SOILSOILSOILSOIL2199-398XCopernicus GmbHGöttingen, Germany10.5194/soil-1-147-2015Permafrost soils and carbon cyclingPingC. L.cping@alaska.eduhttps://orcid.org/0000-0001-6134-0760JastrowJ. D.JorgensonM. T.MichaelsonG. J.ShurY. L.Agricultural and Forestry Experiment Station, Palmer Research Center, University of
Alaska Fairbanks, 1509 South Georgeson Road, Palmer, AK 99645, USABiosciences Division, Argonne National Laboratory, Argonne, IL 60439, USAAlaska Ecoscience, Fairbanks, AK 99775, USADepartment of Civil and Environmental Engineering, University of Alaska
Fairbanks, Fairbanks, AK 99775, USAC. L. Ping (cping@alaska.edu)5February2015111471714October201430October2014–24December2014This work is licensed under a Creative Commons Attribution 3.0 Unported License. To view a copy of this license, visit http://creativecommons.org/licenses/by/3.0/This article is available from https://soil.copernicus.org/articles/1/147/2015/soil-1-147-2015.htmlThe full text article is available as a PDF file from https://soil.copernicus.org/articles/1/147/2015/soil-1-147-2015.pdf
Knowledge of soils in the permafrost region has advanced immensely in recent
decades, despite the remoteness and inaccessibility of most of the region and
the sampling limitations posed by the severe environment. These efforts
significantly increased estimates of the amount of organic carbon stored in
permafrost-region soils and improved understanding of how pedogenic processes
unique to permafrost environments built enormous organic carbon stocks during
the Quaternary. This knowledge has also called attention to the importance of
permafrost-affected soils to the global carbon cycle and the potential
vulnerability of the region's soil organic carbon (SOC) stocks to changing
climatic conditions. In this review, we briefly introduce the permafrost
characteristics, ice structures, and cryopedogenic processes that shape the
development of permafrost-affected soils, and discuss their effects on soil
structures and on organic matter distributions within the soil profile. We
then examine the quantity of organic carbon stored in permafrost-region
soils, as well as the characteristics, intrinsic decomposability, and
potential vulnerability of this organic carbon to permafrost thaw under a
warming climate. Overall, frozen conditions and cryopedogenic processes, such
as cryoturbation, have slowed decomposition and enhanced the sequestration of
organic carbon in permafrost-affected soils over millennial timescales. Due
to the low temperatures, the organic matter in permafrost soils is often less
humified than in more temperate soils, making some portion of this stored
organic carbon relatively vulnerable to mineralization upon thawing of
permafrost.
Introduction
Permafrost soils are unique among the world's soils because their form and
function are greatly affected by cold temperatures and by the presence of
perennially frozen ground beneath their seasonally thawed layer (Karavaeva, 1969;
Tedrow, 1974; Everett et al., 1981; Makeev, 1981; Ping et al., 1998; Tarnocai
et al., 2009; Jones et al., 2010). Low temperatures and seasonal freezing and
thawing lead to strong heaving forces, differential thaw settlement,
development of segregated and massive ice, and a bewildering variety of
patterned ground features. Because the underlying permafrost impedes
subsurface drainage, soils are often wet as well as cold, which greatly
affects oxidation/reduction, decomposition, and other biogeochemical
processes. Low decomposition rates, cryoturbation resulting from freeze–thaw
processes, and depositional environments dating back to the Pleistocene have
led to the accumulation of large stores of organic carbon in the active layer
and underlying permafrost. Due to these unique characteristics,
permafrost-affected soils (commonly termed permafrost soils) are
differentiated at the highest levels in most national and international soil
classification systems.
The enormous organic carbon accumulations in these soils make them important
to the global climate system because of their potential to thaw, decompose
organic matter accumulated over long periods, and release greenhouse gases
into the atmosphere. Soils in the northern circumpolar permafrost region are
estimated to store ∼ 1000 Pg C (carbon) in the surface 0–3 m and
permafrost-affected soils account for 70 % of this carbon (Tarnocai et al.,
2009; Hugelius et al., 2014). Because of the large carbon stocks, there are
increasing concerns about the potential release of significant amounts of
this carbon as greenhouse gases, especially for methane with its greater
warming potential compared to carbon dioxide (McGuire et al., 2009; Schuur et
al., 2011). The release of carbon may be exacerbated by a strong positive
feedback loop where rising soil temperatures and accelerated permafrost
degradation lead to increased decomposition of stored organic matter (Oechel et
al., 1993; Koven et al., 2011; Harden et al., 2012). However, once permafrost
thaws there are radical changes in hydrology, soil conditions, and vegetation
communities that create substantial uncertainty as to how quickly and in what
direction the thawing of permafrost soils will affect soil carbon balance,
greenhouse gas emissions, and climate change (McGuire et al., 2009; Olefeldt
et al., 2013; Jorgenson et al., 2013; Walter Anthony et al., 2014).
In this review, we highlight and discuss important factors affecting the
patterns, processes, and carbon stocks of permafrost soils, and summarize
recent research developments. In Sect. 2 on permafrost characteristics and
transformations, we discuss the formation and distribution of permafrost,
effects of cryogenesis on cryostructure patterns, the importance of
periglacial freeze–thaw processes on patterned ground, and the consequences of
thermokarst landforms resulting from degrading permafrost. In Sect. 3 on
cryopedogenesis, we review how biochemical processes are affected by low
temperatures, how soil structures differ in the active layer compared to the
upper permafrost, and the classification of permafrost soils. In Sects. 4–6
on organic carbon in permafrost soils, we examine permafrost-related factors
affecting soil organic carbon (SOC) accumulation, the quantity of organic
carbon held in permafrost soils, and its quality and decomposability.
Finally, in Sect. 7, we review recent research on the emerging role of
permafrost soils in climatic change.
Permafrost characteristics and transformations Permafrost formation and distribution
Permafrost is defined as ground (soil or rock, and included ice) that remains
at or below 0 ∘C for at least two consecutive years (Sumgin, 1927;
Washburn, 1973). Permafrost underlies approximately 22.79×106 km2, nearly 24 % of the landmass of the Northern Hemisphere
or ∼ 15 % of the global landmass (Zhang et al., 2008). Permafrost
terrain occurs mostly in the lowlands, hills and mountains of the circumpolar region, the
boreal regions, and high alpine and plateau regions in the Northern Hemisphere
(Zhang et al., 2008), in Antarctica, and in limited
areas of high alpine
regions in the Southern Hemisphere (Bockheim, 1995; Brown et al., 2001). The
non-ice-covered land areas of the northern circumpolar region encompass
17.8 × 106 km2, and permafrost-affected soils account for
57 % of this land area (Hugelius et al., 2014).
There are two general types of permafrost. The most common type is the old
permafrost, a remnant of paleoclimate, and the other type formed more
recently, under contemporary conditions. The old permafrost is widespread
throughout the lowlands and hills of the northern circumpolar region (Gubin,
1993; Péwé, 1975; Kanevskiy et al., 2011; Winterfeld et al., 2011).
Much of this relic permafrost is polygenetic in that syngenetic permafrost
developed on epigenetic permafrost. Examples include the Yedoma formations in
northeastern Russia, Arctic and boreal Alaska (Reyes et al., 2010;
Schirrmeister et al., 2011a, b; Kanevskiy et al., 2011, 2014), and the Yukon
regions in Canada (Froese et al., 2008). Yedoma is also called “Ice
Complex” because of the huge ice wedges formed in thick deposits of
syngenetic origin (i.e., concurrent upward growth of deposits and the
permafrost surface). Most epigenetic permafrost (which develops on existing deposits) formed in uplands, such as
the glaciated uplands of North America. The second, more modern kind of
permafrost was formed during the Holocene as a result of deglaciation
followed by vegetation succession under cold climatic conditions, and is
mostly located in the sub-Arctic and boreal regions (Zoltai and Tarnocai,
1975). This more recently formed type has been termed “ecosystem-driven
permafrost” (Shur and Jorgenson, 2007). In general, most permafrost in the
Arctic is of late-Pleistocene age (Brown, 1965; Gubin, 1993; Gubin and
Lupachev, 2008; Kanevskiy et al., 2011; Schirrmeister et al., 2011a). In the
Dry Valley region of Antarctica, the permafrost is even older, as it formed
during the pre–late Quaternary (Bockheim, 1995; Campbell et al., 1998).
However, the upper permafrost of the old formations is subject to change
under contemporary climate fluctuations.
Cryogenesis and cryostructure
Permafrost affects soil formation through physical and biological processes.
Permafrost exerts a controlling role on soil physical and morphological
properties through cryogenesis, i.e., frost cracking and the formation of ice
crystals and lens cryostructures that result in frost heaving and
cryoturbation (Konishchev, 1982; Rieger, 1983; Mueller et al., 1999; Ping et
al., 2008a). Cryostructures are defined as patterns formed by ice inclusions
in the frozen soil (Kanevskiy et al., 2011). These inclusions form as ice
crystals and lenses and as layers of segregated ice in the mineral soil
matrix due to freezing conditions within the soil (Shur and Jorgenson, 1998;
Shur et al., 2005; French and Shur, 2010). There are seven major types of ice
cryostructures in the active layer and the upper permafrost: structureless
(pore), lenticular (micro-lenticular: < 0.5 mm thick; lenticular:
> 0.5 mm thick), layered, reticulate, irregular reticulate (braided),
crustal, and suspended (ataxitic) (Mackay, 1974; Murton and French, 1994;
French and Shur, 2010). Some examples of these cryostructures are shown in
Fig. 1. Upon desaturation after thawing, these cryostructures generally leave
soil with platy, blocky, wedge-shaped, granular, or massive structures (Smith
et al., 1991; Ping, 2013).
