Chapter 3

Groundwater

Subsurface water is found in saturated and non-saturated zones in the ground.

Groundwater is water that occurs in the zone of continuous saturation. Depending

on its origin, groundwater can be meteoric water derived from the atmosphere

through precipitation, juvenile water produced from within the Earth and brought to

the surface by geothermal or volcanic activities, or connate water which is very

ancient water trapped in sedimentary rocks when they were deposited.

The preferred geologic formation for groundwater occurrence is an aquifer, a

stratum that can store and transmit more water than its adjacent layers. An uncon-

fined aquifer is not confined by an overlying layer of impermeable earth material,

and the top of the saturated zone in the unconfined aquifer is marked by the water

table which is under atmospheric pressure. Above the water table is the zone of

aeration which is not saturated but retains varying amounts of soil moisture. Unlike

unconfined aquifers, a confined aquifer is under pressure imposed by both the

atmosphere and by its overlying confining strata.

In an unfrozen state, the richness of groundwater resources is controlled by the

nature of the aquifer. Sedimentary beds of sandstone and coarser formations offer

productive bedrock aquifers, but shale and siltstone have low water-yielding capac-

ity unless they are heavily fractured. Groundwater in crystalline rocks is restricted

to the fissures and weathered zones, and also to solution channels and caverns in

carbonate rocks. Unconsolidated surficial sediments have varying porosity or

volume of openings and pores. Fine-grained materials like silt and clay have high

porosity to store water but they have low yield because water in the small pores is

held against drainage. Coarse sediments with large interstitial openings can store

and yield much groundwater. Thus, alluvial and coarse-grained glacio-fluvial

deposits make good aquifers.

Freezing alters the intrinsic behavior of aquifers as varying amount of ground

ice occupies the interstitial voids to reduce the permeability of the water storage

media.

M.-k. Woo, Permafrost Hydrology,DOI 10.1007/978-3-642-23462-0_3, # Springer-Verlag Berlin Heidelberg 2012

73

3.1 Groundwater Occurrence in Permafrost

The study of groundwater in permafrost terrains has been prompted by the search

for water supply, by problems associated with groundwater in mining and in

construction of buildings, highways, railways, airfields and pipelines, and by

encounters with ice features in the course of permafrost and geological mapping.

Williams (1965) provided a bibliographic documentation of groundwater

investigations in Russian and northern North America up to 1960.

Permafrost may be dry if it contains little water or ice. Permafrost may also be

saturated with unfrozen water if ice formation is prevented by freezing point

depression as, for example, in cases where the groundwater is highly saline. In

most situations, unfrozen water and ice co-exist in permafrost and blockage of the

pores and fissures by ground ice greatly reduces the permeability. Thus, permafrost

is generally considered to behave like an aquiclude (which is poor in the retention

and transmission of liquid water) or an aquitard (which though relatively impervi-

ous, can influence the hydraulics of the non-frozen zone). With permafrost being a

medium of limited permeability, groundwater is normally found in the thawed

zones.

Active groundwater circulation can occur above, within and beneath the perma-

frost, known respectively as suprapermafrost, intrapermafrost and subpermafrost

groundwater (Tolstikhin and Tolstikhin 1976, Williams and Waller 1966).

Figure 3.1 is a conceptualization of the occurrences of groundwater in permafrost

terrain.

subpermafrost groundwater

talik

permafrost

seasonal frost

intrapermafrostgroundwater

lake

icing

frost mound

lake ice

ice

Fig. 3.1 Occurrence of groundwater in permafrost: suprapermafrost (on top of the permafrost),

intrapermafrost (within permafrost) and subpermafrost (below permafrost) groundwater. Also

shown are seasonal frost in the active layer, icing and frost mounds associated with groundwater

flow

74 3 Groundwater

3.1.1 Suprapermafrost Groundwater

Suprapermafrost groundwater occurs mostly in the active layer. As this layer is

subject to seasonal freeze and thaw, water is stored in solid form (ground ice) or

extruded above ground in the cold season; and when thawed, water is released from

seasonal ice storage. Suprapermafrost groundwater is also found in the talik above

the permafrost. Such a closed talik can be maintained if freezing does not reach the

permafrost table, as in the case of water bodies too deep to be frozen to their beds.

There, water remains unfrozen throughout the year.

The water table marks the upper limit of the saturated zone in the thaw season

(Fig. 3.2) Groundwater is normally recharged by meteoric water, with snowmelt

and rain being the main water sources, as well as by glacier meltwater in glacierized

areas. Losses of suprapermafrost groundwater to evapotranspiration and to dis-

charge take place through the active layer. Exceptions are the closed taliks found

below deep water bodies where groundwater is recharged or discharged directly

through the overlying water bodies.

At the beginning of winter when descent of seasonal frost reaches the water

table, the active layer becomes a confined aquifer squeezed in between the imper-

vious frozen cap and the permafrost. Sloan and van Everdingen (1988) noted that

the encroaching frost can produce a ‘quick’ condition in the soil as upward seepage

force of the water exceeds a critical value at which there is no grain-to-grain

pressure. The soil then acquires the properties of a fluid. Such a condition can

also be created when surface pressure is exerted on a thawed active layer saturated

Fig. 3.2 A pit in the active layer revealing the suprapermafrost water table, Teller, Alaska

3.1 Groundwater Occurrence in Permafrost 75

at the base. Figure 3.3 illustrates a similar effect when compression is exerted on the

soil in summer. As some children in Nome, Alaska, jumped up and down on an

active layer thawed to 1.2 m and saturated near the frost table, the compression

induced liquefaction of the diamictic soil. The ground acquired a jelly-like consis-

tency, rolling and rocking to the delight of the children.

3.1.2 Intrapermafrost Groundwater

Intrapermafrost groundwater occurs in taliks within the permafrost. This ground-

water is perennially unfrozen though the water can be below 0�C if it is high in

concentrations of dissolved solids. The unfrozen zone can be an open talik,

extending across the entire thickness of permafrost (van Everdingen 1990), that

provides passage for intrapermafrost groundwater to connect with water above and

below the permafrost. Faults, solution conduits and permeable geologic units

facilitate the development of open taliks while the flow of heated groundwater or

mineralized water maintains hydrothermal and hydrochemical taliks (Fig. 3.4). The

unfrozen zone can also be an isolated talik which may be due to the localized

presence of highly mineralized water, or it may be transitional in nature as

aggrading permafrost below a drained lake encroaches upon the closed talik to

cut off flow connection with its frozen surroundings.

Fig. 3.3 Compression exerted by jumping on the active layer (thawed to 1.3 m and saturated at the

base) leading to liquefaction of the soil, Nome, Alaska

76 3 Groundwater

3.1.3 Subpermafrost Groundwater

Found below permafrost, subpermafrost groundwater is above 0�C but being

confined by the impervious permafrost, it is under hydrostatic pressure and has

artesian conditions (Williams and Waller 1966). Where the permafrost is thin, this

groundwater can be stored in unconsolidated materials or bedrocks; but where the

permafrost is thick, it is limited to aquifers in sedimentary beds or in igneous rocks

sufficiently fractured to store water. Subpermafrost groundwater can be relatively

fresh if it occurs in karst areas that permit fast flow, or in non-soluble bedrocks.

More often, it has high concentrations of dissolved solids, depending on the

composition of the aquifer and the residence time of water in the aquifer. Ground-

water discharged as saline springs at temperature below 0�C would likely have

come from the cryopegs which are actually parts of the permafrost though the high

mineral concentration prevents it from freezing. Subpermafrost groundwater can be

very old. For example, using environmental isotopes, Michel and Fritz (1978)

found some ice within the permafrost in the Mackenzie Valley to be indicative of

water recharged during the glacial period 7,000–10,000 years ago. This groundwa-

ter is post-glacial in age.

Recharge of subpermafrost groundwater in discontinuous permafrost regions

can be directly from adjacent non-permafrost uplands or through supra- and

intrapermafrost connections. In continuous permafrost areas, recharge is restricted

but can be enhanced through fractures or solution conduits such as sinkholes.

drained lakebasin

salinespring

thermalspring

freshspring

basal cryopegT < 0°C

hydrochemical talikT < 0°C

hydrothermal talikT > 0°C

transientisolated

talikT > 0°C

isolated cryopegT < 0°C

lateral talikT > 0°C

seasonal hydrothermal

talikclosed talik

T > 0°C

creek orlake

subpermafrost aquifer T > 0°C

permafrostnon-permafrostfrozen

unfrozenunfrozen

seasonally frozen

Fig. 3.4 Formation of taliks and springs associated with groundwater flow in permafrost

(Simplified after van Everdingen 1990)

3.1 Groundwater Occurrence in Permafrost 77

3.2 Groundwater Recharge and Circulation

3.2.1 Recharge

Water entry into the suprapermafrost reservoir is favored where there is an organic

cover of high porosity, where the unconsolidated deposits are coarse, such as sand

and gravel, or where large conduits in bedrock are exposed at the ground surface.

