Boreal ecosystems store large quantities of carbon but are increasingly vulnerable to carbon loss due to disturbance and climate warming. The boreal region in Alaska and Canada, largely underlain by discontinuous permafrost, presents a challenging landscape for itemizing carbon sources and sinks in soil and vegetation. The roles of fire, forest succession, and the presence (or absence) of permafrost on carbon cycle, vegetation, and hydrologic processes have been the focus of multidisciplinary research in boreal ecosystems for the past 20 years. However, projections of a warming future climate, an increase in fire severity and extent, and the potential degradation of permafrost could lead to major landscape and carbon cycle changes over the next 20 to 50 years. To assist land managers in interior Alaska in adapting and managing for potential changes in the carbon cycle we developed this review paper by incorporating an overview of the climate, ecosystem processes, vegetation, and soil regimes. Our objective is to provide a synthesis of the most current carbon storage estimates and measurements to guide policy and land management decisions on how to best manage carbon sources and sinks. We surveyed estimates of aboveground and belowground carbon stocks for interior Alaska boreal ecosystems and summarized methane and carbon dioxide fluxes. These data have been converted into similar units to facilitate comparison across ecosystem compartments. We identify potential changes in the carbon cycle with climate change and human disturbance. A novel research question is how compounding disturbances affect carbon sources and sinks associated with boreal ecosystem processes. Finally, we provide recommendations to address the challenges facing land managers in efforts to manage carbon cycle processes. The results of this study can be used for carbon cycle management in other locations within the boreal biome which encompasses a broad distribution from 45° to 83° north.

The boreal forest biome is the largest terrestrial biome on Earth, extending roughly from 45° to 83° N, and consisting primarily of coniferous forests with thick organic soils. Much of the boreal biome is underlain by permafrost. Permafrost, defined as any substrate remaining below 0°C for more than two consecutive years, can persist where mean annual air temperatures are as high as +2°C. This is due to localized differences in the soil thermal regime, which is influenced by topography, slope, aspect, hydrology, winter snowfall, ground ice content, soil texture, plant cover, and fire history (Osterkamp and Romanovsky, 1999; Hinzman et al., 2003a; Jorgenson and Osterkamp, 2005; Myers-Smith et al., 2008). Regional temperatures are not low enough to sustain permafrost everywhere (Schuur et al., 2008). Permafrost affects the vertical movement of water and nutrients through the soil column, so its presence can markedly affect surface soil and vegetation processes that are major aspects of the carbon cycle (Petrone et al., 2006; Walvoord and Striegl, 2007).

Climate warming is expected to have pronounced effects on high latitude ecosystems, especially in locations underlain by relatively warm, discontinuous permafrost such as interior Alaska (Arctic Climate Impact Assessment, 2005; White et al., 2007). Large areas of interior Alaska permafrost now show signs of degradation (Racine and Walters, 1994; Jorgenson et al., 2001; Osterkamp, 2007; Osterkamp et al., 2009) and these are expected to continue to degrade with further climate warming (Marchenko et al., 2008). Changes in permafrost distribution have dramatically affected ecosystems across the boreal and arctic regions through widespread drying in some regions (Riordan et al., 2006; Roach et al., 2011) and wetland expansion in others (Racine and Walters, 1994; Osterkamp and Jorgenson, 2006).

Future climate scenarios project a roughly 5°C increase in mean annual air temperatures for the Alaskan interior over the next 80 years (Chapman and Walsh, 2007). This will be enough to initiate permafrost degradation in many areas (Osterkamp and Jorgenson, 2006; Marchenko et al., 2008), and is expected to lead to major changes in vegetation (Potter, 2004; Walker et al., 2006; Wolken et al., 2011), biological and physical processes in soils, and hydrology (Tarnocai and Campbell, 2002; Davidson and Janssens, 2006; Allison and Treseder, 2011; Johnson et al., 2011; Sierra et al., 2010; Douglas et al., 2013). Changes in permafrost extent and stability will alter carbon source–sink dynamics and provide a management challenge for future planning scenarios. There is great concern that permafrost thawing will yield increased carbon emissions from arctic and subarctic landscapes and provide a positive feedback to global climate warming (Frey and Smith, 2005; McGuire et al., 2009; Koven et al., 2011; Schaefer et al., 2011; Schuur and Abbott, 2011).

The goal of this review is to assist land managers in adapting and managing for potential changes in the interior Alaska carbon cycle by incorporating an overview of the climate, ecosystem processes, vegetation types, and soil regimes of the boreal biome in interior Alaska. Our area of interest is the northern portion of the Tanana River Valley spreading from west to east between the Alaska Range to the South and the hills north of Fairbanks (Figure 1; Van Cleve and Viereck, 1983). This area encompasses 15,000 square miles (3.8 million hectares), slightly larger than Maryland. The main federal government land managers include the Department of Defense (primarily the Army) and the National Park Service. State of Alaska lands include forest, recreation, and wildlife areas. The remaining lands are owned either privately or by Alaska Native Corporations. Though our main focus is the region of interior Alaska the studies and information presented in this paper are applicable to boreal regions in Canada and Alaska.