Examples of common cryostructures: (a) lenticular
structures visible below 30–40 cm in a freshly exposed profile of an earth hummock,
Mould Bay, Arctic Canada (reprinted by permission, ASA, CSSA, SSSA);
(b) lenticular structures transition to a reticulate structure at
130–150 cm in a core extracted from the center of a low-centered ice-wedge
polygon; (c) lenticular structure; and (d) suspended
(ataxitic) structure.
According to Gubin and Lupachev (2008) and French and Shur (2010), a
permafrost soil profile commonly consists of three layers referred to as
active, transient, and intermediate layers. But, considering the whole
cryostratigraphy, the profile also should include a fourth layer – the
“true permafrost”, which often contains buried or paleo-genetic horizons
(Höfle and Ping, 1996; Schirrmeister et al., 2002a, 2008, 2011a;
Kanevskiy et al., 2014). The active layer is defined as the zone of the soil
profile above the permafrost that is subject to annual freeze–thaw cycles
(Burn, 1998). The active layer is the modern soil (mostly formed since the
Holocene and a product of the contemporary climate) that undergoes seasonal
thawing. The active layer typically consists of a surface organic horizon (O)
– with or without a humus-rich surface mineral horizon (A) – and a cambic
horizon (Bw, Bg). But, the entire active layer can consist of just mineral or
organic horizons (Ping et al., 1998, 2013). The transient layer (the uppermost
part of the permafrost) results from fluctuations of the permafrost table
that have occurred on a decadal scale and has distinct layered, lenticular
and reticulate cryostructures (Shur et al., 2005). The intermediate layer is
caused by aggradation of permafrost due to soil climate changes resulting
from the buildup of surface organic horizons on a decadal to century scale.
The intermediate layer is characterized by ice-rich cryostructures that are
referred to as “suspended” (or “ataxitic”) because blocks of soil are
suspended in an ice matrix (Fig. 1). Together the transient and intermediate
layers comprise the uppermost section of permafrost (Shur, 1988), and below
it is the “true” permafrost, which has remained frozen and not been
subjected to freeze–thaw cycles on a century to millennial scale (French and
Shur, 2010).
Periglacial processes and patterned ground
Toward the end of the growing season, solar radiation decreases, and the
active layer starts to refreeze. In the continuous permafrost zone with cold
permafrost, the active layer begins to refreeze from both the top down and
the bottom up, while in the discontinuous zone with warmer permafrost, the
active layer freezes only from the top down. The part of the active layer
that has not yet frozen is isothermal and is called the “zero-curtain
envelope” (Rieger, 1983; Davis, 2001). The zero curtain in the active layer
lasts from the initiation of refreezing each year until the layer is
completely frozen. In the Arctic, for example at Barrow, Alaska, the zero
curtain only lasts for several weeks (Outcalt et al., 1990). In the
discontinuous permafrost zone, for example in Fairbanks, Alaska, the zero
curtain occurs for several months. But, in warm winters or those with a thick
snow cover, part of the active layer can remain unfrozen, maintaining the
zero curtain effect through the entire winter (Jorgenson et al., 2000).
During the freeze-up process, lenticular, reticulate, and/or layered
cryostructures form in both the upper and lower parts of the active layer,
drawing soil water away from the middle part of the active layer and leaving
it relatively dry. This leaves a desiccated layer with either coarse platy or
massive soil structure (Shur and Ping, 1994). The exception is in poorly
drained soils, where there is adequate water to sustain layered cryostructure
development throughout the entire active layer.
Frost heave is caused by volumetric changes associated with the phase change
of water to ice. Heaving due to ice segregation and ice-lens formation often
results in deformation of the ground surface. With horizontal expansion
limited, there is often enough stress to produce crooked or tilted lenticular
and reticulate structures (Murton and French, 1994). Differential frost heave
eventually deforms originally flat horizons into warped or wavy horizons.
When freezing occurs in saturated coarse-grained soil, the cryostatic
pressure pushes water out of the soil, producing intrusive-ice features such
as icing blisters and frost blisters (Tsytovich, 1975; van Everdingen, 1978).
Soil deformation occurs during active-layer freezing. When fine-grained soil
of the active layer freezes from above and below on two freezing fronts as a
closed system, water migrates to both freezing fronts. This movement causes
the upper and lower parts of the active layer to become saturated with ice
and leads to formation of vertical cracks in the desiccated part of the
active layer. In contrast, during freeze up in an open system, the flow of
water through the active layer can create artesian pressure that forces water
out of late freezing areas (e.g., near the base of hillslopes). Ice buildup
in these areas produces cracked and broken surface soil horizons (Zhestkova,
1982). When ice-rich layers below the active layer thaw, the release and
movement of water leads to a type of micro-scale diapirism, where water and
saturated materials are forced through brittle layers or cracks while heavier
mineral- and organic-rich soils fill the ensuing voids (Swanson et al.,
1999).
Patterned ground derived from freeze–thaw processes greatly affects soil
formation and the complexity of soils on the micro-scale. Common patterned
ground types include ice-wedge polygons, sorted and nonsorted circles, stone
nets, stripes, and gelifluction lobes (Washburn, 1973). Ice-wedge polygons
form when contraction and expansion of the ice–soil permafrost matrix create
a large-scale net of cracks (5–50 m diameter polygon units) in winter. The
thermal cracks fill with water during the warm season and refreeze
immediately (Lachenbruch, 1966). This process repeats itself annually in new
cracks in roughly the same areas to build wedges that control surface water
distribution and deform developing soils (Kanevskiy et al., 2013; Ping et
al., 1998).
Circle pattern formation generally is initiated by frost cracking of the
active-layer ground followed by repeated freeze–thaw cycles, after which the
thermo-hydrologic conditions around the cracks are modified by vegetation and
ice aggradation (Shur et al., 2008; Walker et al., 2008). As a result of
these modified thermo-hydrologic conditions, the surfaces across circles
heave at differential rates, with greater heaving in the centers relative to
the margins. For soils without rock fragments, materials within and outside
the circles have the same general texture, and these are classified as
“nonsorted” circles or more commonly “frost boils”. For soils with rock
fragments that tend to heave up, these fragments are eventually pushed to the
outer edge of the circle by repeated cycles of frost action, forming
segregated domains of soil and rock that are referred to as “sorted”
circles (Washburn, 1973). In the initial stage, circle formation leaves the
soil surface devoid of vegetation. Progressively, vegetation communities
establish and an organic horizon starts to build. However, with continued
annual differential heaving of the circle surface, the surface organic mat
often becomes ruptured, forming discontinuous surface organic horizons. On
slopes, gravitational forces can cause both nonsorted and sorted circles to
deform downslope as they go through their annual freeze–thaw cycles, and
these deformations develop into patterns known as stripes, or stone stripes
when rock fragments are present (Marr, 1969; Geisler and Ping, 2013).
Oriented rocks are common in sorted circles. Orientation of rocks is caused
by thermally induced frost heave, similar to a common phenomenon observed in
soils wherever seasonal frost occurs. The mechanism driving rock orientation
is that the longest rock axis tends to orient itself perpendicularly to the
frost table, while the short axis lies parallel in order to minimize
resistance (Washburn, 1973; Davis, 2001). Silt-capped rocks commonly occur in
fragmental soils on exposed landscapes in the Arctic and alpine regions (Munn
and Spackman, 1990). The freeze–thaw cycles contribute to the grinding of
rock fragments into silt-sized particles, which are then transported by
percolating water and accumulate on the upward facing surfaces of stones.
Patterned ground formation can cause large- and small-scale variation in soil
types across the Arctic landscape. For example, across ice-wedge polygons,
soils in the polygon troughs often have moderately thick peat underlain by
massive ice; soils on the rims are better drained with well-decomposed
organic matter forming A horizons (Fig. 2); while soils in the polygonal
centers have thick peats over mineral sediment with different degrees of
cryoturbation (Ping et al., 1998, 2011). For circle patterns, the small-scale
net of cracks (0.5–5 m) in the active layer affects moisture availability,
plant community distributions and differential active-layer dynamics (Walker
et al., 2008), and creates different soil profile characteristics across the
circles (Michaelson et al., 2008).
Soils formed in ice-wedge polygons: (a) active layer
deformed by newly formed ice wedge (7 cm wide); (b) cryoturbated
polygon rim of a flat-centered ice-wedge polygon (ice wedge is 85 cm wide);
and (c) an aerial view of the ice-wedge polygon-dominated landscape
of the Arctic Coastal Plain, Alaska.
Thermokarst
Freezing of fine-grained soil attracts water to the freezing front from
unfrozen soil below. As this water freezes within existing pore spaces, it
can increase pore volumes by pushing soil solids apart, often forming lenses
and layers of ice. Ice buildup can greatly increase the volume of pores
present in soil compared to before freezing. When permafrost thaws and ice is
lost from these pores, this causes the soil surface to settle or liquefy, and
the amount of settlement is directly related to the amount and type of ice
(Shur and Osterkamp, 2007). The irregular topography resulting from the
melting of excess ground ice and subsequent ground collapse is called
thermokarst. Czudek and Demek (1970) identified two types of thermokarst. The
first type is permafrost “back-wearing”, which commonly occurs in areas of
dissected relief and develops gullies, thermocirques, parallel retreat of
steep walls with ice veins, and eventually lower-level lowlands. The second
type is referred to as “down-wearing” because it is caused by permafrost
thawing from above. This type is usually found in areas with flat undissected
relief and typically produces flat-bottomed depressions with steep slopes –
such as “alases”, a Russian term used to describe this landform and
thermokarst processes leading to its formation. The patterns and amount of
settlement (or loss) of surficial material are related to complex
interactions among slope position, soil texture, hydrology, and vegetation
over time (Shur and Jorgenson, 2007). The highly variable terrain and an
array of permafrost factors can lead to a wide variety of thermokarst
landforms that include degrading ice-wedge troughs, thermokarst pits,
thermokarst lakes, thermokarst bogs, thaw slumps, active-layer detachment
slides, and thermal erosion gullies (Grosse et al., 2013; Kokelj and
Jorgenson, 2013; Jorgenson et al., 2013; Jensen et al., 2014).