Recharge is limited or curtailed when the freezing of the active layer hinders

infiltration (Sect. 5.2.2) or where flow passages are blocked by ice and sediments.

Furthermore, fissures in bedrocks diminish in number and size with depth and this

restricts downward percolation. However, localized recharge of intrapermafrost or

subpermafrost groundwater is facilitated by deep fractures, lava tunnels or passages

through soluble rocks such as gypsum and carbonates. Van Everdingen (1987)

measured water entry into sinkholes in permafrost west of Great Bear Lake, NWT,

exceeding 1 m3 s–1 in the snowmelt period. One remarkable example of recharge

was given by Brook (1983), who examined the hydrology of three large poljes

(large, flat floored enclosed depressions in karst terrain) in the Nahanni Plateau,

NWT. The karst region there has discontinuous permafrost, as revealed by the ice

and frost found in the tunnels and caves (Fig. 3.5). The poljes studied are 1.4, 0.7

and 1.4 km long and they were dry (Fig. 3.6a) before heavy summer rain events of

19–31 July 1972 deposited 224 mm and filled them to maximum water depths of

8.5, 25 and 8 m respectively. These temporary lakes perched for weeks or months

Fig. 3.5 Cave lined with ice crystals, discontinuous permafrost in carbonate terrain, Nahanni,

NWT

78 3 Groundwater

until the blockages, presumably ice, were cleared from their subterranean passages

in the carbonate rocks. The recharge and drainage of the poljes are repeated

episodes (Fig. 3.6b shows that the Third Polje was again flooded in 1976). The

drainage of these enclosed lakes and other smaller depressions in the vicinity yields

substantial recharges periodically to the subpermafrost groundwater reservoir.

Fig. 3.6 A polje in discontinuous permafrost terrain, North Nahanni basin, NWT, (a) dry before

rain and (b) flooded after a summer rain event in 1972 (Photo courtesy of G.A. Brook)

3.2 Groundwater Recharge and Circulation 79

Lakes and rivers exchange flow with aquifers. Groundwater may feed these

surface water bodies or may receive water from them. Deep lakes with taliks linked

to intra- and subpermafrost zones can recharge the deep aquifers. Along various

reaches of a river there are flow exchanges with the thawed banks and beds during

the summer and autumn seasons. Sokolov (1991) reported that there is intense

exchange between groundwater and river channels during the summer–autumn

low flow period along 90–95% of the Baikal-Amur-Mainline railway route. He

identified three types of river reaches. Groundwater enters the rivers where they

cross fault zones, but in other river reaches water infiltrates the highly jointed

bedrocks to recharge the groundwater. Finally, there are reaches where groundwater

is discharged during rain events but is recharged by river water in times of low flow.

3.2.2 Groundwater Movement

The direction and rate of groundwater flow in permafrost terrain follows the same

physical principles as in non-permafrost areas. Darcy’s Law forms the essential

basis for calculating the velocity of flow in porous media, with hydraulic gradient

and hydraulic conductivity being the prime considerations:

vg ¼ �Kwðdh=dxÞ (3.1)

Here, vg is flow velocity (m s–1), dh/dx is hydraulic gradient and Kw is saturated

hydraulic conductivity (m s–1). Typical hydraulic conductivity values for unfrozen

earth materials are given in Table 1.1. Kw is related to porosity, pore size and soil

cracks, and in frozen soil, it is also much affected by ice occupying pores and found

as lenses. Coldness of soil and chemical concentration of soil water are important in

affecting the freezing process, hence the formation of ice that would block or retard

groundwater flow. With ice in its soils and fissures, permafrost generally plays the

role of an aquiclude or an aquitard embedded in the flow system.

Flow of the suprapermafrost groundwater is seasonal. In continuous permafrost

zone, the flow is confined to the thawed portion of the active layer. In discontinuous

permafrost terrain, there may be exchanges of flow between the water of the

suprapermafrost and the intrapermfrost zones, or with water in non-permafrost

areas. The amount of shallow groundwater in the suprapermafrost zone is small

compared with the deep groundwater that occurs below permafrost. However, the

circulation of shallow groundwater is important to the hydrologic cycle because it is

strongly linked to the processes of infiltration which enters the ground, evaporation

that withdraws water from the active layer, redistribution of water in the soil, and

exfiltration (issuance of water from the ground) in support of runoff or the forma-

tion of ice features in winter. Suprapermafrost groundwater flow is driven by the

hydraulic head which is largely dictated by the topography in the thawed season.

Thus, shallow groundwater tends to drain towards local depressions and bases of

80 3 Groundwater

slopes. When freezing creates an impermeable layer that descends continuously

into the thawed portion of the soil, there is an increase in the pore water pressure

that can lead to liquid water migration to the freezing front in frost susceptible soils

(Sect. 2.6.3).

There is a limited amount of data on the movement of deep groundwater within

and below the permafrost. Mines often provide an opportunity to investigate the

circulation of water, notably within bedrocks. The coal mines of Ny-Alesund in

Svalbard yield information on subpermafrost groundwater flow (Haldersen et al.

1996). A generalized stratigraphic cross section is given in Fig. 3.7. Permafrost of

thickness between 100 and 150 m provides an impermeable confining layer. The

regional groundwater is recharged along the bases of the glaciers which are at

pressure melting point. As the sedimentary rocks in the area have rather low

permeability, fractures are the main conduits for groundwater flow. In the natural

state, groundwater is discharged through springs in karst beds and fractures in

siliceous rocks. Mines create artificial openings for water storage so that after

closure, the cavities are filled with water or ice. Groundwater can flow from the

mines and seals the mine tunnels upon freezing, as shown in the abandoned mine

entrances in Longyearbyen (Fig. 3.8). Within the mining area of Ny-Alesund is

Lake Tvillingvatnet which has artesian water gushing from a depth of 18 m. When a

mine shaft was extended 30 m beneath the lake, permafrost was encountered along

the length of the whole mine, indicating that the lake is not connected to the

subpermafrost zone. The lake is therefore likely to be fed by a transient talik

through the fault.

An example of subpermafrost groundwater circulation in crystalline rocks is

from the gold mines in Yellowknife, NWT. There is large mineralized groundwater

flow at depths exceeding 700 m while the base of the permafrost is just over 100 m

deep (Brown 1970). The flow is concentrated along joints and faults in the Precam-

brian bedrock.

sandstone

shale

glauconitesandstone

limestonedolomite

conglomerate,sandstone

fault

Ny Ålesund

Zeppelin-fjellet

VestreLovénbreen

Ester mine

coal seam and mine permafrost provides a confining layer

confinedsubpermafrost aquifer

phyllite

Fig. 3.7 Circulation and discharge of subpermafrost groundwater in Ny-Alesund, Svalbard, with

recharge along the base of glaciers and groundwater flow in permeable sedimentary rocks and

fractures (Simplified after Haldersen et al. 1996)

3.2 Groundwater Recharge and Circulation 81

Numerical modeling is one approach that can be used to gain understanding of

the flow system. One hypothetical surface water body underlain by a discontinuous

confining bed was used to simulate the hydraulic head and the pattern of ground-

water circulation (Woo and Winter 1993). The confining bed represents permafrost

with low permeability and the model considers a permeable unit (U1) across the

basin, with a discontinuous confining bed (U2) of low permeability to represent

permafrost, and an aquifer below with the highest permeability (U3). The

simulations only consider the situation of downward hydraulic gradient from the

water body to the aquifer below the permafrost. In the three separate simulations,

either the width of the window or gap in the confining bed or the position of the

window relative to the surface water body is changed. Based on the simulated

results (Fig. 3.9), it is deduced that the water body has the largest probability of

interacting with the regional groundwater flow system if a window in the perma-

frost lies directly beneath the water body (a and b in the figure). However, a shift in

the position of the window (comparing a and c) has a greater effect on the flow field

than the size of the window (comparing a and b) in affecting the flow field. If an

upward gradient exists the pattern will be different, with a significant discharge

from the deep aquifer to the water body.