Figure 1
A map of our focus region in interior Alaska.

Boundaries of the major state and federal government landowners. The majority of the non-Federal and non-State of Alaska lands are owned privately or by Alaska Native Corporations.

Figure 1
A map of our focus region in interior Alaska.

Boundaries of the major state and federal government landowners. The majority of the non-Federal and non-State of Alaska lands are owned privately or by Alaska Native Corporations.

Our objective is to provide a synthesis of the most current carbon source and sink estimates and research results to guide policy and land management decisions on how to manage carbon stocks in interior Alaska, and potentially in other boreal landscapes. A number of thorough review articles exist on the carbon sequestration potential of forest soils (Goodale et al., 2002; Lal, 2005) and the vulnerability of soil organic carbon (McGuire et al., 2010; Grosse et al., 2011; Harden et al., 2012), permafrost carbon (Schuur et al., 2008; Schaefer et al., 2014), and the carbon cycle (McGuire et al., 2009) to boreal forest climate change and disturbance. Euskirchen et al. (2010) provide an overview of biogeochemical feedbacks with climate change in Alaska’s boreal forests and Hobbie et al. (2002) focused on the relationship between nutrient availability and carbon dynamics. An excellent synthesis of the impacts of climate change on Alaskan forests was provided by Wolken et al. (2011) and terrestrial ecosystem feedbacks to the climate system are presented in Field et al. (2007). However, the aforementioned studies and review papers were not focused on supporting land management activities with respect to assessing and managing carbon sources and sinks.

To support our objective we: 1) synthesize results from numerous studies on the carbon cycle with a focus on research from the Alaskan boreal biome, 2) provide estimates of carbon sources in soil and vegetation in interior Alaska, 3) pinpoint potential sources and sinks for carbon on landscapes in interior Alaska, 4) identify expected changes in sources and sinks with climate change and human disturbance including how compounding disturbances can affect the boreal system, and 5) provide recommendations to address some of the impending challenges facing land managers in their carbon itemization efforts.

A growing need exists for carbon source and sink inventories as the first step in identifying and implementing effective climate change mitigation strategies. Executive Order 13514 (Federal Leadership in Environmental, Energy, and Economic Performance) was released in October, 2009. One aspect of this Order requires U.S. Government agencies to account for greenhouse gas (GHG) emissions with the long-term goal of establishing emission-reduction targets. Actions taken by the federal government toward itemizing sources and sinks of GHG will likely dictate how state and local governments, non-profit, and for-profit entities follow. Most itemizations of GHG emissions focus on sources of CO2 and CH4 because they are anthropogenic greenhouse gases and major components of the global carbon cycle. CH4 molecules are 21 times more effective at warming earth’s greenhouse than CO2 (Schlesinger, 1997). As a consequence, smaller CH4 sources (or sinks) can play a comparatively large role in carbon itemization assessments.

Federal entities must address Executive Order 13514 across a wide variety of ecosystems, including the boreal biome which encompasses most of the land between 45° to 83° N and is the most broadly distributed terrestrial biome globally and in Alaska. This biome is characterized by dense forests with high carbon content soils and contains 49% of the global terrestrial forest carbon (Lal, 2005). In Alaska the region is characterized by extreme seasonality between winter and summer temperatures, dramatic seasonal variations in sunlight, a long-lasting snow cover, and a subsurface with heterogeneous bodies of discontinuous permafrost.

Atmospheric CO2 and CH4 can be derived from terrestrial ecosystem processes such as aerobic and anaerobic respiration and human activities, predominantly combustion and agriculture. Major CO2 sinks include photosynthetic uptake, stabilization in soils, and the dissolution of CO2 from soil and rock mineral weathering and calcium carbonate formation in the ocean. There is great uncertainty in identifying and quantifying the feedbacks between enhanced atmospheric CO2 concentrations and whether the response of vegetation and soil microbes provide sources or sinks for this additional atmospheric carbon source (Hartley et al., 2012; Higgins and Harte, 2012; Van Huisstden and Dolman, 2012; Krankina et al., 2012).

Carbon dioxide (CO2) and methane (CH4) emissions into earth’s atmosphere enhance the absorption of radiation emitted from earth’s surface, resulting in warming. Anthropogenic activities contributing to elevated GHG concentrations in the atmosphere are dominated by fossil fuel burning, wildfires, and changing soil or vegetation regimes (Karl and Trenberth, 2003). This is compounded by the removal of GHG sinks when natural ecosystems undergo land use change and the attendant surface albedo and energy balance are altered (Pielke et al., 2007; Bonan, 2008).

3.1 Climatic and physiographic setting

The climate in interior Alaska is continental, with a mean annual air temperature of -3.3°C and typical mean monthly temperatures of 20.2°C in the summer (July) and -31.7°C in the winter (January) with extremes ranging from -51°C to 38°C (Jorgenson et al. 2001). The mean annual precipitation is 28 cm, with maximum annual snow fall of 1.7 m (Jorgenson et al., 2001). Approximately 40–45% of the annual precipitation arrives as snow (Liston and Hiemstra, 2011). In interior Alaska the permanent snow line is ∼1,600 meters elevation in the Alaska Range to the south (Takeuchi, 2009). None of interior Alaska is above the permanent snow line.

