Lake and River Ice
An obvious and notable feature of lakes and
rivers in the North is that they are ice-covered for
portions of the year.
Its significant hydrological influence arises
through its effect on the flow and water level in a
stream, the water level in a lake, and through
seasonal storage represented by the ice itself, the
snowcover it carries, and the channel and lake
storage it induces.
Indeed it can be argued the hydrological extremes
of common interest, floods and low flows, are as
much a function of stream processes through the
action of ice, as they are of the catchment
processes of traditional concern.
While the peak discharge is primarily a function
of catchment processes such as snowmelt, the
peak water level (the cause of the flooding), is
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very much a function of the ice conditions on the
stream.
This is particularly so for the North where the
snowmelt peak is the peak discharge event of the
year and can occur while the stream is still icecovered or otherwise influenced by ice in the
channel.
For example, in the period 1983-87, ice jams
were involved in some 30% of the flood events
across Canada.
In New Brunswick ice-jam floods are responsible
for more flood damage than open-water floods.
The 1987 ice jams on the St. John River alone
cause $30 million damages.
On the other side of the country, in northwestern
Canada, the flood threat at almost all riverside
communities is primarily due to ice jams, not
summer floods.
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At the other extreme, low flow at a site on a coldregion stream can also depend heavily on ice
processes.
A striking example of this is the fact that the
discharge over Niagara Falls was halted on
March 29, 1848 by ice obstructing the outlet of
Lake Erie.
A more common circumstance is the minimum
discharge that occurs in October discharge in the
Clearwater River due to ice formation upstream,
rather than in late winter discharge from the
catchment.
The low flow frequency curves for several rivers
in northern Alberta show marked “abnormalities”
in the curves for smaller streams that are
explained by ice effects.
As well as influencing the extremes, ice effects
can have a major influence on the winter
hydrograph of cold-region streams in general. In
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streams the volume of water stored as ice, and as
channel storage due to the increase in water level
caused by the ice, can represent a significant
portion of winter flow which does not become
available until spring.
This may be particularly so for the lakedominated rivers of the Canadian Shield where
slight changes in the resistance to flow from the
outlet due to changes in the ice cover can trigger
enormous changes in lake storage.
Snowfall on lake ice can cause an increase in
flow from a lake.
The weight of water displaced from the lake must
equal the weight of the snowfall on the lake ice
(if the latter is simply floating, with little restraint
from the shore, as is often the case).
Hence a 0.3 m snowfall will displace ≈30 mm of
water from the lake, a flow that can be very
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significant in a stream in mid-winter in a
catchment with a large proportion of lakes.
Therefore, unlike on land, a water equivalent of
snow falling on lake ice is made immediately
available as flow (while a similar amount will be
made available in the spring when the snow
melts, it should not be counted twice when
evaluating the catchment yield).
Similarly, in fall, snow falling on land can remain
until spring.
For small northern lakes the duration of the ice
cover and its thickness (as it affects the volume
of water remaining under the ice) are very
important considerations for aquatic fauna, with
winter kill always being a possibility.
The hydrological effects of floating ice are
therefore of major concern in cold regions.
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As indicated, they are a major cause of floods in
Canada, but these floods are not just significant
because of the damages and loss of life they may
cause.
In other circumstances they can be beneficial; for
example, the multitude of lakes in the vast and
environmentally important Mackenzie and PeaceAthabasca Deltas in western Canada depend on
periodic flooding caused by ice jams to refill and
refresh them.
Freeze-up of a small, well-mixed lake in calm
weather occurs in a straightforward manner (as
discussed in the previous lecture).
When the lake has cooled sufficiently that the
surface water temperature falls to a little below
freezing during the diurnal minimum, a thin and
fragile ice sheet will form over the lake surface.
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It will remain for the rest of the winter if it
thickens fast enough to resist disturbances by
wind, or melt by a subsequent thaw.
When the lake has cooled sufficiently that the
surface water temperature falls to a little below
freezing during the diurnal minimum, a thin and
fragile ice sheet will form over the lake surface.