Permafrost degradation due to natural or human-induced environmental changes

can alter the groundwater flow paths, especially in discontinuous permafrost areas.

Applying a two-dimensional model of coupled heat and water flow, Bense et al.

(2009) simulated the formation of suprapermafrost talik under climate warming,

and suggested also an accelerated increase in groundwater discharge when

remnants of permafrost at depth are thawed. While the link between groundwater

circulation and discharge has not been definitively ascertained through field

Fig. 3.8 Coal mine adit in

Longyearbyen, Svalbard,

providing conduit for

groundwater flow in

permafrost. Emerging water

frozen as ice now seals the

mine entrance

82 3 Groundwater

U1

U3

U2

240230

a

lake

line of equal hydraulic head,interval is variable

stagnation point

direction of groundwater flow

lake sedimentsKw = 1x10–3

unit U1Kw =1

confining bed, U2Kw = 1x10–3

aquifer, U3Kw = 1x103

Kw = hydraulic conductivity

Anisotropy of all units is 103

b

U1

U3

U2

lake

225

225

230240

c

225

240230

lake

U3

U2

U1230.8

250

210

235

235 240230

230227

225

220

215

245

230

230

233

240235

245

250

227

225220

215210

225220215

210230

231

231

235 240

245

250

235

235

230.8

Fig. 3.9 Simulated groundwater flow in discontinuous permafrost, using aquicludes of extremely

low hydraulic conductivity to represent permafrost (After Woo and Winter 1993)

3.2 Groundwater Recharge and Circulation 83

investigations, the simulation offers a plausible connection among climate

warming, change in groundwater movement and increase in winter low flow

(which necessarily comes from groundwater source in permafrost terrain) inferred

for northern Eurasian and northwestern North American rivers in recent decades

(Smith et al. 2007, St Jacques and Sauchyn 2009, Walvoord and Striegl 2007).

3.3 Groundwater Discharge

Groundwater is discharged as springs or seeps that rise above ground. It supports

river flow and supplies water to tundra ponds and lakes. Van Everdingen (1990) and

other researchers suggested several field indicators as evidence of groundwater

discharge:

• Presence of open water sections in ice-covered rivers, maintained by large

discharge of springs

• Discharge of highly mineralized water, formation of terraced mounds, cascade

structures or salt crusts from mineral precipitates

• Production of icing which is layered ice formed by the freezing of water as it

emerges from below ground (Sect. 3.4)

• Formation of frost mounds

• Local change in vegetation.

Highly mineralized groundwater is often tinted. The water that flows into

Engineer Creek along Dempster Highway, NWT (Fig. 3.10) is tan colored and

precipitates with reddish-brown iron oxides. In other areas, white precipitates can

be left by sulfurous springs and thermal springs have whitish deposits around their

vents. Carbonate rich water deposits whitish or tan-colored travertine, sometimes

forming terraces and pools with water color that may be milky white or greenish

(reduced iron or carbonate). Figure 3.11 shows icing formed on a slope in Manners

Creek basin, NWT. The icing is yellowish to light brown, being tinted by organic

matter in the soil water. In Fig. 3.12, frost mounds on a slope of Sukakpak Mountain

along the Dalton Highway of Alaska has domed up the ground and tilted the trees

that grow on some mounds. In the valley of Pilgrim in western Alaska with a

number of hot springs, the micro-climate is altered and the vegetation is enriched.

Plant diversification locally changes the shrub-tundra into deciduous woodlands of

aspen and birch (Fig. 3.13).

3.3.1 Seeps

Seepage is the slow oozing out of water from the ground, through the voids and

cracks in unconsolidated materials, or from fissures and bedding planes in rocks.

Seepage also occurs underwater, and discharges directly into streams, lakes or the

84 3 Groundwater

Fig. 3.10 Iron-rich

groundwater emerging at

Engineer Creek, NWT,

showing tan colored water

and precipitates with reddish-

brown iron oxides

Fig. 3.11 Icing with calcareous deposit, Manners Creek basin, NWT. Leaves and twigs fallen on

the icing have locally enhanced ice-melt

3.3 Groundwater Discharge 85

Fig. 3.12 Frost mounds on a slope of Sukakpak Mountain, doming the ground and tilting the trees

that grow on some mounds

Fig. 3.13 A hot spring in permafrost area, Pilgrim, Alaska, with water flowing under artesian

pressure at 80�C. The springs modify local microclimate and trees grow amid the surrounding

tundra

86 3 Groundwater

sea. Seeps are usually intermittent, spreading over a poorly defined area that

expands and contracts as the water supply changes. Many seeps are supported by

suprapermafrost groundwater and seepage occurs where the water table intersects

the ground surface. Favorable sites are slope concavities, including local

depressions and the break of slope at the bottom of hills (Fig. 3.14). Seepage

from a saturated active layer can cause formation of soil pipes that weakens the

thawed active layer. Detachment and downslope sliding of the saturated soil is thus

facilitated, as illustrated by the loss of a top soil layer in Vendom Fiord area

(Fig. 3.15). Saturated soils or puddles are found at these seeps but they may dry

out after the seepage ceases. Seeps are also one of the water sources for wetlands, as

shown in Fig. 8.37.

3.3.2 Springs

Springs are discrete discharge points where groundwater is issued either directly at

the ground surface or below water bodies such as lakes, rivers and the sea. The flow

is normally higher than seepage though at low discharge, springs and seeps become

indistinguishable. Springs may be intermittent or perennial. Those that are seasonal

are usually fed by suprapermafrost groundwater and lack a steady water supply. On

the other hand, perennial springs are supported by deeper water sources, with flows

maintained continuously along unfrozen conduits in taliks that connect them to

intra- or subpermafrost reservoirs.

Fig. 3.14 Seepage of groundwater at the base of a concave slope (brown patch with puddles), east

of Teller

3.3 Groundwater Discharge 87

Water discharged from springs may be fresh, saline or thermal in nature

(Fig. 3.4). Flow rate varies from less than 0.1 to several hundred L s�1 and can

come from a single source or multiple outlets. Sloan and van Everdingen (1988)

suggest that most springs supported by subpermafrost groundwater have a water

temperature greater than 10�C, or have a total dissolved solid concentration

exceeding 1 g L�1 (1,000 mg L�1), or discharge at rates exceeding 5 L s�1.

Dissolution of bedrock creates a network of tunnels and caves that serve as passages

for deep groundwater to reach the surface to discharge as cold springs.

Springs are much more common in discontinuous than in continuous permafrost

areas as there are greater opportunities for taliks to provide conduits for intra- and

subpermafrost groundwater to reach the surface. The thicker active layer also has

the potential to store and deliver more suprapermafrost groundwater than the thin

thaw zone above continuous permafrost. The discontinuous permafrost environ-

ment at North Fork, Yukon, offers an environment for the emergence of several

groups of springs (Pollard and French 1984). Figure 3.16 shows one such cluster in

which the springs are aligned along the base of a steep slope and in a gully that

crosses the Dempster Highway. Temperature of the springs, measured in March and

August of 3 years, was relatively even at 0.5–1.0�C. Total dissolved solid was in the463–491 mg L�1 range and the electrical conductivity was 540–600 mS cm�1.

Groundwater discharge from the springs is also associated with the formation of

icing in the winter. The tundra pond situated at the mouth of the gully is apparently

fed by springs and seeps along the gully.

Fig. 3.15 Seepage from saturated active layer weakens the thawed soil, leading to detachment and

local loss of top soil layer, Vendom Fiord, Ellesmere Island

88 3 Groundwater

Strongly mineralized water can move through thick permafrost to discharge as

springs. At Expedition Fiord on west-central Axel Heiberg Island where continuous

permafrost thickness exceeds 400 m, several clusters of mineralized springs are

found at locations where diapirs (intrusion of relatively light rock material through

overlying rock) of gypsum are exposed. The groundwater that feeds the springs may

be of subpermafrost origin, diluted or otherwise by lakewater or basal meltwater

from glaciers that circulate through taliks (Pollard et al. 1999). The spring water has

high concentrations of dissolved solids (76–78 g L�1; electrical conductivity of

~100 mS cm-1), notably Na, Ca, SO4 and Cl, that can depress the freezing point by

N

0 50 100

meters

BM 1188.9 m

1185

1187

1189 11

91

1193

1199

1197

1195

1183

1190

1188

1186

Dem

pste

r H

ighw

ay

river channel and icing

tundra pond

seasonal frost mound

spring outlet, seepage line

thermistor cable location

piezometer location

break of slope

1981 icing limit

1982 icing limit

contour line, 1 m interval

Fig. 3.16 Groups of springs along the base of a steep slope and in a gully, Dempster Highway,

NWT (Simplified after Pollard and French 1984)

3.3 Groundwater Discharge 89

7–10�C. The flow rate of individual springs ranges from barely detectable to

1.5–1.8 L s�1 but the flow from moderate and high discharge springs remains

relatively constant (Fig. 3.17). Temperature of different springs ranges from

�4�C to 6.6�C, though there is little temperature fluctuation for individual springs.