The region is underlain by discontinuous permafrost that ranges from a few meters to over 50 m thick and is most commonly found on north facing slopes, in valley bottoms, and under poorly drained soils (Figure 2; Hopkins et al., 1955; Anderson, 1970; Hamilton et al., 1983; Jorgenson et al., 2001; Douglas et al., 2008). In locations with taliks (zones of unfrozen material created as permafrost thaws in areas below the seasonally thawed active layer) the spatial extent of permafrost bodies is extremely difficult to measure or predict. The horizontal and vertical heterogeneity of permafrost distribution in the area prevents the establishment of any simple estimation of permafrost extent.

Figure 2
Extent and type of permafrost for the Tanana Flats and surrounding areas of interior Alaska.

Permafrost classes range from continuous to absent. Data sources for permafrost are varied with the majority from Jorgenson et al. (2008) and from the fine scale maps of the Department of Defense Tanana Flats and Yukon Training areas in Jorgenson et al. (1999). The spatial distribution of permafrost outside of the Tanana Flats and Yukon Training areas is not well mapped and is largely uncertain. Note that the Tanana Flats lowland directly south of Fairbanks and the uplands east of north Pole have been mapped at higher resolution than the rest of the mapped area (Jorgenson et al., 2008). As a consequence, for these areas a far richer dataset is presented.

Figure 2
Extent and type of permafrost for the Tanana Flats and surrounding areas of interior Alaska.

Permafrost classes range from continuous to absent. Data sources for permafrost are varied with the majority from Jorgenson et al. (2008) and from the fine scale maps of the Department of Defense Tanana Flats and Yukon Training areas in Jorgenson et al. (1999). The spatial distribution of permafrost outside of the Tanana Flats and Yukon Training areas is not well mapped and is largely uncertain. Note that the Tanana Flats lowland directly south of Fairbanks and the uplands east of north Pole have been mapped at higher resolution than the rest of the mapped area (Jorgenson et al., 2008). As a consequence, for these areas a far richer dataset is presented.

Permafrost acts as a confining bed in the subsurface that reduces soil water storage capacity and restricts groundwater flow (Hinzman et al., 1991; 1998; Kane et al., 1991; Woo, 2000). The presence of permafrost can greatly affect the stream discharge response to storm activity (Carey and Quinton, 2005), the geochemistry of stream flow (Petrone et al., 2006; Bagard et al., 2011), and seasonal fluxes of nutrients like carbon and nitrogen out of northern watersheds (Carey, 2003; O’Donnell and Jones, 2006; Frey et al., 2007; Frey and McClelland, 2009; Walvoord and Striegl, 2007; Cai et al., 2008a; 2008b; O’Donnell et al., 2012a; Douglas et al., 2013).

In interior Alaska, permafrost distribution correlates with physiography (Figure 3). Lowland landscapes are composed of a complex mosaic of forest, scrub, bog, fen, and open water bodies (Figure 4). They are underlain by sporadic bodies of ice-rich and ice-poor permafrost in gelisols predominately composed of organic material (Pergelic Cryofibrists and Histic Pergelic Cryofibrists) or gravel, silt, and sand (Pergelic Cryochrepts; Brabets et al., 2000). The many open wetland corridors (floating vegetation mats and fens), thought to be permafrost free, are bordered by forests growing above permafrost or perched on top of well-drained gravel substrates. Lowlands comprise 42% of the boreal region in Alaska (574,000 square km), of which 13% is susceptible to collapse-scar bog formation (Jorgenson and Shur, 2007). These bogs form when permafrost acts as an aquiclude (a low flow zone in the subsurface around which flow is channeled) or an aquitard (a zone with no subsurface flow) a few meters below the surface. If a talik penetrates deep enough, the previously closed hydrological system can open up and increase contact with groundwater. In this case fens, wetlands with higher mineral and nutrient content, usually due to subsurface flow, can form. Racine and Walters (1994) suggest the large areas of fen wetlands are fed by groundwater discharge that moves northward from the Alaska Range to the south with permafrost acting as a confining layer channeling subsurface flow. Abundant slumping of trees indicate recent and past permafrost degradation along the upland margins of Tanana Flats.

Figure 3
A cross section through upland and lowland terrains in interior Alaska.

A cross section illustrating the variety of land forms, hydrologic features, and vegetation in interior Alaska and their relationship to permafrost. Adapted from Jorgenson et al., 2010.

Figure 3
A cross section through upland and lowland terrains in interior Alaska.

A cross section illustrating the variety of land forms, hydrologic features, and vegetation in interior Alaska and their relationship to permafrost. Adapted from Jorgenson et al., 2010.

Figure 4
Photographs representing the vegetation typical of lowland ecosystems in interior Alaska.