Thermokarst is widespread throughout the Arctic and boreal regions and has
large implications for soil hydrology and carbon balance (Burn and Smith,
1990; Schuur et al., 2008; Veremeeva and Gubin, 2009; Grosse et al., 2011).
Thermokarst lakes often develop taliks (unfrozen thaw bulbs) underneath the
deep water, and resulting heat-transfer dynamics can lead to degradation of
adjacent permafrost, lake-bottom subsidence, and lateral lake expansion.
Through these processes, organic matter that has long been sequestered in
permafrost and reworked by shoreline erosion can decompose and release
greenhouse gases (Walter et al., 2007; Grosse et al., 2013). But, thermokarst
lakes and their expansion can also provide good conditions for enhanced
primary productivity and carbon gain (Walter-Anthony et al., 2014).
Thermokarst bogs in the boreal region also develop thick taliks, where
organic matter in previously frozen soils becomes susceptible to
decomposition. However, rapid sedge and Sphagnum colonization in the
depressions can add new peat, and the net effect on carbon balance is
uncertain (Sannel and Kuhry, 2011; Jorgenson et al., 2013). Degradation of
ice wedges in the Arctic has increased, creating trough-like depressions that
collect water and provide an anaerobic environment for new peat accumulation
while the adjacent polygon becomes better drained (Jorgenson et al., 2006).
Thermal erosion gullies often develop along thawing ice wedges, channelize
surface water flow, and can lead to drying of adjacent soils (Godin et al.,
2014). Thaw slumps and active-layer detachment slides are common on slopes
with an ice-rich intermediate layer or where buried glacial ice is abundant
(Kokelj and Jorgenson, 2013). This disturbance quickly removes surface
organic layers, which can be exported to rivers and lakes or re-buried in
thick debris lobes. The fate of this material is uncertain, but some portion
may be re-stabilized in new environments. Furthermore, in some cases, it
appears that stabilization and recovery of newly exposed surfaces can occur
within several decades as changes in vegetation composition and ecosystem
biogeochemistry promote greater productivity and relatively rapid
accumulation of new organic matter (Pizano et al., 2014).
Cryopedogenesis
Cryopedogenesis refers to soil formation processes as affected by freezing
temperatures and freeze–thaw processes or cryogenic processes. Some of the
most important direct effects result from the physical barrier the permafrost
zone imposes on (1) water movement and biogeochemical processes at low
temperatures and (2) the expansion and contraction of the active layer during
the formation and melting of seasonal ice. Permafrost also exerts seasonal
cooling effects on the bottom of the active layer, which affects ice
formation and the freeze up of the active layer.
Biochemical processes at low temperatures
The accumulation of soil organic matter (SOM) in permafrost regions is
enhanced by two factors that slow the rate of decomposition: (1) low
temperatures and (2) anaerobic conditions caused by high moisture content in
the active layer (Kaiser et al., 2007; Rodionov et al., 2007). Cool soil
temperatures also slow chemical alteration of soil minerals and biological
activity. The zero-curtain zone of the active layer provides favorable
conditions for heterotrophic soil respiration during the late fall to early
winter. Thus, appreciable gas fluxes have been measured on the snow surface
during winter (Zimov et al., 1993, 1996; Fahnestock et al., 1999). It is
likely that frost cracks provide passages for gases to reach the atmosphere
as there is often a strong pulse of gases during early spring thaw. Under
saturated conditions, the processes and rates of SOM degradation involve
complex biogeochemical interactions that can lead to release of either
methane or carbon dioxide (Megonigal et al., 2004). While such processes have
been shown to occur at low temperatures (Gilichinsky and Rivkina, 2011),
efforts to understand the dominant microbial processes and abiotic factors
controlling these processes and their interactions in permafrost soils are
currently the subject of increasing research efforts (e.g., Rivkina et al.,
2007; Wagner et al., 2007; Lipson et al., 2013; Frank-Fahle et al., 2014).
In addition, mineral elements can be reduced even at sub-zero temperatures.
Most commonly manganese (Mn) and iron (Fe) are first reduced followed by
sulfur and then carbon dioxide, with reduction of the latter producing
methane (Patrick and Jugsujinda, 1992). Elements like Mn and Fe in minerals
are reduced by microbial processes and can be released into the aquatic
system (Lipson et al., 2010). When the concentration of SOM is high, reduced
Fe (Fe2+) in soil solution is rapidly oxidized upon contact with air,
forms a poorly ordered Fe3+ oxide film on water surfaces, and
eventually precipitates as orange-colored deposits called ferrihydrite
(Schwertmann and Taylor, 1989). The hydromorphism also results in the gleyed
color of the mineral matrix in the lower active layer due to Fe reduction
(Ping et al., 1993, 2008a).
Soil structure in the active layer
Cryogenesis commonly results in some unique soil structures that are limited
to the Arctic, sub-Arctic, Antarctic, sub-Antarctic, and alpine regions. Ice
formation, as the active layer refreezes in the fall and winter, varies
depending on the conditions that occur in different parts of the active
layer. This variation results in the development of an assortment of soil
cryogenic structures. On the exposed ground surface of silty mineral soils
with no (or only thin) surface organic horizons, crumb and granular
structures develop due to needle ice formation (Ping et al., 2008a). Ice
crystals form on frozen soil particles or structural units near the surface
when water is drawn from below to form or build the needle-like ice crystals
that lift and move soil particles or structural units (Brink et al., 1967).
Under snow cover, sublimation ice can build near the surface from water vapor
in the soil atmosphere, and this can also produce needle ice. In this
process, small amounts of surface-accumulated organic matter can be mixed
with the top few centimeters of mineral soil – promoting aggregation,
disrupting root establishment, and favoring biotic crust formation (Pawluk,
1988; Michaelson et al., 2008, 2012). In the High Arctic, where vegetative
and snow cover are scant, a persistent net of soil surface cracks form due to
freeze-desiccation contraction, creating small frost polygons that in turn
produce microenvironments that support different vegetation communities
(Walker et al., 2008). In areas where snow banks occur on slopes, the
micro-cracking and vegetation pattern can interact with erosion and
deposition patterns to form small < 1 m diameter turf hummocks (Tarnocai
et al., 2006; Broll and Tarnocai, 2002; Lewkowicz, 2011). But, these small
polygons and hummocks caused by frost-desiccation cracks result in little
cryoturbation, largely due to rapid freeze up and insufficient moisture to
develop segregated ice during the freeze up. These surface cryogenic
processes are most important on the surface of exposed mineral soils, such as
those of active frost boils in the Low Arctic and more generally across the
landscapes of the High Arctic.
Soil structure in the upper permafrost
As mentioned above, the transient layer is the upper zone of permafrost that
has formed more recently. This layer can be thawed during extremes of the climate
cycle but remains frozen most years. For this reason, the transient layer may be more
ice rich and be similar in ice and soil structures to the lower active layer
when it is frozen. Angular blocky, lenticular, or platy soil structures are
common. The upper permafrost below the intermediate layer, however, is
usually devoid of visible or larger ice bodies or lenses. Rather, it is ice
cemented or has porous visible or invisible ice crystals (Murton and French,
1994). The cryostructure of upper permafrost deposits is often structureless.
The most common exceptions occur in permafrost landscapes that have developed
for long periods of time or under many thaw-lake cycles or periodic
depositional cycles, where the permafrost has moved steadily upward
encompassing many previous transient and active layers (Hinkel et al., 2003).
Permafrost soils can contain very large proportions of ice in the
upper-permafrost horizons (the transient and intermediate layers of Shur,
1988). Investigations over large exposed areas of Alaska's Arctic coastline
found the overall average ground ice content of permafrost soils to be 77 %
by volume in the upper few meters of soil, ranging from 89 % in Yedoma
formations to 43 % in eolian deposits (Shur and Zhestkova, 2003; Kanevskiy
et al., 2013). Schirrmeister et al. (2008) investigated the characteristics
of the Yedoma suite in northern Siberia and offshore islands and found that
gravimetric ice contents averaged 60–120 % in sequences of permafrost
soils. Although most of the studies on cryostructure are focused on its
physical aspects, cryostructure also has a role in soil biogeochemical
processes. As the unfrozen water in ice lenses or ice layers in the
cryostructure engage in heat and mass transfer in the winter (Romanovsky and
Osterkamp, 2000), unfrozen water can serve as a channel for biological
activity and thereby affect carbon cycling during winter months (Fahnestock
et al., 1999).
Classification of permafrost soils
Permafrost-affected soils are referred to as Cryosols in both the International Union of Soil Sciences (IUSS) World
Reference Base (WRB) (IUSS Working Group WRB, 2014) and the Canadian soil classification system (Soil
Classification Working Group, 1998), Gelisols in the US system (Soil Survey
Staff, 1999), and permagelic suborders in the Chinese system (Gong, 2001).
The Russian system has no equivalent for Cryosols/Gelisols. In the new
Russian system (Shishov et al., 2004), the presence of permafrost (within 1
or 2 m of the surface) and cryoturbation is taken into account at lower
taxonomic levels. Despite differences in nomenclature and the properties used
for classification, these systems share the basic requirement of the presence
of permafrost at a certain depth, generally within 2 m of the surface.
However, these systems differ substantially in their classification
hierarchy. In the WRB and Russian systems, the genetic lineage or soil
material type (such as organic vs. mineral) supersedes soil thermal regimes.
For example, organic soils (Histosols) key out first and those with
permafrost key out in the subclass as Cryic Histosols. But, in the Canadian
and US systems, Cryosols and Gelisols, respectively, key out first because
permafrost is recognized as the controlling factor in land-use interpretation
and ecosystem functions.