It will remain for the rest of the winter if it
thickens fast enough to resist disturbances by
wind, or melt by a subsequent thaw.
On a large lake the process is much less well
defined.
The first water to cool sufficiently for ice to form
is that in the shallow protected areas around the
shore.
Here sheet ice, or border ice, forms in much the
same manner on a small lake.
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However in the centre of the lake the surface is
continually disturbed by wind that prevents
formation of the initial thin sheet and instead the
heat demand from the cooling atmosphere is
satisfied by the formation and growth of small
(typically a millimeter or so in diameter) fine
discs of ice called frazil.
Being ice, the particles are less dense than the
water and, although they can be entrained and
carried to considerable depth by the windgenerated turbulence, they eventually accumulate
at the surface to form frazil slush.
Here they are exposed to the cold air temperature
and with the wind agitation, causes the frazil
slush to gradually agglomerate to form frazil
pans.
With air temperatures cold enough, the
concentration of pans, slush, and sheets broken
from the shore ice, can eventually become high
enough to dampen the surface disturbance by the
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wind and finally allow an initial solid ice cover to
form over the lake surface.
This ice cover will thicken by ice growth on the
underside of the ice cover.
The ice so formed is coarse-grained, with a
distinct columnar structure, and is almost
transparent.
The growth of this clear ice will be terminated by
a snowfall of sufficient weight to “sink” the ice
cover.
Typically an accumulation about half the ice
thickness is sufficient to do this.
Then water can flood out over the ice, saturating
the lower layers of snow.
This saturated snow begins to freeze, forming
snow ice.
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This ice is fine-grained and, due to the fine air
bubbles incorporated into it, is white and opaque.
If another snowfall occurs before the saturated
snow layer is completely frozen, an unfrozen
slush layer will be left behind as the process is
repeated.
The presence of the saturated snow renders the
ice cover isothermal after a short time, so there
will be no further heat transfer from the bottom
of the ice cover and ice growth there will cease
until all the saturated snow layers are frozen.
Because of the different grain structure and vastly
different transparency to solar radiation, the
relative proportions of clear and white ice in a
snowcover has considerable significance for the
life of the cover and events in the water below it.
While a large proportion of clear ice can be
expected on northern lakes where the initial ice
thickens rapidly and there is little snowfall, the
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ice cover on lakes in regions of heavy snowfall
can be predominantly white.
Even on the same lake, however, the proportions
of clear and white ice can vary significantly,
primarily due to the effects of snow drifting.
As the white is formed from snow, its presence
must be considered when conducting a snow
survey for a catchment with a significant
proportion of lakes and rivers.
While the water temperature at the under-ice
surface in a lake is at freezing, that just below be
significantly above freezing due to the winter
“inversion” caused by the fact that water reaches
its maximum density at 4oC.
Because of this “warm” water within the lake, the
flow at the outlet of the lake is above freezing.
The outlet can therefore remain open long after
the remainder of the lake is ice covered. This can
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have significant repercussions on the variation in
flow from the lake, and the winter hydrology of
the outlet stream.
The situation at freeze-up in a river is somewhat
similar to that of a large lake, with two major
differences: the turbulence in a river is generated
by its own flow, and is therefore ever-present
except in pools above rapids, bars, weirs, or
dams.
It is sufficient to prevent any thermal
stratification of the flow so that the water
temperature remains within a few hundredths of a
degree throughout the flow depth.
Again the first ice to form is sheet ice over the
quiet water of the shallows along the banks.
Out in the central region of the stream, the flow
and turbulence is usually sufficient to prevent the
formation of sheet ice on the surface.
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Instead frazil forms and is mixed throughout the
depth.
When first formed this frazil is actively growing
and has a propensity to attach itself to anything it
touches.
This includes the bed material where frazil will
accumulate if it is large enough to resist the
buoyant force of the accumulation; this is known
as anchor ice.
Such deposits can blanket rapids and gravel bed
rivers to a thickness of the order of a metre.
In a shallow stream, these deposits can grow to
project through the water surface to freeze solid
and directly obstruct the flow.
The deposits are then known as ice dams.