A) An overview of an area that burned on the Tanana Flats in 2010. Green regions are locations where the vegetation is typical of a collapse-scar bog. The brown and black standing dead trees are characteristic of burned spruce and birch. B) A view showing the fen to forest transition in an unburned location in the fall. C) A close up of the fen from B. D) The vegetation typical of a bog. This area burned in 1988 and the dense forest in the background is typical of a birch forest.

Photos by Thomas A. Douglas

Figure 4
Photographs representing the vegetation typical of lowland ecosystems in interior Alaska.

A) An overview of an area that burned on the Tanana Flats in 2010. Green regions are locations where the vegetation is typical of a collapse-scar bog. The brown and black standing dead trees are characteristic of burned spruce and birch. B) A view showing the fen to forest transition in an unburned location in the fall. C) A close up of the fen from B. D) The vegetation typical of a bog. This area burned in 1988 and the dense forest in the background is typical of a birch forest.

Photos by Thomas A. Douglas

Upland landscapes contain well-drained rocky or loess covered soils with permafrost occurring only on north-facing slopes and valley bottoms (Figure 5; Osterkamp et al., 2000; Jorgenson et al., 2001; Douglas et al., 2008). Soils include Alfic Cryochrepts on warmer, well drained south facing sites, and Histic Pergelic Cryochrepts at colder locations underlain by permafrost (Viereck et al., 1983). Upland permafrost tends to be more ice-poor and generally more thaw stable than lowland permafrost. Uplands have higher elevation gradients, drier soils, and more exposed bedrock than lowlands (Jorgenson et al., 2001). South facing and/or well-drained upland slopes tend to be free of permafrost (Osterkamp et al., 2000). The north slopes of the Alaska Range are composed of uplands containing rocky moraines with thick loess deposits.

Figure 5
Photographs representing the vegetation typical of upland ecosystems in interior Alaska.

A) A winter time view of the Stuart Creek watershed showing the rolling hills and elevation gradients. B) A close up of the hills above Stuart Creek in late fall with mixed stands of birch-aspen forest and spruce (green). C) A valley bottom in the Caribou-Poker Creek Research Watersheds underlain by discontinuous permafrost with a spruce forest. D) A recent photo of a side hill in upland terrain in the Caribou-Poker Creek Research Watersheds that shows standing dead aspen trees (white colored crowns) from a 1999 fire. Standing dead spruce trees are visible in the foreground.

Photos by Thomas A. Douglas

Figure 5
Photographs representing the vegetation typical of upland ecosystems in interior Alaska.

A) A winter time view of the Stuart Creek watershed showing the rolling hills and elevation gradients. B) A close up of the hills above Stuart Creek in late fall with mixed stands of birch-aspen forest and spruce (green). C) A valley bottom in the Caribou-Poker Creek Research Watersheds underlain by discontinuous permafrost with a spruce forest. D) A recent photo of a side hill in upland terrain in the Caribou-Poker Creek Research Watersheds that shows standing dead aspen trees (white colored crowns) from a 1999 fire. Standing dead spruce trees are visible in the foreground.

Photos by Thomas A. Douglas

3.2 Ecosystems of the boreal biome in interior Alaska

Slope, aspect, soil type and disturbance history play major roles in controlling vegetation composition in the Alaskan boreal forest just as they do in other places in the world. Vegetation in the boreal forest of interior Alaska (Chapin et al., 2006; Figure 6 and Tables 1, 2, and 3) is predominantly (40%) black spruce (Picea mariana) on permafrost, lowlands, and north-facing upland slopes, especially after low severity fires (Hollingsworth et. al 2006). Deciduous Alaska paper birch (Betula neoalaskana) and aspen (Populus tremuloides) with a matrix of either pure or mixed white spruce (Picea glauca) stands are common on uplands and south-facing slopes. Sphagnum spp. dominates the groundcover in poorly drained lowland black spruce forests, while feather mosses (Pleurozium schreberi, Hylocomnium splendens) are more common in drier sites. Mosses directly influence permafrost stability in the black spruce ecosystems by insulating the ground against summer heat (Jorgenson et al., 2010; O’Donnell et al., 2009; Turetsky et al., 2012). Moss cover, coupled with thicker organic soils (0.5 to >3m) in black spruce stands are more protected against permafrost thaw than birch forests with no moss cover and little organic soil (Osterkamp et al., 2000). Mature black spruce stands typically have a well-developed carbon-rich surface organic layer dominated by a nearly continuous Sphagnum spp. ground cover in poorly drained locations and feathermoss and lichen on well- and moderately well drained sites (Trumbore and Harden, 1997; Hollingsworth et al., 2008).

Figure 6
Predominant land cover classes for the Tanana Flats and surrounding areas of interior Alaska.

From the Alaska 2001 National Land Cover Database (Homer et al., 2007) including evergreen (34%) and deciduous (12%) forest, shrubland (24%), woody wetland (13%), and barren (7%). Percentages are calculated from Department of Defense-owned lands within the visible domain.

Figure 6
Predominant land cover classes for the Tanana Flats and surrounding areas of interior Alaska.