Examples of permafrost-affected soils from different areas:
(a) Glacic Aquiturbel formed in weathered shale with small patterned
ground features (polar desert), ice wedge at 58 cm, Ellef Ringes Island,
Nunavut, Arctic Canada; (b) Haploturbel associated with a nonsorted
circle, formed in weathered siltstone (High Arctic), Mould Bay, Arctic
Canada; (c) Ruptic–Histic Aquiturbel formed in Yedoma deposit under
nonacidic tundra, Mys Chukochi, Northern Yakutia, Russia; (d) Ruptic
Histoturbel associated with a nonsorted circle, formed in loess over Tertiary
glaciofluvial deposit under tussock tundra, Arctic Alaska; (e) a
well-drained Haploturbel formed in weathered sandstone under alpine tundra,
northern–central Qinghai–Tibet Plateau, China; (f) Histel formed
in wet Kobrecia meadow (note the limnic layer at 112 cm),
Qinghai–Tibet Plateau, China; (g) Histoturbel formed in late
Pleistocene glaciofluvial
deposit under boreal forest, northern Alaska; (h) Aquorthel formed
in loess deposit under boreal forest, central Alaska; and (i) Histel
formed in bog, central Alaska. Soils shown in (a), (b),
(f), (h), and (i) are formed in epigenetic
permafrost and soils shown in (c), (d), (e), and
(g) are formed in syngenetic permafrost.
Gelisols of the US classification system have three suborders: Histels,
Turbels, and Orthels. Histels are organic soils, Turbels are mineral soils
affected by cryoturbation, and Orthels are mineral soils lacking
cryoturbation (Soil Survey Staff, 1999). Examples of each great group in the
High and Low Arctic and boreal regions are shown in Fig. 3. Gelisols are
ubiquitous throughout the continuous permafrost zone of the northern
circumpolar Arctic region. In this zone, Gelisols commonly form in lowlands
with restricted drainage that favor near-surface permafrost, but Gelisols
also form in well-drained upland sites where the permafrost is deeper
(Rieger, 1983; Ping et al., 2004; Li et al., 2008). Within the discontinuous
and sporadic permafrost zones of the boreal regions, the slope and aspect of
uplands and mountains play a controlling role in distributions of permafrost
and Gelisols (Péwé, 1975; van Cleve et al., 1983; Ping et al.,
2005a). In the Southern Hemisphere, permafrost and Gelisols occur on the
Antarctic continent, on the sub-Antarctic islands, and in mountain areas
(Beyer et al., 1999, 2000; Blume et al., 1997; Bockheim, 1990; Campbell and
Claridge, 2004a, b; Simas et al., 2007; Gilichinsky et al., 2010). However,
the organic carbon content of permafrost-affected soils in the southern
circumpolar region is extremely low when compared to the northern permafrost
region (Bockheim, 1990; Claridge and Campbell, 2004), except for some
isolated areas affected by maritime climate in coastal Antarctica and sub-Antarctica
(Beyer et al., 1995). Hence, the focus of this review is on permafrost soils
of the northern permafrost region.
Modes of soil organic carbon accumulation particular to permafrost
soils
Similar to other soils, the deposition of surface litter, rhizodeposits, and
the turnover of belowground biomass are sources of organic carbon inputs to
SOM for permafrost soils. The organic matter inputs from the existing plant
community are largely deposited in the near-surface active layer. However,
one of the most striking differences between permafrost soils and other soils
is the large amount of organic carbon stored in the active layer and
perennially frozen portions of the soil relative to the annual production of
plant biomass. The northern circumpolar region contains about 30–40 % of
the global soil carbon pool to a depth of 3 m, yet the region covers just
15 % of the global land surface and supports only 10–20 % of the global
vegetation carbon pool (Jobbágy and Jackson, 2000; McGuire et al., 2009;
Hugelius et al., 2014; Köchy et al., 2014). The disproportionate ratio of
soil carbon to plant biomass carbon in the permafrost region is due to the
slow decomposition rates for organic matter under the low temperatures and
often
saturated (thus reducing) conditions in these soils (Hobbie et al., 2000;
Davidson and Janssens, 2006; Ping et al., 2008b, 2010). However, in
permafrost soils, other processes, such as crystal and massive ice formation,
freeze cracking, freeze jacking, diaparism, and gelifluction can result in
deformation and displacement of soil materials within the profile. The
deformation and displacement of soil materials caused by these general
processes are commonly referred to as cryoturbation (Williams and Smith,
1989; Bockheim and Tarnocai, 1998; French and Shur, 2010). Cryoturbation does
not cause direct input of organic carbon to soils; instead, it displaces and
redistributes SOM within the active layer of the soil profile. Over time and
with changing conditions, cryoturbated organic matter in the active layer can
become encased in permafrost by processes associated with either intermittent
burial or syngenetic permafrost growth (Schirrmeister et al., 2002a, 2011a;
Kaiser et al., 2007; Ping et al., 2008a, b). The unique feature of most
permafrost-affected soils is the storage of large quantities of this surface- and near-surface-produced SOM at depth – mainly in Turbels, which account
for 35 % of the areal extent and 46 % of the organic carbon stocks of all
Gelisols (Hugelius et al., 2014). However, cryoturbation has no (or only a
minor) effect on organic carbon storage in Histels, Orthels, and the large
areas of Histosols in the northern circumpolar region. Nevertheless, other
processes can also redistribute and stabilize carbon in these
permafrost-affected soils. For example, dissolved organic carbon that is
transported from the active layer to the permafrost table can accumulate
there due to “cryochemical precipitation” (Ostroumov et al., 2001;
Gundelwein et al., 2007).
Cryoturbation
Cryoturbation is the term used to describe the lateral and vertical
displacement of soil during seasonal and/or diurnal freezing and thawing
(French, 1988). The genesis of cryoturbation has been attributed to loading,
cryohydrostatic pressure and cryostatic pressure, and diaparism (Washburn,
1973; Rieger, 1983; French, 1988; Vandenberghe, 1988; Swanson et al., 1999).
However, Mackay (1980) diffused the commonly accepted cryostatic pressure
theory and proposed an alternative equilibrium model as a mechanism for
cryoturbation in hummocks. His equilibrium model was based on field and
laboratory evidence for cell circulation within the confines of bowl-shaped
frost tables under hummocks. He explained the upward displacement of material
in a hummock by a gravity-induced cell-like movement caused by the
freeze–thaw of ice lenses at both the top and bottom of the active layer.
The cell circulation occurs because the top and bottom freeze–thaw zones
have opposite curvatures, as described by Shilts (1978). Under this
mechanism, the period of greatest activity occurs in early summer at the top
of the active layer and in late summer at the bottom. In contrast, activity
induced by cryostatic pressure would be greatest during the early winter
freeze-back period, but Mackay (1980) could not find any theoretical,
experimental, or field evidence to support that mechanism for hummocks.
Hummocks are nonsorted circles that have uniform soil texture inside and
outside the circle and are one of the most widely distributed forms of
patterned ground in permafrost regions (Washburn, 1973, 1980). However,
cryoturbation also manifests in sorted circles. Kessler et al. (2001) used a
similar circulation model to describe a mechanism for sorted circle formation
in areas where a layer of soil underlies a uniform layer of stones. They
proposed that frost heave deforms the stone–soil interface and that this
causes local instabilities, where soil plugs begin to form. Through a variety
of transport mechanisms driven by freeze–thaw cycles, the soil plugs grow
upward as soil is drawn in from surrounding areas. The soil plugs develop
into sorted circles as they contact the ground surface, simultaneously
elevating a surrounding circle of stone or gravel. Once formed, sorted
circles are maintained by continued dynamic circulation within the soil and
stone domains.
Cryoturbated soil profiles of the thawed active layers of an earth
hummock (top) and a nonsorted circle (bottom). Note the distribution of
organic matter in the low micro-topographic positions and the insulating
effect of the organic horizon on thaw depths.
Indeed, the most striking cryoturbation features are associated with
patterned ground (especially nonsorted and sorted circles) and include
broken, involuted or warped horizons within the soil profile (Tarnocai and
Smith, 1992; Ping et al., 2008a, 2013). These features are caused by
differential frost heave processes in the soils. Commonly, circle centers are
devoid of vegetation or are only sparsely vegetated. Whereas, the areas outside
and at the margins of circles have tundra vegetation with thick organic
horizons (Fig. 4). Due to the lack of surface thermal insulation, the
mineral-dominated circles tend to thaw faster and deeper, resulting in deeper
active layers, but they also freeze faster. In contrast, areas outside the
circle, which are insulated by vegetation cover and organic layers, have
shallow active layers and tend to have higher soil moisture content during
thaw cycles (Shilts, 1978). These factors contribute to the development of
the bowl-like frost table central to the Mackay (1980) equilibrium model
described above. Because the mineral-dominated circles freeze faster, they
draw available water from the circle margins to form ice lenses, which
displace greater volumes of soil in circle centers. Meanwhile, freeze
contraction allows cracks or voids to develop around the circles. As a
consequence, greater heaving and deformation occurs within the circles than
in the adjacent insulated inter-circle soil. Frost heave in extreme cases can
cause the surface to heave as much as 20 cm in circles compared to 3 cm in
vegetated soils surrounding the circles (Romanovsky et al., 2008). This
differential frost heave often causes the organic mat and, in some cases, the
vegetative cover, to be entrapped in the cracks along the edge of patterned
ground features. From there, this material is frost-churned down with the circulation described by
Mackay (1980) from the depressed inter-circle zone to the lower active layer
under the circle, where it eventually can become part of the upper permafrost
(Ping et al., 2008a). Generally, nonsorted circles form in fine-textured
(loamy) soils, particularly silt loam soils. With increasing clay contents,
the centers of such circles tend to be more elevated. Circles with elevated
centers are termed “earth hummocks”, with hummock height directly related
to clay content (Ping et al., 2008a).