If the turbulence generated by the flow is not too
intense, frazil particles will accumulate at the
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surface to form slush, which in turn will
agglomerate into frazil pans.
Hence, as the cold weather continues, border ice
grows out from shore, while the concentration of
frazil pans increases on the remaining free
surface.
Eventually, at some point along the river, the
remaining space between the border ice edges
will be insufficient for the passage of the frazil
pans.
Then the pans will lodge between the border ice,
and other pans moving down from upstream will
begin to accumulate behind the lodgement.
This initial pack gradually extends upstream as
more frazil pans arrive, forming the initial ice
cover on the river.
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Downstream of the lodgement the production of
frazil and the growth of frazil pans must begin
again, which delays freeze-over downstream.
As the ice cover develops on the water surface
the resistance to flow in the remaining water way
under the ice increases.
The water level must therefore rise to provide a
water way sufficient to carry the discharge with
the increased resistance. This depth increase is
typically greater than 30% of the mean depth of
the open-water situation.
Furthermore, the ice cover or pack floats with
more than 90% of its thickness submerged.
This necessitates a further increase in water level
to provide the required water way.
The nature of the initial pack that forms depends
on weather conditions during its formation and
on the discharge and geometry of the river.
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If the river is steep and fast, and the air
temperatures mild during formation, the initial
pack will be thick and the rate of frazil-pan
production slow.
In these circumstances the increase in water level
as the ice cover passes a section will be large but
the pack progression slow.
For example, on the Nelson River, a stream
characterized by alternating pools and rapids, the
initial pack must develop thicknesses of up to 12
m to allow the pack to progress over the many
stretches of fast water.
Because of the enormous volume of ice required
for progression of a pack of such thickness, the
progression is slow, despite bitterly cold air
temperatures during freeze-up.
On the other hand, if the river is flat, the flow
slow and the cold intense during the formation of
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the pack, the pack will be thin and frazil
production high.
In these circumstances the pack that forms will be
simply one pan in thickness and the progression
upstream rapid.
The variation of water level as the pack passes
can have considerable hydrological significance.
The water required to fill the channel storage
associated with the increase in water level is
abstracted from the flow moving into the reach
from upstream.
This means the discharge moving downstream
below the pack is less.
This effect can be responsible for the minimum
discharge of the year.
After the pack is in place, it will begin to freeze
solid.
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The ice so formed is fine-grained and a little
translucent.
If the pack is thin enough, and the season long
and cold enough, the freezing front will progress
through the pack into the flow underneath.
The ice that is formed then is coarse-grained and
transparent, as described for lake ice.
Also as for lake ice, the freezing of the pack, or
growth of clear ice under it, will be terminated by
a snow accumulation heavy enough to submerge
the ice surface.
If open water exists upstream, frazil will be
generated and carried under the ice cover to be
deposited to form a slush-ice layer.
As the freezing front moves through this, finegrained ice is formed that is somewhat
translucent and often has a smoky hue due to
sediment particles incorporated in the ice.
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The above processes continue throughout the
winter. Because the ice cover is initiated at
different times in different places, and its
character depends on local stream characteristics,
it can be expected the final ice cover will be quite
heterogeneous, possibly varying from a relatively
thick pack at one location, to thin, totally clear
ice over a polynya at another location.
If there are reaches of fast water, areas of warm
ground water or lake outflow, or a large polynya,
there can also be areas of open water that survive
the winter.
With the onset of warm weather, the temperature
of the ice cover increases until melt begins at the
bottom and, if the ice is exposed to solar
radiation, within the ice cover itself.
This initial melt within the
grain boundaries, a process
and which, particularly in
columnar structure, can
cover occurs at the
known as candling,
clear ice, with its
cause a dramatic
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reduction in the ice-cover strength while having
little effect on its thickness.
Being fine-grained and opaque, white ice resists
the debilitating effects of solar radiation on its
strength with the result that areas of the ice cover
that are predominantly white ice can remain
strong when areas of clear ice, such as in a
freeze-up polynya, have decayed completely.