From the Alaska 2001 National Land Cover Database (Homer et al., 2007) including evergreen (34%) and deciduous (12%) forest, shrubland (24%), woody wetland (13%), and barren (7%). Percentages are calculated from Department of Defense-owned lands within the visible domain.

Table 1.
Land cover types and their areal extent and percent coverage for the interior Alaska domain for three major U.S. Army training ranges (Tanana Flats, Donnelly, and Yukon), and for the U.S. Army Corps of Engineers Chena Lakes project

Physiographically, the Tanana Flats, Donnelly, and Yukon training lands are characterized as lowlands, uplands, and a mosaic of lowland and wetland terrains, respectively.

Table 2.
Lowland ecosystem types with their characteristic vegetation, permafrost regime, total area, and areal change information

1Adapted from Jorgenson et al., 2001 

Table 3.
Upland ecosystem types with their characteristic vegetation and permafrost regime

Alaska paper birch (Betula neoalaskana) and aspen (Populus tremuloides) are canopy dominants on well-drained, often permafrost-free mineral soils common in uplands, particularly on south facing slopes. White spruce ecosystems with moderately thick (20 cm) organic soil layers can also be found in uplands and in lowlands and north-facing upland slopes and are often underlain by permafrost. Typical understory vegetation includes high bush cranberry (Viburnum edule), dwarf dogwood (Cornus canadensis), willows (Salix spp.), lingonberry (Vaccinium vitis-idaea), bear berry (Arctostaphylos uva-ursi), and prickly wild rose (Rosa acicularis). Mixed forests typically have a less well-developed organic soil layer and only contain permafrost on north-facing slopes.

The ∼8,000 km2 Tanana Flats Lowland is an alluvial fan built from silt and gravel extending northward from the northern slopes of the Alaska Range (Cleve et al., 1993; Walters et al., 1998;). It forms a complex matrix of land cover types due to variability in the mineral substrate (silts to gravels), the local hydrology, and the presence of discontinuous permafrost. The patchy forest developed on the lowlands reflects differences in stand age related to time since last fire, the permafrost regime, and local flooding frequency. Collapse-scar bogs are common in the lowland black spruce forests and are dramatic features easily discerned with aerial photography. Deciduous forest stands generally have greater seasonally thawed active layer depths due to an absence of a moss groundcover that renders these locations more vulnerable to summer seasonal thaw. The ice contents of the birch forest vary widely and can reach greater than 50% while ice contents are closer to 20% in the black spruce stands (Osterkamp et al., 2000). Groundwater upwelling, ultimately sourced from the Alaska Range to the south, results in a matrix of nutrient-rich fens where permafrost is absent (Jorgenson et al., 1999). The fens are dominated by sedges and brown mosses (Table 2).

Permafrost depths have been documented up to 47 m thick within the Tanana River floodplain (Chacho et al., 1995), but they generally range from 0.5 to 12 m (Racine and Walters, 1994). A study undertaken in the late 1990s presented information on the extent of permafrost underlying the 263,759 hectares of U.S. Army Fort Wainwright training lands on Tanana Flats south of Fairbanks and between Nenana and Salcha (Figures 1, 2, and 4; Osterkamp et al., 2000). They reported 17% of the Tanana Flats training area lands were unfrozen, 48% had stable permafrost, 31% was partially degraded, and 4% was totally degraded.

3.3 Disturbance and succession in the boreal biome

3.3.1 Interior Alaska fire dynamics

Fire is the greatest ecologically important disturbance in the boreal biome (Viereck, 1973a; Zoltai et al., 1998; Kasischke et al., 2000a; Wurz et al., 2006; Barrett et al., 2011) affecting stand structure and species composition through patterns of mortality and regeneration (Johnstone and Chapin, 2006). Fire exerts significant influence on species composition and carbon cycling (Bond-Lamberty et al., 2007; Turetsky et al., 2011; Bernhardt et al., 2011; Table 4). The degree to which fire influences long-term boreal biome dynamics depends on fire severity, return interval, and burn depth (Viereck, 1973b; Johnstone et al., 2010; Beck et al., 2011). Recent study results suggest fires are growing larger, more intense, and more frequent in interior Alaska as a result of climate warming (Kasischke and Turetsky, 2006; Kasischke et al., 2010). Together, these factors have the potential to alter the long-term trajectory of forest regeneration (Johnstone and Chapin, 2006). The last decade experienced double the number of fires of any decade within the previous 40 years (Kasischke et al., 2010), a trend expected to continue as a result of potentially warmer and longer growing seasons (Duffy et al., 2005; Chapin et al., 2008; Balshi et al., 2009).

Table 4.
Soil carbon stocks for burned and unburned interior Alaska black spruce forest and peatland sites

The horizon thickness indicates the amount of carbon for the total measured depths.

Fire is a natural part of boreal forest ecosystem dynamics and can regenerate food sources for large mammals such as moose and decrease insect outbreaks. Historically, most fires were caused by lightning strikes. Today, humans cause most fires but lightning-caused fires consume a greater areal extent. Based on our analysis of lightning strike density information provided by the Alaska Fire Service (Figure 7), we determined that 40% of the number of fires since 1947 were caused by lightning which has an uneven distribution across the landscape. Human activities, including accidents and prescribed burns, caused 47% of the fires, and the remaining 13% are attributed to a variety of other causes. Lightning strikes caused 57% of the area burned since 1947, while 26% were human caused, and 17% were attributed to other causes.