The importance of cryoturbation in soil formation has long been studied and
stressed by soil scientists and geocryologists. Classic examples of these
studies were carried out in the 1970s in the Canadian Arctic by
Tedrow (1974), Pettapiece (1974), Tarnocai and Zoltai (1978), and
Mackay (1980), and summarized by Tarnocai and Smith (1992). These researchers
found that the depth of the permafrost table is inversely related to the
surface microtopography; i.e., the active layer is deepest in the center of
the earth hummock and shallowest in the depressions around the hummocks. They
also noted that buried and discontinuous organic horizons occurred at depth,
and attributed this “burial” phenomenon to cryoturbation. Concurrently in
northeastern Russia, Makeev and Kerzhentsev (1974) also found similar
accumulation of cryoturbated organic matter at the permafrost table, which
often occurs at depths of 100 to 200 cm. Gubin and Lupachev (2008) found a
concentrated layer of organic matter in the upper part of the transient layer
under earth hummocks. Rieger et al. (1979) also recognized the occurrence of
“black streaks of frost-churned organic matter” at depth and discontinuous
surface organic horizons in association with nonsorted circles in the tundra
of Arctic Alaska. This work led to the creation of the subgroup of
“Ruptic–Histic” in the US taxonomic system (Soil Survey Staff, 1975).
Later, more detailed pedological studies in Arctic Alaska asserted that
cryoturbation is the primary factor controlling the amount of carbon
sequestered in the soils of tundra dominated by nonsorted circles (Bockheim
et al., 1998; Ping et al., 1998, 2008b; Bockheim and Hinkel, 2007). Thus, the
presence/absence of cryoturbated features was adopted as differentia for the
Turbel suborder of Gelisols in US taxonomy (Bockheim and Tarnocai, 1998,
Ahrens et al., 2004) and for the prefix of Cryosols in the WRB (IUSS Working
Group WRB, 2014). However, these phenomena are not limited to the continuous
permafrost zone. Cryoturbation and carbon translocation/sequestration
associated with sorted circles, stripes, and solifluction are also commonly
found in landscapes affected by periglacial processes, such as alpine
environments in the boreal regions of Alaska (Geisler and Ping, 2013),
Mongolia (Maximovich, 2004), and southern Siberia (Gracheva, 2004). Deeply
subducted organic masses are even found in rocky permafrost soils lacking
surface evidence of patterned ground (Jorgenson et al., 2013). Furthermore,
when permafrost is absent under current climate conditions, cryoturbated
profiles are often still present as remnants of past permafrost environments
(Munn, 1987; Rieger, 1983).
Redistribution of organic matter within the soil profile due to cryoturbation
is most common in the Arctic, where patterned ground processes are active –
such as sorted and nonsorted circles and ice-wedge polygons (Washburn, 1973;
Michaelson et al., 2008; Walker et al., 2008; Kanevisky et al., 2011;
Zubrzycki et al., 2013; Ping et al., 2008b, 2014). In early studies, SOC
stocks were measured only at shallower depths – mostly limited to the
rooting zones at depths less than 50 cm, due to logistics (Brown, 1969;
Everett and Brown, 1982). However, as concerns about global warming
increased, this raised fears that the organic carbon stored in
permafrost-region soils might become a source of rather than a sink for
atmospheric carbon (Oechel et al., 1993). Consequently, a series of studies
was conducted in the northern circumpolar regions to explore the depth
distribution of stored biogenic carbon in Gelisols (Michaelson et al., 1996,
2008, 2013; Ping et al., 1998, 2008b; Bockheim and Hinkel, 2007; Bockheim et
al., 1999; Tarnocai et al., 2009; Hugelius et al., 2010, 2013a, b; Strauss et
al., 2012, 2013). Generally, on gentle to moderate slopes of glaciated
uplands, SOM was cryoturbated to depths of mostly 80 to 120 cm. But,
cryoturbated SOM was found to reach depths of 3 m or more on exposed ridge
tops where vegetation is sparse and protection from snow cover during winter
is lacking (Michaelson et al., 1996), on floodplains (Shur and Jorgenson,
1998), in thaw-lake basins (Hinkel et al., 2003; Ping et al., 2014), and in
degraded permafrost in the boreal region (Jorgenson et al., 2013). Most of
these mineral soils with cryoturbated SOM and broken surface organic horizons
are classified as Turbels (Turbic Cryosols). But, in the poorly drained
valleys or basins among the hilly uplands, thick organic horizons
(> 40 cm) can build up due to fen or bog formation, and these soils are
classified as Histels (Cryic Histosols) (Ping et al., 2004, 2005a).
Ping et al. (2008b) measured SOC stores along a north–south transect through
five bioclimatic subzones from the High Arctic to the boreal regions in North
America as part of a larger study investigating the interrelationships
between patterned ground formation and vegetation zonation (Walker et al.,
2008). As the patterned ground transitioned from simple frost cracking in the
High Arctic to well-formed nonsorted circles in the Low Arctic, the
land-cover types changed from polar desert to tundra to shrub tundra, with increased dominance of vascular plants in the south (Walker et al., 2008). Ping
et al. (2008b) found that SOC stocks were directly related to biomass
production but that the proportion of cryoturbated carbon did not follow the
same trend. Rather, the proportion of cryoturbated carbon stored in the upper
permafrost reached a peak near the middle of the transect, where tussock
tundra dominates the landscape, and then dropped off further south, where
shrub tundra becomes dominant (Michaelson et al., 2008; Ping et al., 2008b).
This trend corresponds to reduced frost heave and cryoturbation under shrub
tundra because the taller vegetation canopy, thicker ground moss layer, and
increased snowfall provide greater thermal insulation (Rodionov et al., 2007;
Kade and Walker, 2008; Walker et al., 2008).
Deformation by massive ice formation
Ice-wedge polygons, with diameters ranging from a few meters to more than
20 m across, dominate the landscapes of the Arctic coastal plains, which are
widespread lowlands throughout the northern circumpolar region. Ice-wedge
polygons also dominate river floodplains, valleys, lowland areas and
thermokarst depressions in Arctic foothills and the sub-Arctic, e.g.,
interior Alaska, northern Canada, and central Yakutia, Russia (Péwé,
1975; French, 2007; Schirrmeister et al., 2011b; Strauss et al., 2012;
Zubrzycki et al., 2013). During ice-wedge development, soils on both sides of
the ice wedge are pushed apart and heaved to form rims on both sides,
creating a trough between the rims and over the ice wedge that delineates the
polygon (Fig. 2). The polygon center is less affected and remains flat or
slightly lower than the rims. At this developmental stage, the polygons are
flat- or low-centered, where the trough (and often the center) is wet during
the growing season. Three different main soil types have been observed to
develop across this micro-toposequence: (1) organic or organic-rich soils
over the ice wedge (generally less than 50 cm across) in the polygon trough
(Glacistels), (2) cryoturbated soils along the polygon rims (Histoturbels or
Aquiturbels), and (3) soils usually lacking cryoturbation in flat or low
polygon centers (Aquorthels, Historthels). However, with time, the ice wedges
can degrade, forming deep troughs (Jorgenson et al., 2006). Relatively,
polygon interiors then become high centered, surface cracking increases, and
greater cryoturbation associated with the cracks leads to formation of soils
(Histoturbels) with greater organic matter accumulations (Ping et al., 2011,
2014; Zubrzycki et al., 2013).
Intermittent burial and syngenetic permafrost
Intermittent burial by eolian, alluvial, colluvial, and lacustrine sediments
can lead to accumulation of significant SOC stocks in these depositional
environments (Shur and Jorgenson, 1998; Schuur et al., 2008; Tarnocai et al.,
2009; Grosse et al., 2011; Schirrmeister et al., 2008, 2011b; Zubrzycki et
al., 2013). These processes are often accompanied by syngenetic permafrost
growth, where the upward growth of the permafrost surface follows the
accumulation of sediments and peat layers at the surface (Shur, 1988).
Deltas and thaw lakes with ice-rich permafrost are abundant along the Arctic
coastal plains, such as at Barrow in northwestern Alaska (Hinkel et al.,
2005), the Colville Delta in northern Alaska (Shur and Jorgenson, 1998; Ping
et al., 2011) and the Lena River Delta in northern Russia (Boike et al.,
2013; Zubrzycki et al., 2013). In the Colville Delta, buried organic soils
are found as deep as 3 m (Shur and Jorgenson, 1998). In addition to Arctic
coastal plains, thaw lakes are also common in the sub-Arctic, such as
interior Alaska (Péwé, 1975), northern Canada (Sannel and Kuhry,
2011), and several areas in Russia (Grosse et al., 2013).
