon and Mackenzie delta
area, particularly Jim and Shann Carmichael of Mayo,
Scott Smith of Whitehorse, and Les Kutny and Alan
Fehr of Inuvik is acknowledged with gratitude. I thank
J.R. Mackay and M.W. Smith for constant stimulation
and encouragement, and Joan Ramsay Burn for her
support of these endeavours. J. Brown, H.M. French,
J.R. Mackay, T.E. Osterkamp, D.W. Riseborough and
M.W. Smith provided helpful comments on the manuscript. PCSP contribution 00498.
The 7th International Permafrost Conference
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