Within each cluster, colder and lower-discharge springs surround a central area

with springs of warmer and higher discharge. This suggests a steady cooling of the

water by the permafrost that surrounds the talik through which the groundwater

flows (Pollard et al. 1999).

Thermal water also passes through permafrost without freezing. Figure 3.13

shows a thermal spring which is one of several hot springs that occupy the broad

valley floor of Pilgrim or Kruzgamepa River in Seward Peninsula. The water is at

about 80�C and flows through permafrost under artesian pressure.

3.3.3 Baseflow

Baseflow is that portion of streamflow that is supported by groundwater. For

continuous permafrost area, only suprapermafrost groundwater supplies water to

baseflow. With limited storage, this is not a reliable water source and baseflow can

cease during dry periods in summer. As winter approaches, the suprapermafrost

groundwater may turn into ice within the soil or is extruded and freezes above

ground as icing. With little water to feed the streams, discharge terminates during

winter. In discontinuous permafrost region, taliks provide passages for intra- and

subpermafrost groundwater to the streambed in support of icing formation and low

flow that continues through the winter. Williams and van Everdingen (1973)

estimated that groundwater contribution ranges between 2 and 5 L s�1 km�2,

with the larger discharge values associated with drainage from lakes. Baseflow

Aug Sep Oct

1997

0.0014

0.0013

0.0012

0.0011

Flo

w (

m3 s

-1)

Jul

Fig. 3.17 Flow rate of a spring, Colour Peak, Axel Heiberg Island (After Pollard et al. 1999)

90 3 Groundwater

decreases as permafrost increases in a basin. However, where river icing occurs,

both the amount of icing and the winter baseflow are part of the groundwater

discharge, which can constitute a large portion of total annual runoff. Clark and

Lauriol (1997), for example, estimated from baseflow analysis and icing

measurements that up to 50% of annual runoff of Firth River, Yukon, is derived

from groundwater. This is comparable to the 50–80% groundwater contribution to

flow in karst terrain of non-permafrost regions.

An increasing influx of mineralized water as a river flows downsteam is an

evidence of groundwater contribution to streamflow. Van Everdingen (1974)

analyzed the chemical composition of a set of water samples taken along the

lower 11 km of a small stream, Vermillion Creek, 38 km southeast of Norman

Wells, NWT. There is a tendency of downstream rise in ionic concentrations

(Fig. 3.18). Samples from the springs along the creek indicate elevated chemical

concentration compared with the river water. This suggests an increase in

Ca

Mg

SO4

Cl

Na + K

Downstream direction

10

1.0

0.1

Ion

conc

entr

atio

n (e

quiv

alen

ts p

er m

illio

n)

Fig. 3.18 Ionic concentrations in river water (shown as lines with symbols) along Vermillion

Creek, NWT, and concentrations in spring water (symbols only) between river reaches. Horizontal

axis is not to scale. Downstream rise of ionic concentration in river water suggests increased

contribution of groundwater to streamflow (Modified after van Everdingen 1974)

3.3 Groundwater Discharge 91

groundwater inflow that leads to enrichment of chemical concentration in the

stream. For some streams in discontinuous permafrost area, groundwater discharge

maintains a steady baseflow, and such flow is the principal if not the only source of

river water in the winter (Fig. 3.19).

Groundwater supplied to streamflow can come from several sources. Based on

water samples collected from the predominantly carbonate basins of Upper Firth

River and Joe Creek in Yukon, Clark and Lauriol (1997) found that the groundwater

has two sources. Methanogenic groundwater is derived from saturated soil with an

anaerobic environment found in the suprapermafrost zone. Karst groundwater is

recharged through unsaturated soils, moves through the talik of the carbonate rock

fissures and discharges to the valleys as streamflow (Fig. 3.20). At another site near

Aklavik, NWT, where Big Fish River intersects limestone and shale bedrocks,

Clark et al. (2001) used geochemical and stable isotope techniques to distinguish

three water types that contribute to baseflow. The thermal groundwater, at 16�C and

rich in sodium chloride, is from the subpermafrost zone. It is discharged at constant

rates year round and represents 85% of the winter baseflow. The ‘shallow ground-

water’ (probably intrapermafrost in nature) has high calcium sulfate content,

indicative of its circulation through taliks in the carbonate bedrock. Finally,

the suprapermafrost groundwater, low in salinity and high in calcium bicarbonate,

is the major component of runoff during spring melt but makes little contribution

in winter.

Fig. 3.19 Groundwater issued at the stream gaging station of Wolf Creek, Yukon, maintains

baseflow through the winter in a discontinuous permafrost basin. Note the uneven snow cover on

the boreal forest floor, with tree wells (see Chap. 4) around individual stands of spruce tree

92 3 Groundwater

3.3.4 Ponds and Lakes

There are myriad ponds in permafrost regions. In continuous permafrost areas,

transient talik linked with faults can tap into the subpermafrost groundwater, as

exemplified by a lake in Ny-Alesund (Fig. 3.21). For most ponds and lakes,

however, their groundwater supply comes entirely from the supraperamafrost

layer. Owing to the limited supply from the suprapermafrost source, a large

catchment area is required to maintain a high water level in the ponds. Marsh and

Woo (1977) used a water balance approach to calculate the amount of groundwater

flow needed to support a tundra pond in Vendom Fiord

dSp=dt ¼ Pp � Ep þ Q (3.2)

where dSp/dt is the change in storage in the pond, Pp and Ep are rainfall on and

evaporation from the pond surface, and Q is suprapermafrost groundwater

discharged into the pond. All variables are expressed in volumetric terms. The

surface area of this pond varied from 210 to 800 m2, depending on the water level

(Fig. 3.22). The period before 6 July 1975 was dry, the active layer storage and pond

icing

saturated zone

water

talik

permafrost

fissures

karstgroundwater

methanogenicgroundwater

Fig. 3.20 Methanogenic groundwater is derived from saturated soil with an anaerobic environ-

ment found in the suprapermafrost zone. Karst groundwater is recharged through unsaturated soils,

moves through the talik of the carbonate rock fissures and discharges to the valleys as streamflow

(After Clark and Lauriol 1997)

3.3 Groundwater Discharge 93

level were low. The following 2 weeks received 56% of the total summer rainfall,

adding 4.1 m3 of water directly to the pond, but evaporation loss amounted to

13.6 m3. Yet, the pond had a net storage increase of 24.0 m3. A groundwater inflow

of 33.5 m3 was needed to maintain this water balance. The discharge of

suprapermafrost groundwater depends not only on water added to the active layer

(through such sources as rainfall or snowmelt) but is also predicated upon its

dryness. The period 27 July to 2 August received only 22% of the seasonal rainfall

but the antecedent storage was high in the active layer. Pond level rose by 48 m3. In

general, (1) suprapermafrost groundwater discharge is the primary water source for

ponds, and (2) when dry, much of the rain replenishes the suprapermafrost ground-

water storage; but when the active layer is fully saturated, it cannot hold additional

rain and groundwater is readily discharged.

In the discontinuous permafrost region, intra- and subpermafrost groundwater

may discharge to lakes through taliks. Near Fairbanks, Alaska, the permafrost is

relatively warm (about �0.5�C), and thin (often <70 m). There are many small

lakes, one of which, the Isabella Creek bog lake (area 0.02 km2), was studied by

Kane and Slaughter (1973). This lake has a floating vegetation mat and a bottom of

mixed silt and organic matters. Temperature profile of the lake and its bed shows a

gradual increase in temperature with depth (Fig. 3.23). This indicates the existence

of a talik that links the lake to the subpermafrost water. Piezometers installed in the

lake bed show a downward increase in hydrostatic head which allows an upward

Fig. 3.21 A lake in

Ny-Alesund supported

partially by subpermafrost

groundwater discharged

through transient talik along

faults in continuous

permafrost

94 3 Groundwater

discharge of groundwater to the lake. While the lake studied is a point of ground-

water discharge (estimated to be 44 m3 day�1) to the subpermafrost aquifer, other

small lakes may serve the functions of recharge or discharge or both.