Thaw lakes and drained thaw-lake basins are prevalent on flat landscapes in
the Arctic and boreal regions where ground ice is sufficiently abundant to
allow the surface to thaw, collapse, and be filled with water to form shallow
(< 2 m) and deep (> 2 m) lakes. Thaw lakes are most abundant on
ice-rich, fine-grained deposits, such as abandoned floodplains, colluvial
lower slopes and basins, peatlands, and lowland loess deposits (Veremeeva and
Gubin, 2009; Grosse et al., 2013; Jorgenson, 2013). The age of thaw lakes can
be extremely variable with some thaw lakes in Yedoma persisting since the
late Pleistocene (Grosse et al., 2013, Kanevskiy et al., 2014), while others
are newly formed. Veremeeva and Gubin (2009) identified two stages of active
thaw-lake formation in Northern Russia; one during the early Holocene around
9000 to 8000 years BP and the second one during the late Holocene around
5000 to 4000 years BP. Formation of these lakes has been largely attributed
to thermokarst and thermal erosion processes related to past and modern
climate changes and more recently to land use. In addition, not all lakes in
permafrost regions are thaw lakes; some can simply be abandoned channel
lakes, inter-dunal lakes, or impoundments in depressions of undulating
surficial deposits (Jorgenson et al., 2006). In many Arctic coastal plain
regions, there are overlapping lakes and lake basins, which have been
attributed to “thaw-lake cycles” as short as 3000 years in northern Alaska
(Hinkel et al., 2003). However, Jorgenson and Shur (2007) argued that ice
aggradation and degradation are generally too slow to account for this. They proposed that the
process is less cyclical and more evolutionary throughout most of the Arctic
coastal plain of Alaska, with secondary thaw lakes being a common outcome of
developmental sequences that depend on factors such as topography, reworking
of surficial deposits, drainage, and ground ice conditions. In some cases,
soils formed in drained-lake basins reflect multiple thaw episodes, as
evidenced by the presence of multi-layered organic-mineral horizons with
cryoturbated organic matter to depths of 2–4 m (Fig. 5) (Ping et al., 2014;
Kanevskiy et al., 2014). Thaw lakes can also form in landscapes where there
are no ice wedges. For example, thaw lakes are commonly associated with the
degradation of palsas and peat plateaus in sub-Arctic peatlands, such as
those observed in the northern boreal region of Alaska (Brown and Kreig,
1983). As a consequence of thaw-lake dynamics, large SOC stocks can
accumulate – as high as 90 kg C m-2 in profiles up to 3 m deep
(Ping et al., 2011). Much of the organic matter in the central portions of
drained-lake basins is derived from limnic sediments comprised of algal
material and detrital peat eroded from collapsing lake shores (Jorgenson,
2013).
Gelisol profile from the Arctic coastal plain of northwestern Alaska
exhibiting a complex stratigraphy developed through a sequence of different
environments coupled with cryoturbation caused by freeze/thaw/heave forces.
The strata suggest that (1) old coastal plain sediments below 80 cm thawed
and collapsed into a thaw lake as indicated by the mixed-up sediments with
scattered fragmental peat chunks; (2) after drainage, an organic mat
developed at 50 to 80 cm; (3) this mat was flooded or collapsed within the
lake basin and was covered by limnic silts at 30–50 cm; (4) after a second
drainage, another organic mat covered the limnic silts at 15 to 30 cm;
(5) windblown eolian silt covered this mat at 10–16 cm; and (6) a final
organic mat developed at the surface. The fluctuating and rising permafrost
table disrupted the underlying organic mats.
Extremely ice-rich silt deposits of Pleistocene age, known as Yedoma or Ice
Complex, are particularly noteworthy because of the significant stocks of
“fossil” organic carbon (Schirrmeister et al., 2011a). These deposits
(average depth ∼ 20–25 m) developed during the late Pleistocene in
unglaciated areas of the Beringia, including northern Siberia, Arctic and
Interior Alaska, and the Yukon Territory, Canada (Zimov et al., 2006; Froese
et al., 2008; Schirrmeister et al., 2002a, b, 2011b; Kanevskiy et al., 2011;
Vonk et al., 2012). Yedoma deposits are polygenetic accumulations of eolian
or water-borne origin that settled in the interglacial periods under
syngenetic conditions when permafrost was already present or in an
environment that already experienced sub-zero temperatures for most of the
year (Kanevskiy et al., 2011; Schirrmeister et al., 2002a, b, 2011b, 2013;
Strauss et al., 2012). The volumetric ice content of Yedoma can be as high as
80–90 %, including both segregated ice within soil and massive ice wedges
that can reach depths of 50 m or more (Kanevskiy et al., 2013). Mineral
horizons in Yedoma typically contain over 70 % silt, although sandy
deposits also occur. Fast sedimentation rates and simultaneous freezing often
buried and preserved organic matter that is potentially more decomposable
than that of non-permafrost mineral soils. The concentrations of organic
carbon in these deposits average 1–5 % (Strauss et al., 2013), but
concentrations can reach 15 % in peaty horizons within some deposits
(Schirrmeister et al., 2002b). Strauss et al. (2013) estimated an organic
carbon pool of 83 (+61/-57) Pg C for the Yedoma deposits of Siberia and
Alaska, with an average organic carbon density of 19
(+13/-11) kg C m-3 that was reduced to 10
(+7/-6) kg C m-3 when ice-wedge volumes were included in the
estimate.
The large ice content of Yedoma makes this soil particularly vulnerable to
climate warming, and deep thermokarst lakes are common on the Yedoma
landscape. Some thermokarst lakes eventually turn into drained basins, or
“alases”, a Yakutian term for thermokarst basins with steep slopes and flat
floors. Alas formation passes through a sequence of landforms – lake, swamp,
wet meadow, and grassland – with associated soil types (Czudek and Demek,
1970; Veremeeva and Gubin, 2009; Desyatkin, 2008). In the intermediate stage,
permafrost organic soils (Histels) were found in alases at Duvaany Yar, the
extensively studied Yedoma formation along the Kolyma River upstream from
Cherskiy (Smith et al., 1995), and in northern Alaska (Kanevskiy et al.,
2011). Morgenstern et al. (2013) suggested that alas formation was most
extensive across the coastal lowlands of eastern Siberia due to abrupt
thaw-lake drainage during the early Holocene. However, on a local scale and
often within primary alases, more dynamic but less intensive thermokarst
processes during the late Holocene have shaped the modern thermokarst
terrain, which continues to evolve in response to present-day climatic and
land-use changes.
Quantity of organic carbon in permafrost soils
Major efforts to estimate global SOC storage during the twentieth century
yielded values ranging from 400 to 9120 Pg C stored in soils worldwide
(Amundson, 2001). Amundson (2001) calls attention to the wide variation in
the amounts of soils data, analyzed depths, and summation methods used in
these early analyses of the global SOC pool. Some estimates applied overall
soil averages to the entire land surface (Rubey, 1951), others averaged SOC
storage according to ecosystem types (Schlesinger, 1977; Post et al., 1982;
Jobbágy and Jackson, 2000), while others were based on soil maps and soil
taxonomy (Atjay et al., 1979; Bohn, 1982; Eswaran et al., 1993). However, the
influence of permafrost on SOC storage was not directly recognized until
relatively recently in studies such as Tarnocai and Lacelle (1996),
Tarnocai (1998), Michaelson et al. (1996, 2013), Ping et al. (2008b),
Hugelius et al. (2010), Bliss and Maursetter (2010), Johnson et al. (2011),
and Mishra and Riley (2012) for Alaska and North America; Stolbovoi (2002),
Kuhry et al. (2002), Rodionov et al. (2007), Hugelius (2012), Schirrmeister
et al. (2011a), and Zubrzycki et al. (2013) for Eurasia; and Tarnocai et
al. (2009) for the entire northern circumpolar region. Indeed, the Gelisol
order of Soil Taxonomy was only established in 1999 (Soil Survey Staff,
1999).
As more soil pedon data became available, especially pedons evaluated for SOC
storage, estimates of the permafrost-region SOC pool increased substantially
– primarily due to a fuller accounting of deeper cryoturbated carbon and
permafrost carbon stocks. As an example, Jobbágy and Jackson (2000)
estimated that only 294 Pg C were stored at 0–3 m depth in the boreal
forest and tundra biomes combined, whereas Tarnocai et al. (2009) estimated
1024 Pg C stored at 0–3 m for the northern circumpolar permafrost region.
Although the estimate by Tarnocai et al. (2009) was based on about 1700 total
pedons, only 25 of these pedons encompassed the full 0–3 m depth. At
present, the largest spatially distributed soils data set from the northern
circumpolar region is the Northern Circumpolar Soil Carbon Database (NCSCD),
which was initially developed for the Tarnocai et al. (2009) estimate. The
NCSCD was recently updated with new data for soil depths of 1–3 m as well
as deeper Yedoma and deltaic-alluvium deposits (Hugelius et al., 2013a, b).
Hugelius et al. (2014) used the updated NCSCD to produce the current best
overall estimate of 1307 Pg C stored in permafrost region soils (0–3 m
depth together with carbon in the deeper Yedoma and deltaic deposits). Of
this total, roughly two-thirds (822 Pg C) is estimated to occur in permafrost.
However, the overall estimate has a substantial uncertainty range
(1140–1476 Pg C) and significant data gaps. Among the reasons for the
large uncertainty range are the remoteness and vastness of the region (17.8×106 km2 or 16 % of global soil area). In the updated NCSCD
database, the carbon estimates at 0–1 m depth were derived from 1778
pedons. However, deeper soils were represented by only 341 Gelisol pedons and
177 pedons of non-permafrost soils at 1–2 m and by 246 Gelisol pedons and
105 non-permafrost soil pedons at 2–3 m (Hugelius et al., 2014). Although
these numbers constitute a significant improvement over the deep soil pedons
available for the Tarnocai et al. (2009) estimate, the data are still quite
limited considering the overall size of the region.
In addition to the low sampling density due to difficulties associated with
remote access and the vast extent of the circumpolar permafrost region, there
are technical and sampling challenges unique to Gelisols. These challenges
can contribute to increased variability and decreased confidence in SOC
storage calculations and the overall carbon stock estimates for Gelisols. However, these issues
were not accounted for by Hugelius et al. (2014) in their uncertainty
estimates. Two of the more significant technical challenges are the lack of
measured soil bulk densities and differences in the methods of SOC
determination. Both of these challenges are common in most soil databases,
necessitating the estimation of these key soil properties, which are essential
to the calculation of SOC storage (Tarnocai et al., 2009; Johnson et al.,
2011; Hugelius et al., 2012; Mishra and Riley, 2012).
