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the conductivity. When considering two areas with the same climatic conditions, several
other locally variable factors influence infiltration:
• Soil texture.
• Orography and slope.
• Snow: thicker snow cover leads to enhanced insulation and thus weaker freezing of the
soil, and vice versa.
• Vegetation: similar effects as snow, via weaker or stronger insulating properties of the
vegetation cover, and via its roots' effects on the soil structure.
These influencing factors can lead to patterned areas with surface-water-impeding zones
in some parts, and with permeable zones, where conditions for infiltration are much more
favourable, in others. In consequence, the impact of frozen soil on infiltration and on
hydraulic conductivity is strongly scale dependent (Niu and Yang 2006 ; Koren et al. 1999 ).
Thus, it is important to notice that the falling below the 0 C threshold does not lead to a
complete blocking of infiltration and percolation. These characteristics complicate the
implementation of cold regions' soil processes in land surface schemes for climate and
Earth system modelling.
The response of the soil to freezing leads to specific variations in the annual cycle of soil
hydrology. The snowmelt, which is usually constrained to a very short period of sometimes
less than two weeks, delivers a large water input to the land surface, which at this time of
the year is still frozen. Infiltration capacity is thus low, and much of the snowmelt water is
channelled into surface runoff. The thawing of the active layer begins immediately upon
the completion of snowmelt (Boike et al. 1998 ), dependent upon a number of factors
including soil material, duration of snow cover, soil moisture and ice content, and con-
vection of heat by groundwater (Woo 1986 ). The beginning of the thawing coincides with
high surface moisture values, and ice melting in near-surface layers occurs on the top of
still frozen, and thus less permeable, deeper layers. Consequently, subsurface water flows
are weak, and high soil moisture values develop within the still thin thawed upper layers.
Refreezing of infiltrated snowmelt water also contributes to this (Swenson et al. 2012 ).
Over the course of the warm summer season, the thawing and deepening of the active layer
increase the water-holding capacity of the soil, resulting in a decreasing surface water
contribution during precipitation events and a steadily increasing baseflow contribution
(Hinzman and Kane 1991 ). The latter is a lateral slow subsurface runoff that can develop as
the permafrost table forms a barrier to the deeper soil, where again water cannot easily
percolate. Due to the enlarged water storage and increased baseflow, upper soil layers can
also become drier in this part of the year. The autumn precipitation often coincides with the
start of the freezing season; thus, again high surface runoff rates are produced, yet much
lower than in spring. During winter, the decreased hydraulic conductivity in frozen soils
leads to the observed very low winter baseflow. Permafrost degradation due to a warming
trend will likely lead to a decreasing seasonal variability of water flows (Frampton et al.
2011 ), whereat results of Frampton et al. ( 2013 ) show that total runoff will first increase
and then decrease as the permafrost degradation progresses further to total thaw.
Apart from the above-mentioned effects of the soil processes on hydrological quantities,
perennially frozen ground shows some unique features that are examples for processes that
act on both long and short timescales and that are often highly nonlinear. Massive ice
wedges are one of these features, which occur in permafrost-dominated landscapes (French
1990 ). Water enters the soils through frost cracks and, through volume expansion during
freeze-up, further increases the cavities in the ground. Cryosuction leads to movement of
unfrozen, supercooled water towards the freezing front, and, over time, the ice body can
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