3.4 Icings

Icings (aufeis in German and naled in Russian) are formed by the freezing of water

that seeps from the ground, flows from a spring or emerges from beneath a river bed

or through fractures in river ice. Icings are not restricted to permafrost areas. Their

30

20

10

0 July 1975

m3

cumulative rainfall

cumulative evaporation

cumulative groundwater inflow

cumulative pond volume change

201510

silty organicdeposits

pond

gravels

contours (m)

100 m

N

Bedrock

70

80

80 90

110

100

68

68

70

66

66

Fig. 3.22 Water balance of a tundra pond, Vendom Fiord, showing that suprapermafrost ground-

water supply is essential in maintaining a high pond level in the summer (After Marsh and Woo

1976)

3.4 Icings 95

formation is due to processes that are common in areas where winter is cold enough

to freeze the exfiltrated groundwater. The exfiltration of water can be continuous or

intermittent.

Icings may be transparent or bubbly, and are often colored due to the presence of

dissolved minerals. When tinted by organic matter, ground icing may acquire a

yellowish-brown hue (Fig. 3.24). Mineral deposits frequently found on the icing

surface are likely caused by precipitation of salts during freezing (Akerman 1982).

3.4.1 Ground and Spring Icings

After surveying a plethora of ice terminology, Carey (1973) generalized icings into

ground, spring and river icings according to their mode of occurrence. Ground icing

is produced by seepage that simply saturates the ground surface in summer while

spring icing is due to the discharge of water that leads to well defined channels on

the ground. However, Carey also noted that the distinction between ground and

spring icing may not be entirely clear cut.

Water for ground icings may seep from soil pores or rock cracks, root channels

or animal burrows; whereas water for spring icings are issued from more defined

vents that may be single or multiple. The supply of water to ground icing is

normally limited and is often exhausted before the end of winter. Thus, they do

1 20

floating organic mat

mixed clasticand organicmaterials

0

1

2

3

4

5

6

7

Dep

th (

m)

Temperature (oC)

water

25July 1969

Fig. 3.23 Temperature

profile of Isabella Bog near

Fairbanks, Alaska, with cool

water overlying warmer water

at bottom, indicating the

presence of a talik that links

the bog to subpermafrost

groundwater (After Kane and

Slaughter 1973)

96 3 Groundwater

not grow to large sizes. Ground icings take on various surface expressions. They

can be flat to slightly arched. They can mantle steep slopes to form a cascade of ice

or icefalls.

Spring icings are usually larger than ground icing, being fed more reliably by

deeper groundwater. Geological conditions favorable for spring icing include fault

zones and exposed water-bearing rock strata. Yoshikawa et al. (2007) studied the

preferred locations of icing sites in Brooks Range, Alaska. They found a correspon-

dence between limestone areas and icing source springs. In addition, icings occur

preferentially between 200 and 900 m a.s.l. Above 900 m, the potential piezometric

head is too far below the ground surface and below 200 m, permafrost in the thick

Quaternary sediments inhibits upwelling of groundwater.

The presence of taliks beneath rivers and lakes can conduct water to the surface

to produce springs and icings. In the continuous permafrost environment of the

Canadian Shield, Vellette and Thomas (1979) found several icings on granular

surficial deposits. These icings are close to and at elevations lower than their source

lakes. Through drilling and ground temperature measurements, they deduced that

taliks exist below the lakes to act as subpermafrost conduits that feed the springs

and the icings.

Glaciers can enrich groundwater storage to sustain discharge. Akerman (1982))

noted that subpolar glaciers insulate the ground to maintain taliks from which water

is issued in the cold season to create icing at glacier terminals. Past glaciers can play

a role in facilitating icing occurrence. Yoshikawa et al. (2007) noted that the current

Fig. 3.24 Seepage of

groundwater from a northern

Ontario wetland (tinted by

organic matters to give a

brownish color) freezes to

form icing, which also

incorporates snow that drifts

or falls on the water

3.4 Icings 97

distribution of some springs and icings in Brooks Range is related to the existence

of Wisconsinan ice caps, the bottom of which potentially yielded large quantities of

groundwater to the subpermafrost aquifer to support a number of the present icing

sites.

3.4.1.1 Formation and Development

Groundwater that seeps or emerges slowly from springs will freeze as icing as it is

cooled by the low winter temperature to below its freezing point. Schohl and

Ettema (1986) presented a theoretical treatment of icing growth, which is here

simplified for a one-dimensional (vertical) case. Figure 3.25 is a conceptualized

profile of an icing topped by an overflowing water layer. It also provides definition

of the symbols used in the following equations.

Net heat flux from water to the air (Qwa in W m�2) is responsible for the

formation of ‘ice platelets’ which constitute a slush layer (mixture of water and

ice). Strictly speaking, this flux should be obtained by the energy balance at the

water surface. Such a physically based approach demands a substantial set of

micro-meteorological data. For practical considerations, it is expressed empirically

as a function of air temperature (T in �C) and temperature at the water surface (Twsin �C)

Qwa ¼ twaðTws � TÞ (3.3)

z

x

Frozenground

Ice

Water

T

Tws

Tw q'

Qwa

Qwi

Qice

hplatelet

hw

hi

Tf = freezing temperature

temperatureprofile heat flux discharge

Ti

Ice-water slush

Fig. 3.25 Conceptualized profile of an icing topped by an overflowing water layer (Modified after

Schohl and Ettema 1986). See text for explanation of symbols

98 3 Groundwater

with twa being a heat transfer coefficient that depends on meteorological conditions.

The volumetric growth of ice platelets per unit surface area of the slush layer that is

exposed to the air (dhplatelet/dt, in m s�1) is obtained by converting the latent heat of

fusion released as the platelets grow

dhplatelet=dt ¼ Qwa=ðrilf Þ (3.4)

with ri being the density of ice (kg m�3) and lf the latent heat of fusion (J kg�1).

The heat flux convected by the water to the ice–water interface is

Qwi ¼ twiðTw � Tf Þ (3.5)

where Tw is water temperature and Tf is the freezing temperature of water (e.g. 0�C),and twi is an empirical heat transfer coefficient. Temperature of ice beneath the

surface sheet of water is at or below Tf, and conduction from the water–ice interface

downward into the ice (Qice) is

Qice ¼ KiðdTi=dzÞ (3.6)

with Ki (in W m�1 C�1) being the thermal conductivity of ice and dTi/dz is the

temperature gradient in the ice. The change in ice thickness (dhi/dt) is

dhi=dt ¼ ðQice � QwiÞ=ðrilf Þ (3.7)

Here, icing will melt at its surface if Qice < Qwi (i.e. heat is absorbed by the

icing), and icing accretion occurs if Qice > Qwi (i.e. heat is released).

To evaluate the length to which an icing spreads, the magnitude of water flow

and the heat fluxes have to be considered. The vertical heat flux is to be linked to

horizontal heat transport. The amount of heat convected by water is governed by its

temperature and the rate of flow per unit width, q’ (m2 s�1), which diminishes with

distance down from the water source. A two-dimensional continuity equation for

icing growth is

@q0=@xþ @ð’hwÞ=@tþ ðri=rwÞ@ðhi þ hplateletÞ=@t ¼ 0 (3.8)

Taking Qwi ¼ 0 (for negligible heat convected from the water to the ice–water

interface) and substituting Eqs. 3.4 and 3.7 into Eq. 3.8 gives

@q0=@xþ @ð’hwÞ=@tþ ðQice þ QwaÞ=ðrwlf Þ ¼ 0 (3.9)

in which ’ (dimensionless) is porosity of the ice-water slush, hw is thickness of

water on the icing and rw (kg m�3) is water density.