Modern high-temperature dry-ignition methods provide direct measurements of
soil carbon. But, prior to 2000, SOC was commonly estimated indirectly by
using variations of the chemical acid-dichromate oxidation method (the
Walkley–Black method), which required calibration or estimation of a
recovery factor for the method and group of soils being analyzed. In general,
recovery factors have been derived for soils from temperate and warmer
environments and have varied substantially. For example, Amacher et
al. (1986) reported recovery factors ranging from 0.46 to 0.87 with a mean of
0.71, which is close to the recovery factor of 0.77 used for the US
Department of Agriculture Natural Resources Conservation Service (USDA-NRCS)
Alaska soil pedon data set (Soil Survey Staff, 2014). Working with this data
set, Michaelson et al. (2013) found that a recovery factor of 0.77
overestimated SOC in Alaska Gelisols by an average of 12 % (R2=0.98)
when compared to analysis by dry-ignition methods. Soil carbon values for
other soil orders found in the Alaska data set were overestimated by
3–18 %. Yet, these data have been commonly used for pedon carbon
assessments of permafrost regions (Tarnocai et al., 2009; Hugelius et al.,
2014). Thus, new pedotransfer functions were developed and used by Michaelson
et al. (2013) to update the Alaska USDA-NRCS data set because SOC data for
57 % of the pedons were determined only by the chemical
dichromate-oxidation method. The problem of overestimating organic carbon
concentrations in cold-region soils, such as those found in Alaska, is likely
due to the application of recovery factors that were developed for more
highly oxidized soils. Overall, the organic matter in cold-region soils tends
to be less decomposed and less oxidized compared to more temperate soils.
Given a higher proportion of more readily oxidizable (decomposable) organic
matter, many cold-region soils could be more completely oxidized by the
dichromate-oxidation method, leading to overestimation when the standard
recovery factor is used. Higher concentrations of Fe2+ in cool wet soils
might also increase estimates by dichromate-oxidation methods. More research
into individual soils would be needed to sort these factors out. But, data
users should be aware of these aspects of data sets, even though newer data
sets are more likely to contain soil carbon data determined by the more
direct high-temperature dry-ignition methods.
As an additional complication, SOC concentrations are commonly used to
estimate and fill in missing bulk density data. But, for Gelisols, the
correlation between SOC concentration and bulk density is poor in the
permafrost layer (R2=0.46) because of the large and variable ice
contents (Michaelson et al., 2013), especially in the often carbon-rich
transient layer of the upper permafrost. Ping et al. (1998, 2013) outline
state-of-the-art protocols for sampling permafrost soils that address the
issues of sampling across pedon cryogenic patterns of variability,
calculation of SOC storage in cryoturbated profiles, and the sampling and
measuring of bulk density in frozen Gelisol horizons.
Another major issue affecting estimates of the permafrost-region SOC pool is
fine-scale vertical and horizontal heterogeneity. Estimates of the vertical
distribution of organic carbon stocks in permafrost region soils are
extremely important for predicting the amount of SOC that will become
vulnerable to mineralization upon top-down thaw or losses due to increased
incidence of fire (Koven et al., 2011; Harden et al., 2012). Yet, the content
and distribution of organic carbon within the soil profile vary widely among
Histels, Turbels, and Orthels (Harden et al., 2012), making this variation a
major source of uncertainty in current estimates. Furthermore, as described
above, patterned ground features can exhibit various combinations of these
suborder soils, resulting in large local and landscape heterogeneity. A
number of studies have looked at the SOC content of patterned ground
features, for example frost boils (Dyke and Zoltai, 1980; Walker et al.,
2004; Kaiser et al., 2005; Michaelson et al., 2012), circles (Kimble et al.,
1993; Hallet and Prestrud, 1986), stripes (Walmsley and Lavkulich, 1975;
Horwath et al., 2008), ice-wedge polygons (Ping et al., 2014; Zubrzycki et
al., 2013), and earth hummocks (Kimble et al., 1993; Landi et al., 2004). In
addition to these pedon-scale studies, explicit efforts to account for and
incorporate local-scale spatial heterogeneity into sampling designs and
upscaling approaches are being explored for the estimation of SOC stocks on
landscape or regional scales (Horwath et al., 2008; Hugelius and Kuhry, 2009;
Zubrzycki et al., 2013). Hugelius et al. (2011) used the highest resolution
upscaling to date of any permafrost SOC study (2 m resolution) and compared
local scale heterogeneity vs. regional scale upscaling. Hugelius (2012) also
provided an in-depth discussion regarding the topic of upscaling accuracy and
spatial resolution in upscaling.
Characterization of the quality and decomposability of organic matter
in permafrost soils
Soil organic matter is a complex and heterogeneous mixture derived mostly from
plant litter inputs and microbial residues that exists in various stages of
decomposition and associations with soil minerals (Baldock and Skjemstad,
2000; Sutton and Sposito, 2005; Kelleher and Simpson, 2006; Jastrow et al.,
2007; Lehmann et al., 2008). Traditionally, the “quality” of SOM (and by
inference its intrinsic potential to be further decomposed, transformed and
mineralized) has been evaluated by assessing its molecular composition and
by applying some type of physical, chemical, or biological fractionation
approach to partition the bulk SOM pool into “labile” vs. “recalcitrant”
or “stable” forms (Kleber and Johnson, 2010; von Lützow et al., 2007;
Simpson and Simpson, 2012). Although the chemical composition of SOM can
affect the rate of decomposition, it is now recognized that (1) readily
degradable “labile” carbon forms can be stabilized in soil by a variety of
mechanisms or conditions that physically or chemically limit microbial
access or that impact microbial activity, and (2) even the most
“recalcitrant” carbon forms can be mineralized given the “right”
conditions (von Lützow et al., 2006; Marschner et al., 2008; Kleber,
2010).
Thus, while precise definitions and standardized measurements of “labile”,
“recalcitrant”, or “stable” SOM pools are not possible, these terms are
nonetheless useful in a relative sense for comparing the intrinsic
degradation state or potential decomposability of SOM – particularly for
permafrost-affected soils. This is because the dominant factors limiting
decomposer activity in permafrost soils are the cold and often wet anoxic
environments (Davidson and Janssens, 2006). In addition, low soil pH can
limit decomposition rates in some soils (Grosse et al., 2011). Hence, in
permafrost soils, organic matter is often preserved in a relatively
undecomposed state, both in surface horizons and in buried forms (e.g., peat
deposits, cryoturbated organic matter). In addition, much of this organic
matter is uncomplexed or only poorly associated with soil minerals (Diochon
et al., 2013; Höfle et al., 2013). Furthermore, subzero microbial
activity in unfrozen water films can lead to accumulation of easily
decomposable, soluble microbial by-products in frozen soils. Plus, soluble
organics produced in surface horizons (e.g., root exudates) might accumulate
due to drainage limitations imposed by the permafrost (Michaelson et al.,
1998; Mikan et al., 2002; Michaelson and Ping, 2003). Alternatively, in some
soil environments, cryochemical precipitation of dissolved organic carbon
might result in the deposition of rather stable SOM near the permafrost table
(Gundelwein et al., 2007).
Because so much SOM is currently stabilized simply by environmental
conditions, most efforts to evaluate the quality of organic carbon stored in
permafrost region soils have focused on characterizing the chemical
composition of the organic matter (e.g., Dai et al., 2002a; Turetsky et al.,
2007; Pedersen et al., 2011). Chemical characterization studies have run the
gamut of methodologies in efforts to assess the current state of degradation
– from wet chemical fractionations (acid/base extractions, molecular
biomarkers, and water or solvent extractions; e.g., Uhlířová et
al., 2007; Paulter et al., 2010; Hugelius et al., 2012) to various types of
spectrochemical analysis (nuclear magnetic resonance spectroscopy,
pyrolysis–gas chromatography/mass spectrometry, mid-infrared spectroscopy;
e.g., Dai et al., 2002a, b; Andersen and White, 2006; Waldrop et al., 2010;
Pedersen et al., 2011; Pengerud et al., 2013). Even though organomineral
associations might play a role in the relative persistence of SOM upon
thawing and warming (Davidson and Janssens, 2006), relatively few studies
have employed physical fractionations to characterize the amount of SOC
stabilized by mineral associations or by aggregation (e.g., Dutta et al.,
2006; Xu et al., 2009a, b; Höfle et al., 2013).
These varied techniques provide different perspectives of the composition of
organic matter stored in permafrost soils and its potential resistance to
decomposition. In general, wet chemical and physical fractionations enable
estimates of pool sizes for SOM of different qualities, whereas
spectrochemical analyses provide more specific information on SOM chemistry.
From the perspective of the permafrost region as a whole, application of
these methods to a range of soils and horizons provides similar overall
conclusions for permafrost soils. In general, wet chemical and
spectrochemical methods indicate that (1) the bulk of the organic matter in
permafrost-affected soils is relatively less decomposed or “humified” than
that of more temperate soils; (2) plant-derived materials are often more
prevalent than microbial residues; and (3) environmental conditions have
promoted the preservation of relatively labile SOM constituents (e.g., Ping
et al., 1997; Dai et al., 2002a, b; Pedersen et al., 2011; Hugelius et al.,
2012; Paulter et al., 2010). Physical fractionation approaches have provided
comparable information. For example, Diochon et al. (2013) found that lightly
decomposed particulate organic matter (POM) was a greater component (usually
> 30 %) of the total SOC in active-layer and permafrost horizons of
Canadian Turbic Cryosols than is typical for temperate soils. For similar
horizons in two Alaskan Turbels, Xu et al. (2009a, b) reported even greater
proportions of total SOC (> 70 %) in the POM fractions, and analysis of
this material by pyrolysis–gas chromatography/mass spectrometry indicated
that it was only lightly decomposed. Moreover, Höfle et al. (2013)
concluded that organomineral associations and aggregation are of lesser
importance for SOM stabilization in permafrost soils than in temperate and
tropical soils. Indeed, this is not surprising, because most clay minerals in
Arctic tundra soils are inherent from the parent material rather than
pedogenic (Borden et al., 2010). The lack of reactivity of these clay
minerals was further demonstrated by the lack of correlation between clay
content and cation exchange capacity (Ping et al., 2005b).