At about the same time, snowmelt begins and
causes an increase in water level.
The bottom-fast ice along the shore is not free to
follow the increase in water level and becomes
flooded, exposing the water to warming by solar
radiation and air temperature.
Any open-water areas expand as the heat the
water absorbs melts the ice cover downstream.
Areas of thin ice become weak and either melt or,
if at the upstream end of an open-water area,
break away due to the drag of the current, to
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accumulate at the downstream end of the open
water.
If the discharge remains reasonably steady these
processes of melt, decay, and fragmentation will
continue until the river reach is free of ice.
However, if the discharge increases substantially,
or an accumulation of fragmented ice develops to
the extent that the load it exerts on the
downstream ice cover causes the ice cover to fail,
an ice run can be initiated.
As the ice moves the depth falls, releasing water
from channel storage, the reverse of what
happened during freeze-up.
This releases a surge, or river wave that causes an
increase in discharge and water level downstream
and helps sustain the ice run.
Eventually this ice run will stall, either because it
encounters a sudden increase in water way width
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or depth, such as the confluence with a reservoir,
lake, or larger main stream, or at a sudden
reduction in channel slope, that absorbs the surge
maintaining the ice run.
When the ice run stalls an ice jam has formed and
the water level will increase substantially.
Eventually the ice jam will fail or move, possibly
releasing another surge that will trigger an ice run
again if any ice remains downstream.
This process is repeated, not necessarily
sequentially, until the whole river is finally free
of ice.
On a lake the process of ice decay and melt
begins as on a river.
It is typical that the ice closer to the shore melts
first, releasing the main body of ice to be moved
about the lake by the wind.
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As on a river, the rate of melt, and particularly
decay, of a lake ice cover depends on the
proportion of clear and white ice.
This effect can be sufficient to reverse the shoreto-centre progression of melt mentioned above.
On a large lake, wind can assist break-up by
blowing large ice floes about the lake once they
have been freed from shore by melt.
However, on more moderate-sized lakes the ice
more-or-less decays and melts in place, only
disturbed by wind when it is in a very frail state.
The above events are typical of a truly cold
region, so that the water body experiences only
one freeze-up and one break-up each year.
In more temperate regions there may be more
than one freeze-up and break-up cycle in a given
year, whereas other years there may be none at
all. In such situations events become a strong
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function of the quantity of ice that can be
generated in each cold spell. In North America
such a situation is typical of the Maritimes,
southern Ontario and New England, and of
British Columbia and the northern Pacific States
of the USA. Inland and north of these locations
the former scenario is more typical.
On lakes in the High Arctic the situation can be
such that there may be no break-up at all in a
particular year.
Chronologies of river and lake ice formation and
disappearance provide broad indicators of climate
change over extensive lowland areas.
Broad scale patterns of freeze-up are available for
Russia from 1893 to 1985.
In general, freeze-up in western Russia is 2-3
weeks later now than at the turn of the century,
whereas further east there is a slight trend toward
earlier freeze-up.
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Similar patterns are available for ice break-up
dates, with western Russia rivers breaking up 710 days earlier now than in the 19th century.
In North America, records from 1823 to 1994 at
six sites on the Great Lakes show that freeze-up
came later and break-up was earlier until the
1890s, but they have remained constant during
the 20th century.
Freeze-up and break-up dates of ice on lakes and
rivers provide consistent evidence of later freezeup and earlier break-up in the northern
hemisphere from 1846 to 1995.
Under conditions of overall annual warming, the
duration of river ice cover can be expected to be
reduced.
Many rivers within temperate regions would tend
to become ice-free, whereas in colder regions the
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present ice season could be shortened by up to
one month by 2050.
Warmer winters would cause more mid-winter
break-ups as rapid snowmelt becomes more
common.
Warmer spring temperatures could affect the
severity of the break-up, but the effect is the
result of a complex balancing between
downstream resistance (ice strength and
thickness) and upstream forces (flood wave).
Although thinner ice produced by a warmer
winter would tend to promote a thermal break-up,
this might be counteracted by the earlier timing
of the event, reducing break-up severity.
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