If the rate of water freezing equals to the rate of water supply, there will be no

water accumulation in the icing slush, and ∂(’hw)/∂t ¼ 0. Consider the situation

that both the heat flux from water to the air (Qwa) and the heat conducted from the

3.4 Icings 99

ice–water interface to the ice below (Qice) do not vary, and assuming that the rate of

discharge from source is fixed at q’0, an integration of Eq. 3.9 for x ¼ (0, Xeq) yields

Xeq ¼ ðq00 rwlf Þ=ðQice þ QwaÞ (3.10)

where Xeq is the distance along the icing reached by the flow q’0 when water supplyequals (or is in equilibrium with) the rate at which water freezes. This equation also

implies that if the air gets colder (i.e. Qwa increases but Qice is held unchanged), the

length of icing decreases as the icing grows thicker.

The above theoretical treatment may be compared with experiments on icing

growth conducted in the field. Hu and Pollard (1997a) observed the following

sequence.

1. A thin water sheet of <2 mm flows downslope.

2. As water temperature has not equilibrated with the substrate, much of the heat

convected by water is conducted downward and the water travels only a limited

distance, being <2 m in the experiments.

3. When water reaches the freezing point, slush is formed with ice crystals attached

to the bottom.

4. Water flows at a reduced speed over and within the slush section, thus raising the

water level upslope.

5. Heat loss to the atmosphere leads to the formation of thin ice on the water

surface but the water flowing below it (interlayer flow) continues to convey heat

to the slush, causing it to melt and thus allowing the flow to travel a greater

distance.

6. When the flow is obstructed by freezing that closes the conduit for interlayer

flow, the water below this barrier drains off without replenishment, leaving

behind intra-layer cavities.

7. Upslope of the barrier, a second icing cycle begins, but since the underlying ice

temperature has been raised during the first cycle of icing formation, the time to

complete the cycle is longer. However, as the icing slope becomes steeper, the

flow speed increases and water can travel greater distances downslope. Further-

more, the flow may concentrate on a narrow section of the icing surface and

repeated melting can produce a small channel. It is noted that the experiments

were conducted on a plot with flows confined to a 1 m-wide strip. Under natural

situations, water can spread over a broader width.

3.4.2 River Icings

When the hydrostatic pressure of water is increased by the hydraulic gradient in a

stream channel and by the constriction of flow, water is expelled above the

confining cap which may be the frozen streambed or river ice. This water freezes

as river icing.

100 3 Groundwater

Kane (1981) described the sequence of river icing development over the cold

season, focusing on the interactions of icings with groundwater along the river

banks (Fig. 3.26).

1. In late summer when active layer thaws to a maximum, groundwater enters the

stream.

2. Early winter cooling of river water leads to the formation of an initial ice cover.

3. With increased ice thickens to constrict the channel cross section and with

snowfall that adds weight on the ice, water is under pressure. A simplified

equation to estimate the thickness of snow that has to be deposited on the ice

before the water level will reach the snow–ice interface is presented by Kane

(1981):

hs ¼ ðrw � riÞhi=rs (3.11)

where h is thickness, r is density and the subscripts w, i and s stand for water, ice

and snow, respectively.

icing water permafrost

hydrostatic head

hydrostatic head

water tableLate summer

Early winter

Late winter

seasonal frost

Fig. 3.26 Sequence of river icing development during freezing of the active layer, with interactions

between the icing and groundwater flow along river banks (Simplified after Kane 1981)

3.4 Icings 101

4. When the pressure head is above the ice–snow interface and where there are

fractures in the ice, river water will overflow the ice cover. Freezing of this

overflow water forms the initial icing layer.

5. As the ground freezes, the frozen soil exerts pressure on groundwater in the

remaining thawed zone of the active layer. The pressure is found to be greater

along the banks than further away from the river (Kane et al. 1973). This leads to

the flow of river water into bank storage. However, further away from the river,

suprapermafrost groundwater still moves towards the banks.

6. With an increased hydrostatic pressure adjacent to the river, upward forces

increase at the base of the icing and at the base of the seasonal frost. The icing

is arched and the frozen soil layer is lifted. The void left by the uplifted soil is filled

with water which subsequently freezes as horizontal ice lenses. Melting this ice

during breakup or in the summer causes subsidence or collapse of the river banks.

7. With hydrostatic head being higher than the ice surface in the channel (which

can be a composite of river ice, icing and snow incorporated into the ice), icing

growth continues. It is noted that suprapermafrost groundwater supply is limited

and icing growth relies on water from non-permafrost areas or through taliks that

tap into the subpermafrost sources.

8. During cold spells that last several days, fluid pressure in the vicinity of a river

tends to decrease and there can be a reverse in flow. However, with moderated

temperature at the end of winter, the fluid pressure will remain high.

Along a river channel reach, the supply of water and the loss of heat from the

water are the primary considerations for river icing formation. Hu and Pollard

(1997b) proposed a three-stage process in the growth of river icing. Initially, an

ice cover is formed in the stream channel. In the second stage, ice growth is a

combination of continued ice growth, incorporation of snow that falls on the ice and

freezing of the net seepage (inflow minus outflow) within the reach. The final stage

is the freezing of water that drains as recession flow after the cessation of seepage

input to the reach.

Extensive icing can be formed in discontinuous permafrost where there is

continuously winter discharge to the river bed (Fig. 3.27a). For rivers not sustained

by baseflow, such as those in continuous permafrost regions, streamflow ceases

when suprapermafrost groundwater is exhausted. Even after the cessation of flow,

there may be water trapped in the pools between the frozen riffles (Fig. 3.27b).

Continued freezing will force the pool water to move through the bed materials that

are still not frozen, or break through the ice cover to freeze as icing. In this case,

icing activity is localized and brief.

The growth and decay of river icings over the winter and spring seasons have

been described by Kane and Slaughter (1972), van Everdingen (1982) and

Yoshikawa et al. (2007). The river channels are filled with ice in early winter,

followed by the development of a smoother and thicker icing surface in mid-winter,

and by continued thickening and downstream expansion of the icing body in late

winter. Snow that falls on the ice surface is incorporated into the icing when the

seepage water freezes with the snow. Figure 3.28 compares the growth and decay

102 3 Groundwater

Fig. 3.27 (a) Residual icing in Erdaogou, Tibetan Plateau, undergoing surface ablation,

undercutting and tunneling by the river during summer. Icing reaches a maximum thickness of

3 m in the spring, the edge of its former extent is marked by the trim-line of tundra vegetation on

river banks. (b) A small river icing formed from water in a pool trapped in frozen streambed,

Eidsbotn Fiord, Devon Island

3.4 Icings 103

rates of selected icings from discontinuous permafrost (Burlap Creek west of

Burwash, Yukon, and Caribou-Poker Creek, Alaska) and from continuous perma-

frost (Kuparuk River in Alaska and Kolyma River in Siberia) regions. Rivers in

colder permafrost areas produce earlier and longer lasting icings than those in

discontinuous permafrost areas. The maximum volume of icing produced varies

considerably between rivers and from one winter to the other. Unless groundwater

supply is limited (in which case, icing development ceases before the winter is over),

icing thickness reaches a maximum just before melting begins. During the initial

melt period, meltwater from the icing and from snowmelt may refreeze in shaded

location. This ice growth on the existing icing surface is called spring-time icing

(Froehlich and Słupik 1982) though strictly speaking it is only re-worked ice

superimposed on the icing and is not produced directly by groundwater discharge.

Melting of most river icing is completed by mid-summer. In some cool summers,

however, extensive icings such as the one in Moma valley (a tributary of Indigirka

River in Siberia) may not melt entirely but are incorporated in the new icings formed

in the subsequent winter. It should be emphasized that rivers without subpermafrost

groundwater inputs or inflows from upstream do not form thick icings.