More direct assessments of the potential decomposability of SOM have been
made by measuring SOC mineralization in laboratory incubation studies (e.g.,
Michaelson and Ping, 2003; Lee et al., 2012; Elberling et al., 2013;
Knoblauch et al., 2013). To date, most incubation studies have been
relatively short term (weeks to a few months) and thus are assessments of the
most bioavailable components of SOM. Some studies have combined chemical or
physical characterizations with incubations in an effort to relate SOM
composition to decomposition rates or to identify indices of decomposability
(e.g., White et al., 2002, 2004; Weintraub and Schimel, 2003; Waldrop et al.,
2010; Diochon et al., 2013; Paré and Bedard-Haughn, 2013; Pengerud
et al., 2013; Treat et al., 2014). In general, SOM of higher quality (less
decomposed and rich in polysaccharides and proteins) is positively related to
mineralization rates.
Long-term (> 1 year) incubation studies provide better assessments of the
integrated effects of multiple SOM chemistries and stabilization mechanisms
on mineralization rates and also can inform the estimates of turnover rate
functions used in process models. But, to date, the number of studies of this
length for permafrost soils has been quite limited (Schädel et al.,
2014). Initial or short-term mineralization is often related to the amount of
dissolved or water-extractable organic carbon. But for
longer-term
incubations, the mineralization observed in thawed permafrost soils has been related to
total SOC concentration and to differences in SOM quality at the time of
permafrost incorporation (Dutta et al., 2006; Lee et al., 2012; Knoblauch et
al., 2013). In a synthesis of eight long-term aerobic incubation studies
(> 1 year) including 121 samples from upland sites in 23 high-latitude
ecosystems, Schädel et al. (2014) estimated SOC pool sizes and turnover times
from a three-pool decomposition model. They projected a 20–90 % loss of
initial SOC within 50 incubation years at 5 ∘C, with greater losses
occurring for soils with higher carbon to nitrogen ratios. However, the
importance of oxygen availability for mineralization rates was demonstrated
by Elberling et al. (2013). At the end of their 12.5 year incubation study,
only 9 % of the initial organic carbon in a saturated upper permafrost soil
from a wet grassland site was mineralized compared to 75 % of initial
carbon lost when the same soil was drained before incubation.
In addition to intrinsic soil differences (i.e., SOM quality and mineral
composition) and incubation time, the outcome and conclusions drawn from
incubation studies are highly dependent upon experimental conditions
including sample handling/disturbance, temperature, moisture conditions,
oxygen availability, and experimental additions of labile carbon substrates
and nutrients. For example, Dai et al. (2002b) found that the most
bioreactive compound was polysaccharides at 4 ∘C, whereas at
25 ∘C, more resistant fractions such as lignin were consumed. Such
shifts in substrate utilization at higher temperatures might reflect changes
in the active microbial community that may not be realistic under most field
conditions. In another case, Wild et al. (2014) showed that the
mineralization rates of surface organic soil, mineral subsoil, and cyroturbated organic material in the subsoil (all from the active layer) were each
differentially limited by the availability of labile organic carbon
substrates or nitrogen. Importantly, however, many of the conditions
impacting mineralization rates in incubation studies are indicative of
factors responsible for variations in overall SOC storage, turnover times,
and the intrinsic relative degradation state of SOM on subregion, landscape,
and local scales (e.g., Kaiser et al., 2007; Hugelius et al., 2012; Paré
and Bedard-Haughn, 2013; Pengerud et al., 2013).
Role and future of permafrost soils in climatic change
Climatic change at high latitudes is causing region-wide warming, hydrologic
changes, and other related disturbances (e.g., fires) that are triggering
widespread degradation and thawing of permafrost with potential global
impacts (Jorgenson et al., 2010; Romanovsky et al., 2010; Rowland et al.,
2010). One of the most likely and important consequences of sustained warming
in circumpolar regions is the thawing of permafrost soils and the subsequent
release of carbon as carbon dioxide and methane to the atmosphere due to
enhanced microbial mineralization of previously frozen SOC stocks (Oechel et
al., 1993; Zimov et al., 1993; Goulden and Crill, 1997; Melillo et al., 2002;
Eliasson et al., 2005; Zhuang et al., 2007; Schuur et al., 2009, 2011).
Model simulations of carbon losses from thawing permafrost are highly
uncertain and vary widely – with predictions of cumulative net transfers to
the atmosphere ranging from 7 to 250 Pg C by 2100, 121 to 302 Pg C by
2200, and 180 to 380 Pg C by 2300 for intermediate to high fossil fuel
emission scenarios (Zhuang et al., 2006; Koven et al., 2011; Schaefer et al.,
2011; Burke et al., 2012; MacDougall et al., 2012; Schneider von Deimling et
al., 2012). Such large releases of greenhouse gases are expected to have a
positive feedback to Earth's atmosphere, leading to further warming (Schuur
et al., 2008, 2011; Schaefer et al., 2011; Burke et al., 2012; MacDougall et
al., 2012; Schneider von Deimling et al., 2012). But, there are large
uncertainties in model predictions of carbon–climate feedbacks caused by
warming of northern circumpolar regions (Koven et al., 2011; Burke et al.,
2012). Furthermore, a discrepancy exists between the baseline SOC stock
estimates generated by Earth system models and observation-based estimates of
SOC stocks for the permafrost region (Mishra et al., 2013). Resolving these
issues will require improved empirical estimates as well as better model
representations of the unique processes controlling the formation and
stabilization of SOC stocks in permafrost regions (Mishra et al., 2013;
Zubrzycki et al., 2014).
In addition to the uncertainties surrounding current estimates of SOC stocks,
even less is known about the potential vulnerability of organic carbon stored
in permafrost soils to climatic change, and how this varies across land-cover
classes or soil types within different eco-regions (Schuur et al., 2008;
Kuhry et al., 2010, 2013). Upon thawing, the initial intrinsic degradation state of
previously frozen SOC pools depends on the type of permafrost formation, the
origin and chemistry of the SOM, and the extent of mineralization and
transformation that occurred before these materials were incorporated into
permafrost (Hobbie et al., 2000; Kuhry et al., 2009; Hugelius et al., 2012;
Knoblauch et al., 2013). Often, SOM stored in permafrost has undergone some
level of decay before its incorporation in perennially frozen horizons
(Pedersen et al., 2011; Hugelius et al., 2012). This may be particularly true
for SOM stabilized in epigenetic permafrost deposits (permafrost formed after
deposition of soil material) (Schuur et al., 2008). But, in cryoturbated
soils or syngenetic permafrost deposits (permafrost formed more or less
concurrently with deposition of soil material, such as Yedoma or some peat
deposits), relatively undecomposed organic materials were sometimes rapidly
buried and frozen (Zimov et al., 2006; Sannel and Kuhry, 2009). Even in
epigenetic permafrost, the extent of SOM degradation can be limited if anoxia
or other factors constrained decomposition before permafrost formation
(Hugelius et al., 2012).
Thus, there is a need to develop biogeochemical indicators that reflect
differences in the genesis and past history of SOM before incorporation into
permafrost and that represent its intrinsic relative degradation state (Kuhry
et al., 2010; Hugelius et al., 2012). With such indicators, maps of the
spatial and vertical distributions of SOC stocks could be coupled with an
indication of the potential decomposability of these carbon stocks. At the
same time, additional targeted mechanistic and process-based investigations
are needed to better understand the complex interactions and feedbacks among
SOM, vegetation, decomposers, nutrients, and hydrology induced by climatic
changes. Taken together, this information would substantially improve
observationally based predictions of the impacts of soil warming and
permafrost thawing on SOC stocks and their future dynamics as well as
contribute significantly to the parameterization, calibration, and validation
of regional and Earth system models (Schuur et al., 2008; Burke et al., 2012;
Hugelius et al., 2012; Mishra et al., 2013; Zubrzycki et al., 2014).
Conclusions
Soils in the permafrost regions have received considerable recent attention
because of their disproportionally high carbon stocks in comparison to soils
of non-permafrost regions and because of the attendant implications for
global climate change. One of the most likely and important consequences of
sustained warming in circumpolar regions is the thawing of permafrost coupled with subsequent effects on soil hydrology and resulting transformations of the
permafrost landscape. Increased paludification is expected due to increased
thermokarsting on the coastal plains and lowlands and due to the shifting of
water, carbon and nutrients from uplands to basins or valleys. In addition to
elevated fluxes of carbon dioxide and methane to the atmosphere, there will
be a shifting of the proportional releases of these two gases, with increased
methane emissions likely occurring in areas where the spatial extent of
wetland soils increase. These changes can be evaluated through landform
modeling based on newly developed remote sensing technologies such as
high-resolution lidar. Meanwhile, the vulnerability of organic carbon stocks
in permafrost soils to changing climatic conditions will depend on the
interactions of SOM composition with numerous other controlling factors that
also are likely to respond to climatic changes (such as temperature,
hydrology, nutrient availability, new carbon inputs from changing plant
communities, and changing associations with soil minerals). These
interactions will affect the integrated activities of the microbial
community, as well as the physical access of decomposers and their enzymes to
thawed carbon pools. All of these factors and interactions can only be
assessed through a combination of laboratory studies, integrated
observational and manipulative field studies, geospatial upscaling and
mapping of SOC stocks and indicators of their decomposability, as well as
ecosystem, regional, and Earth system modeling studies.
Acknowledgements
This material is based upon work supported by the US Department of
Agriculture NIFA Hatch funds and by the Alaska State Soil Classification
project (CLP and GJM), and by the US Department of Energy, Office of Science,
Office of Biological and Environmental Research under contract no.
DE-AC02-06CH11357 (JDJ).Edited by: D. Weindorf
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