Figure 7
Lightning strike density and historical fire area information.

Top: lightning strike density (strikes per square km) occurring within a 5 km search radius from the center of each 500 m cell for the Tanana Flats and surrounding areas of interior Alaska. Data from the Alaska Fire Service (1986–2011; http://fire.ak.blm.gov/afs). Bottom: historical fire area data for interior Alaska from the Alaska Fire Service from 1947–2011 distinguished by the cause of each fire.

Figure 7
Lightning strike density and historical fire area information.

Top: lightning strike density (strikes per square km) occurring within a 5 km search radius from the center of each 500 m cell for the Tanana Flats and surrounding areas of interior Alaska. Data from the Alaska Fire Service (1986–2011; http://fire.ak.blm.gov/afs). Bottom: historical fire area data for interior Alaska from the Alaska Fire Service from 1947–2011 distinguished by the cause of each fire.

Over the period from 1947 to 2011, larger fires have occurred predominantly on lowlands, and the period from 2001 to 2011 experienced the largest area burned of any decade (Wendler et al., 2011), which also experienced an overall decrease in the number of fires per year compared to the 1970s and 1990s. These larger, more severe fires are generally associated with greater burn depths, higher carbon emissions, greater destruction of surface soils, enhanced permafrost degradation, and initiation of deciduous forests where the surface organic material burned down to mineral soils (O’Neill et al., 2002; 2003; Zhuang et al., 2003). These processes stimulate the decomposition of carbon that was previously locked frozen in permafrost and this leads to additional CO2 emissions (Goulden et al., 1998).

In interior Alaska the number of lightning derived fires has increased over time (Kasischke and Turetsky, 2006; Turetsky et al., 2011) and lightning strikes are more prevalent in upland terrain versus lowland terrain (Figure 7). As would be expected, lightning strikes predominantly occur in the summer months with June and July accounting for most of the lightning in a given year (Wendler et al., 2011).

Fire has a large direct and immediate impact on carbon cycling as the combustion process causes the release of trace gases such as CO2, CH4, and CO and generates black carbon (Harden et al., 2000). In addition to the direct influences on carbon fluxes through biomass combustion, fires also influence successional patterns, which can have long-term consequences on carbon cycling (Shenoy et al., 2011) and permafrost stability (O’Donnell et al., 2012b). Historically, black spruce forests are burned through stand-replacing fires every 70–130 years with an average return interval for the overall boreal forest of about 29–300 years (Yarie, 1981; Dyrness et al., 1986; Kasischke et al., 2000b). Fire frequency and severity depend largely on the climate (i.e., meteorology) and on human activities such as suppression and ignition (Burn, 1998). Recent increases in burn severity have led to changes in successional trajectories that break the legacy of black spruce regeneration and this has led to a shift toward deciduous forest (Johnstone et al., 2010). These ecological shifts are facilitated by the combustion of the moss understory and organic soil layer that increases the mineral soil seedbeds that favor deciduous recruitment (Chapin et al., 2000; Johnstone and Kasischke, 2005; Kasischke and Johnstone, 2005; Johnstone and Chapin, 2006). Furthermore, unlike moderately-burned stands, where the organic soil remains and spruce trees have higher recruitment success, severely burned stands favor faster-growing deciduous tree species which outcompete spruce trees because their taproots, absent in spruce, help buffer against moisture stress.

Black spruce ecosystems are often underlain by permafrost that is thermally protected by the moss groundcover, particularly where hummock-forming Sphagnum mosses are present (Turetsky et al., 2010, 2011; Nossov et al., 2013). The thick organic layer in permafrost-impacted environments actively protects permafrost from thawing as part of the “ecosystem-protected permafrost” identified by Shur and Jorgenson (2007). The type of moss groundcover exerts an important control on the soil thermal regime since mosses with higher moisture content, such as Sphagnum mosses, protect the soil from warm summer temperatures better than feather mosses, which have a higher thermal conductivity (Yoshikawa and Hinzman, 2003). Since feather mosses are more likely to burn during fires, their presence or absence has a greater thermal impact on the thermal regime post fire. The depth of the active layer is largely dependent on the thermal conductivity of the soil, which is a function of density, moisture content, and thermal phase.

In a study of post-fire soil climate dynamics near Delta Junction, Alaska, Harden et al. (2006) found that the coldest, wettest soils were accompanied by the thickest organic mats. They also report that with every centimeter of organic mat thickness the temperature at 5 cm depth was 0.5°C cooler during the summer months. Model studies suggest that if an organic layer can remain >7–12 cm thick following wildfire, the impact of the fire on permafrost stability will be minimal (Yoshikawa and Hinzman, 2003). In fact, the largest driver of active layer thickness is the thermal conductivity of the organic layer, and as long as the organic thickness is not significantly altered, even with a decrease in surface albedo from the fire, the active layer is not significantly impacted. The importance of the organic layer and surface vegetation was also demonstrated by a study in which bulldozing of vegetation for fire lines showed more impact on active layer thickness than the severely burned black spruce during the 1971 Wickersham Dome fire. In this study the active layer was 161% deeper in the burned compared to the unburned forest (Viereck, 1982).