20

0

-20

-40

2.0

1.6

1.2

0.8

0.4

0

Air

tem

pera

ture

(°C

)Ic

ing

thic

knes

s (m

)

Oct Nov Dec Jan Feb Mar Apr

19801979

Burwash, Yukon

Burlap Creek near Burwash

Ice

accu

mul

atio

n (%

)

100

80

60

40

20

0

Kuparuk, Alaska

Hulahula, Alaska

Caribou-Poker, Alaska

Kolyma, Siberia

Sep Oct Nov Dec MarFebJan Apr May Jul AugJun Sep

Fig. 3.28 Growth and decay

of selected icings from

discontinuous permafrost

(Burlap Creek, Yukon, and

Caribou-Poker Creek,

Alaska) and from continuous

permafrost (Hulahula and

Kuparuk Rivers in Alaska and

Kolyma River in Siberia)

regions (After van

Everdingen 1982; Yoshikawa

et al. 2007)

104 3 Groundwater

3.4.3 Icing Dimension

Icings can measure from several square meters on a slope to over tens of square

kilometers in a valley. While ground icings are limited in size as they cease growing

when the suprapermafrost water supply is exhausted, spring icing can be much

larger. Carey (1973) noted that a spring discharge of only 1 ft3min-1 (0.47 L s�1)

can create an icing covering an acre (0.4 ha) of ground to a depth of over 1 ft (0.3 m)

within a month. River icings, supported by higher levels of discharge than ground

and spring icings, tend to be thicker and larger. An enormous icing was reported for

the valley of Moma River. It had a length of about 25 km, a width of 5.5–8 km and a

thickness of up to 4 m (quoted by Carey 1973). Icings can also infill narrow valleys

to a thickness of 7 m (tributary valleys of Firth River in Yukon reported by Clark

and Lauriol 1997) but they tend to be thinner when found on broad valleys with

braided streams. Although icings recur at the same general locations, their extent

and shape vary between years depending on the amount of water feeding the icing

and where the water freezes (Akerman 1982). Hu and Pollard (1997b) reported that

the volumes of two icings formed in 1993–1994 at two sites in Yukon, being

7.10 � 105 m3 at North Klondike River and 6.10 � 105 m3 at East Blackstone

River, were half of those reached in the previous winter.

Icing formation is usually limited by groundwater supply in continuous perma-

frost area during winter. There are exceptions, such as in West Spitsbergen where

winter discharge comes from local springs (thermal and non-thermal), rivers,

seepage around pingos and groundwater in taliks beneath some glaciers (Akerman

1982). At Expedition Fiord, Heldmann et al. (2005) reported highly saline springs

that formed an icing that reached 300 � 700 m2. Icing begins to develop in late

September when temperature drops below �7�C (freezing point of this saline

water). Networks of pipes and tunnels provide conduits for water flow to enlarge

the icing. Icing found at the furthest reach has the highest salinity. Maximum size is

attained in late April but then, air temperature rises to the point where the icing and

salt mixture melts and growth is replaced by ice deterioration.

It is tedious to measure the thickness of an icing but its area can be readily

surveyed or, for very large icings, obtained from satellite imagery. Based on long-

term studies of 310 icings in permafrost regions of Siberia, Sokolov (1973) used

icing area to estimate the volume of an icing by applying the empirical equation:

Vi ¼ aAib (3.12)

where Vi is the ice volume (in thousand m3) and Ai is the icing area (in thousand m2).

The coefficients of a ¼ 0.96 and b ¼ 1.09 were obtained by Sokolov andwere found

to be suitable also for the icing fields in the Brooks Ranges (Yoshikawa et al. 2007).

Icings can represent a large portion of groundwater discharge. Harden et al.

(1977) noted that the icings in some basins in northeastern Siberia store up to

25–30% of total river discharge and up to 60–80% of the subsurface drainage. For

Firth River which is typical of several major rivers on the North Slope of Alaska

3.4 Icings 105

and Yukon, Clark and Lauriol (1997) calculated that the volume of icing represents

38% of annual groundwater discharge, or 12% of total basin runoff. Similarly,

Yoshikawa et al. (2007) noted that icings constitute 27–30% of annual groundwater

discharge in the Kuparuk River.

3.4.4 Icing Problems

Carey (1973) indicated that ground icing is rare under natural conditions, but where

the environment is disturbed, it can become the dominant icing type. Ground water

seeped from cut slopes along highways can spread and freeze on the road surfaces

to form a sheet of icing that is hazardous to traffic. The construction of flow barriers

such as raised roadbeds, bridge footings or pipelines would increase the water

pressure upslope of the structures, leading to initiation of icing. Conduits designed

to allow slope drainage or to accommodate streamflow can be filled by icing.

Figure 3.29a shows a small culvert in Fort Simpson, NWT, blocked by icing and

Fig. 3.29b shows large culverts at a river crossing of the Dempster Highway to

accommodate icing. Despite their large size, icing formed in the culverts has to be

thawed by steam during the winter to prevent excessive build up. The pipes at the

top of the culverts allow a hose to be connected to pass heat into the culverts.

Fig. 3.29 Icing in discontinuous permafrost area: (a) Terry Prowse standing next to a small

culvert choked with icing, Fort Simpson, NWT. (b) Large culverts installed at a river crossing of

the Dempster Highway to accommodate icing

106 3 Groundwater

Buildings are not immune to icings. Muller (1945) gave an example of a heated

building which prevented the formation of an impermeable cap of seasonal frost,

thus causing groundwater to burst through the floor to form icing that filled the

house and spilled through its windows (Fig. 3.30).

3.5 Domed Ice Features

The freezing of groundwater concentrated under pressure at a particular site can

lead to the formation of domed features that range from small low mounds (frost

and icing mounds) to hillocks (pingos). Pollard and French (1984) emphasized that

the injection of free water and its freezing are important processes in their formation

though ice segregation can play a major role, especially in the case of pingo

development. Another feature is the palsa which is a peaty frost mound, often

considered to grow by the build-up of segregation ice (Sect. 8.3.1).

3.5.1 Frost Mounds and Icing Mounds

Van Everdingen (1978) suggested that frost mounds that are seasonal in nature

(i.e. form within one winter) be termed frost blisters. However, it is not easy to

unfrozen active layer

water flow under hydrostatic pressure

permafrost

Fig. 3.30 A heated building preventing formation of an impermeable cap of seasonal frost,

allowing downslope groundwater flow (under hydrostatic pressure) to rupture through the floor

and produce icing that fills the house and overflows from the windows (Redrawn after Muller

1945)

3.5 Domed Ice Features 107

distinguish between blisters and mounds on the basis of morphology and structure.

Since the formation of frost and icing mounds or blisters are subject to similar

thermal and hydrologic processes, the following section will not separate them on

the basis of whether they are preserved seasonally or perennially.

Van Everdingen (1982) conducted a detailed study of frost mound growth and

decay in the Bear Rock area near Fort Norman using time-lapse photography and

field measurements. For this area, a freezing index of at least 1,100�-days (runningsum of number of below-freezing degrees for each day during the winter season) is

needed to initiate a frost mound. Air temperature, however, may not be a direct

indicator of temperature at the ground surface as snow cover of different thickness

offers varying degree of insulation (Sect. 1.3.3.2). As frost penetrates the saturated

active layer, the spring water freezes to form icings. Icing eventually blocks the

spring outlet, restricting water flow but increasing the hydraulic pressure. Pollard

and French (1984) installed piezometers to measure pressure built up in several

seasonal frost mounds in North Fork Pass, and obtained values of 50–82 kPa in an

icing blister with a height of 2.19–2.30 m. Increased pressure leads to doming of the

frozen cap above the injected water and the rate of uplift was observed to be as fast

as 0.55 m per day. Dilation cracks are formed during the rupturing of the frozen soil.

Air entry through the fissures is possible but the cracks are then sealed off by

freezing, trapping the air inside the mound. Should compression of the entrapped air

build up large enough pressure within the mound, explosive rupture can occur.

Suslov (1961, p. 157) noted that groundwater flow through unfrozen ground from

slopes to the river bed of Onon River, Siberia, builds mounds 3–4 m high and

15–60 m in diameter. When delivered gradually, water may issue quietly through

the cracks of the mounds but if large hydrostatic pressure is built up beneath the ice,

the mounds would crack and shudder, then burst with a sound like gunfire. Ice

fragments from the mound are carried by the gushing water until the flow subsides

several hours later. Carey (1973) quoted such an event on this river when in 1928,

the explosion moved a block of ice that measured 19 � 5 � 1.7 m3.

These mounds can be a composite, such as icing overlying a frost mound. The

stratigraphic sequence of a typical frost mound has a frozen soil cover which may

contain segregated ice, overlying a layer of intrusive ice sometimes with pure ice

that suggests freezing of pooled water (Fig. 2.18), and an air cavity below left after

the water has drained. Frost mounds and icing mounds are formed by similar

processes (Fig. 3.31). The major difference is that for a frost mound, water does

not extrude above ground but freezes beneath a layer of sediments (Fig. 3.32)

whereas an icing mound relies on a previously formed icing cover to provide a

confining layer for the injected water (Michel and Paquette 2003).