Fire chronosequence studies have shown that carbon fluxes are most dynamic in the first 20 years since fire with forests changing from a carbon source to a sink within roughly 10 years (e.g., Litvak et al., 2003; Welp et al., 2006; Randerson et al., 2006; Liu and Randerson, 2008; Amiro et al., 2010; Goulden et al., 2011). The largest control on the carbon cycle exerted by fires is the burn severity (Goulden et al., 2011; Iwata et al., 2011; Genet et al., 2013; Jafarov et al., 2013). An eddy covariance study of carbon exchange at a severely burned black spruce forest at the Poker Flat Research Range revealed the forest became a carbon sink within 5 years (Iwata et al., 2011). Despite the lower gross primary productivity, compared to other burned forest studies, respiration was also much lower because the fire consumed the organic soil layer.

In recent years the increased intensity of boreal fires has led to greater consumption of the protective moss layer resulting in enhanced permafrost thaw (Jorgenson et al., 2010). When this leads to ice-wedge degradation and thermokarst (subsidence of the ground surface following thaw of ice-rich permafrost) lakes are formed or relatively ice-rich permafrost thaws. These responses can cause collapse-scar bogs or other wetlands to form. Thawing permafrost compounds carbon cycle changes related to wildfire as ground subsidence changes forests to wetlands and lakes. An assessment of soil drainage in Alaska, as defined by water-holding capacity, hydraulic conductivity, and the position of the seasonal water table, showed a statistical correlation between poorly drained wet areas and historical burn (since 1950) which is likely related to the high C, and thus fuel load, of poorly drained soils (Harden et al., 2001). Punctuated drought or short-term seasonally dry soil conditions of these typically poorly drained soils could therefore promote burning and loss of belowground carbon.

Forest succession following fire, human disturbance, or floodplain development on the Tanana River floodplain initiates with willows (Salix spp.), followed by willow-alder (Alnus tenuifolia), then balsam poplar (Populus balsamifera) and white spruce (Picea glauca), and black spruce (Picea mariana) after the forest floor develops and organic matter begins to accumulate in the soil (Viereck et al., 1993; Hollingsworth et al., 2010; O’Donnell et al., 2012b). The emergence of black versus white spruce on floodplain soils may be due to differences in site drainage, which, in turn, is generally controlled by geomorphology (Mann et al., 1995; Hollingsworth et al., 2010). This succession process is associated with a steady accumulation of carbon in the vegetation and surface soils following flooding disturbance (Van Cleve et al., 1971; 1983; Nossov et al., 2011). Net primary productivity (NPP) decreases as the forest transitions from deciduous to needleleaf forest since nutrients become less available as organic matter builds up in the soil and a change to more recalcitrant species occurs (Berg, 2000). Development of black spruce/feathermoss systems constrains decomposition thermally, and this limits nutrient availability for net primary productivity (Lavoie et al., 2005; Wickland and Neff, 2008). Consequently, CO2 uptake is low under slow-growing black spruce forests, but respiration (and thus CO2 release) also slow as organic carbon accumulates in the soil. Minimal release of CH4 occurs from these forests and in some cases they can act as a carbon sink (Euskirchen et al., 2010). Fire may alter this source-sink dynamic (Bond-Lamberty et al., 2007).

3.3.2 Nutrient dynamics

Changes in nutrient (e.g., N and P) availability can have large consequences on ecosystem structure and function, which impacts carbon cycling. Nutrient availability is low in boreal lowlands, such as bogs and black spruce forests, and species found in these systems are adapted to grow under reduced nutrient conditions. Fertilization has been shown to negatively impact evergreen shrubs growing in both bogs and black spruce forests in favor of grasses (Manninen et al., 2009). Along with understory vegetation, white spruce shows minimal growth response to nutrient addition (Nams et al., 1993). While increased spruce growth has been observed in Yukon Territory forest stands, cone and seed production of the spruce did not increase (Turkington et al., 1998). After disturbance, the amount of ammonia increased in the soil after fertilization and microbial C decreased, suggesting greater microbial C consumption following fertilization (Manninen et al., 2009). In a black spruce forest post-fire N addition resulted in a significant decline in soil C initially but total microbial biomass declined and lower respiration rates increased ecosystem C storage in black spruce forests experiencing frequent fires (Allison et al., 2010).