3.5.2 Pingos

Pingos are relatively large perennial frost mounds with a core of massive ice and

covered with soil and vegetation (Permafrost Subcommittee 1988). They grow

108 3 Groundwater

Fig. 3.32 Wayne Pollard describing a frost mound along Dempster Highway, split open to reveal

the ice that domed up a covering layer of sediment

Early winter: icing

Winter: frost blister

Spring: thawing Spring: collapse

Winter: icing blister

permafrost

seasonally frozen

icing and ground ice

stored water

nonfrozen

discharge

subsurface flow

Fig. 3.31 Processes through which frost and icing mounds are formed (Modified after van

Everdingen 1982)

3.5 Domed Ice Features 109

slowly over a period of decades or more, to attain heights ranging from tens to over

100 m. The term is derived from pinguryuaq, an Inuit word for hill, and is

equivalent to the Russian name bulgunniakh.The classical form of a pingo is dome shaped with an oval base (Fig. 3.33), but

some are elongated (Fig. 3.34). Some pingos also resemble flat-topped hills

(Mackay 1979) or have a ridge-like outline. Pingos occur singly or align in a

group. In terms of their origins, there are two types of pingos: open system

(hydraulic or Greenland type) and closed system (hydrostatic or Mackenzie type)

pingos. Both types are created by the uplifting of a layer of frozen ground by the

pressure of water freezing in the substratum to form large ice masses. However,

they differ in terms of water source and processes of formation.

Open system pingos are found on steep hillslopes or on the lower gentle slopes of

valleys and in broad valley bottoms particularly those with alluvial materials.

Continuous or thin permafrost where taliks in sub- and intrapermafrost zones can

supply water under hydraulic pressure favor the development of open system

pingos, an example of which is provided in Fig. 3.35, as conceptualized by M€uller(1968).

Yoshikawa et al. (2003) gave examples of open system pingos along Caribou

Creek near Fairbanks. The artesian pressure in a collapsed pingo was at 50 kPa, as

measured in the winter of 2001–2002. Springs are found around this pingo. A spring

with water temperature of�0.01�C discharged at a rate of 1.24 L s�1. Injection ice,

possibly formed during the Holocene, likely froze near the interface of bedrock and

its overlying silt sediments. At present, the ice is at least 5 m thick and the top of this

pingo has collapsed to produce a crater-like depression. Disintegration at the pingo

Fig. 3.33 Ibyuk Pingo near Tuktoyaktuk, NWT, a thousand year old ice-cored hill with an oval

plan form, rising to 49 m above its surrounding drained lake bed, which is occupied by ice-wedge

polygon fields, wetlands and ponds

110 3 Groundwater

Fig. 3.34 Several open

system pingos aligned along

seepage zone in Reindalen,

Svalbard. Pingos exhibit

circular and elongated plan

forms

talik

active layer

pingo ice

ice

permafrost

talik

lifting force (artesian pressure)

water underhydraulicpressure

permafrost

ice

groundwaterflow

Fig. 3.35 Diagrammatic section of a slope with lateral drainage supplying groundwater to feed an

open system pingo, and cross section of a pingo (Modified after M€uller 1968)

3.5 Domed Ice Features 111

summit is common, and the summit depression may or may not contain a pond, as

shown by an example from Caribou basin (Fig. 3.36).

The largest known pingo in China is an open system pingo located at the

intersection of two set of faults with northeastern and west-northwestern

orientations. Situated along the Qinghai-Xizang Highway at the Kunlun Shan

Pass on the Tibetan Plateau, it reached 20 m height, 40–50 m length and 20–30 m

width in the 1960s (Qiu and Cheng 1995, Zhou et al. 2000, p. 349). The active layer

thickness is about 1.3 m and the massive ground ice below is of injection and

segregated types, inter-bedded with sediment layers. When a 57.35 m borehole at

the summit penetrated the permafrost layer, artesian discharge rose to over 20 m

and at a water temperature of 0.1�C. After the pingo was blasted two to three times

by highway construction crews, the position of the pingo has apparently shifted.

The pingo is now 18 m high, with length and width that reach 140 m and 45 m

(Fig. 3.37a). The summit is a crater-like depression with vents through which

groundwater gushes out (Fig. 3.37b). In winter, the discharge of subpermafrost

groundwater produces icing that can reach 5 m thickness.

Closed system pingos are restricted to the continuous permafrost region. Most of

them are found on drained lake beds with fine material overlying sand. The lakes

were deep enough to be underlain by taliks composed of saturated materials. When

a lake is drained, its bottom is exposed to subfreezing temperatures that enable the

formation of frozen soil. This creates an impervious cap for the saturated talik

which then becomes confined by permafrost on all sides. In addition to drained lake

beds, abandoned river channels can also have ancient taliks that were much larger

Fig. 3.36 A pond occupying the collapsed summit of an open system pingo, Caribou-Poker basin,

Alaska

112 3 Groundwater

Fig. 3.37 (a) Light-colored hill at center is an open system pingo located at the intersection of two

sets of faults, Kunlun Shan Pass, Tibetan Plateau, China. Pingo was 18 m high before an attempt

was made to blast it. (b) On the floor of the depression in the pingo, groundwater is seen (in 1986)gushing out from a vent which has shifted position over the years

3.5 Domed Ice Features 113

than those of the present (Pissart and French 1976). A change in the fluvial regime

and freezing of the sediments can provide the mechanisms for pingo formation

equivalent to the draining of lakes.

The next stage is permafrost aggradation that proceeds from the side, the bottom

and particularly the top of the transient isolated talik (Mackay 1998). The hydro-

static pressure thus created expels water from the pores in the unfrozen soil. This

water moves as groundwater to a residual pool which becomes the site of pingo

growth (Fig. 3.38). Water forms segregated ice, or is squeezed through the existing

ground ice and frozen soil to form injection ice. An increase in the volume of the

massive ice results in doming of the land to form a pingo. Initial growth of a pingo

may be 1.5 m per year, with growth later decreasing to several centimeters per year.

PT – PW < C

PT – PW > C

PT ≈ PW < C? ?

1

2

3

??

-10°C

0°C

-10°C

0°C

-10°C

0°CPw > PT

spring

residual pond drained

lake bottom

unfrozen frozen water ice

a

b

c

Fig 3.38 Formation of a closed system pingo (After Mackay 1979). (a) Permafrost is thinnest

beneath the center of a drained lake and is the preferred location for pingo growth. (b) Growth of

segregated ice. (c) Growth of intrusive ice and formation of peripheral failure. In these situations,

consider the relative magnitudes of PT (total resistance to heaving), PW (water pressure) and C

(soil constant). PT – PW < C: lens ice formation, PT � PW < C: lens ice and pore ice formation,

PT – PW > C: pore ice formation with water expulsion, PW > PT: sub-pingo water lens accumu-

lation, PW >> PT: pingo ruptures and water escapes

114 3 Groundwater

Doming leads to slumping of the sides of the pingo and the formation of cracks

which may contain dilation crack ice. Sometimes, water under pressure can erupt

through vents and shoot into the air (Fig. 3.39), or intrude in the areas peripheral to

the pingo to form ice mounds (so called baby pingos though this is not

recommended as a proper term).

Fig. 3.39 Pressurized

groundwater erupting from a

hole drilled on the side of a

closed system pingo, 22 m

below its summit,

Tuktoyaktuk, NWT (Photo

courtesy of J.R. Mackay)

gravel

?

?

??clay / ice

medium sand

medium sand

clay

creek

permafrost

lake silt

diamicton

0 200 250 300 350 400100 15050

0

20

40

-20

-40

North South

Distance (m)

Dep

th/h

eigh

t (m

) drillhole

?pingo ice

Fig. 3.40 Stratigraphic cross section of Ibyuk Pingo (a closed system pingo), Tuktoyaktuk (After

Mackay 1998)

3.5 Domed Ice Features 115

The Ibyuk Pingo near Tuktoyaktuk, NWT, is the largest closed system pingo in

Canada (Mackay 1979). It rises to 49 m above the surrounding flat (Fig. 3.33) which

is the bed of an old lake, possibly drained after 1,650 years BP. The bottom of the

ice core is at least 15 m above the level of the adjacent flat. Its overburden includes a

top layer of lake silt, above a diamicton that rests on glaciofluvial sand (Fig. 3.40).

Dilation cracks, enlarged by erosion to form gullies, radiate from the summit.

Thawing of the ice in the cracks has produced tunnels that extend into the ice.

This pingo may be 1,000 years or more old but its present growth is restricted to a

small area beneath the pingo center.

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