It is well known that bryophytes are negatively impacted by increased nutrients, especially Sphagnum mosses (Turkington et al., 1998; Güsewell et al., 2002; Juutinen et al., 2010; Turetsky et al., 2010). Bryophytes are more recalcitrant than the herbaceous understory vegetation that replaces them with increased nutrient fertilization. As a consequence, the amount of belowground carbon storage decreases with decreased bryophyte prevalence. A 9-year study of nutrient addition to the Mer Bleue peatland in eastern Canada showed a loss of bryophyte biomass over time and an associated decrease overall ecosystem photosynthesis rates within five years. By the eighth year, net ecosystem exchange was nearly the same in the perturbed and control plots due to increased shrub growth where nutrients were added. However, ecosystem respiration rates increased 24–32%. This suggests high nitrogen deposition lessens the CO2 sink strength of the bog (Juutinen et al., 2010). Ecosystems have also been found to recover quickly following nutrient drawdown (Limpens and Heijmans, 2008). Nutrient additions to certain ecosystems, such as deciduous forests, have the potential to increase net primary productivity and thus strengthen the CO2 sink. In contrast, with moss-dominated bogs and black spruce understory, nutrient addition increases carbon cycling rates and has the potential to minimize the net carbon sink.

3.3.3 Human and biological disturbance

Any removal of the organic soil layer or moss ground cover will increase the ground heat flux, which promotes permafrost thaw. This includes road building, clearing fire lines, developing trails, airboat use in shallow water, and infrastructure development. On the Tanana lowlands, particularly in the wet fens and floating mats, disturbance by airboats has resulted in mortality of the plants on trails and ultimately more open water areas. The open water has a lower albedo than the highly reflective fen vegetation, which promotes further permafrost degradation. Evidence exists for an expansion of Typha latifolia and Menyanthes trifoliata along airboat trails (Racine and Walters, 1994) suggesting airboat traffic not only impacts plant growth but also changes the fen community. The presence of deeper water and more frequent disturbance decreases the ability of the fen to uptake carbon.

Extensive insect outbreaks have been associated with climate warming in Alaska. Most notable are the spruce bark beetle (Ips typographus) outbreaks that have reached epidemic levels and caused widespread spruce mortality on the Kenai Peninsula in south-central Alaska (Berg et al., 2006). There is currently little indication of spruce bark beetles in interior Alaska. Spruce budworm (Choristoneura fumiferana) outbreaks, which affected white spruce populations in interior Alaska in the late 1990s and mid-1980s, were attributed to climate warming and an increase in the rate of larval development (Han et al., 2000; Volney and Fleming, 2000). These insects caused a reduction in tree growth and density which likely reduced forest carbon uptake. Leaf miners (Phyllocnistis populiella Chambers) have extensively infested aspen stands in interior Alaska. They do not result in tree mortality but result in leaf abscission four weeks earlier in heavily mined leaves than healthy leaves, impacting rates of photosynthesis and carbon uptake (Wagner et al., 2008). As agents of disturbance, insects can decrease NPP and this can lead to a reduction in CO2 uptake by forests (Fleming, 2000). A recent review by Hicke et al. (2012) covers the role of insect disturbance on the forest carbon cycle in a variety of locations including the boreal forest of Canada and the United States. They discuss the contradictory responses to disturbance whereby some forest stands exhibit increased primary productivity because surviving vegetation experiences increased growth while stands with repeated growth reductions experienced increased tree mortality and decreased productivity. They conclude that though biotic disturbances can have a major impact on forest C stocks and fluxes there are many uncertainties associated with quantifying the effects of disturbance on C budgets.

Alaska has relatively few invasive species, and most are limited to road margins and other clearings. One exception impacting the interior is Melilotus alba, an aggressive early successional monospecific colonizer, now found along several glacial river flood plains (Wurtz et al., 2010), including the Nenana River (Conn et al., 2011). M. alba has a significant advantage as an early colonizer because of its ability to fix large quantities of nitrogen, up to 100 kg/ha (Sparrow et al., 1993; 1995). At high densities the species can reduce seedling survival, cover, and density of native species (Conn et al., 2011). Another aggressive invasive in interior Alaska is the narrlowleaf hawkweed (Hieracium umbulatum) which spreads aggressively on burned soil and along roadways (Cortés-Burns et al., 2007). If invasive species outcompete forest stands for water or nutrients and cause forest dieoffs and an increased presence of standing dead trees, fire prevalence and severity could increase, particularly in the immediate years following tree die-off.

Lodgepole pine (Pinus contorta var. latifolia), though not an invasive species, has been shown to have strong ecosystem effects where it has expanded into southern boreal forests, particularly as a dominant species following fire (Johnstone and Chapin, 2003) due to its high growth rates (Gutsell and Johnson, 2002). Interior Alaska has no species listed as threatened or endangered by the U.S. Fish and Wildlife Service. The American peregrine falcon (Falco peregrinus anatum), common in interior Alaska, was delisted in 1999.

3.4 Permafrost degradation in uplands and lowlands

Permafrost is common in lowland areas, covering roughly 80 percent of the lowland landscape (Figure 2). The presence (or absence) of permafrost is intricately linked to the local vegetation, soils composition, and surface hydrology. The extent of this permafrost is expected to decline in coming decades as a response to climate warming (Jorgenson et al., 2001). Thermokarst development increased 21% between 1949 and 1998 on Tanana Flats and the expected continuation of climate warming in the region could lead to permafrost elimination by 2100 (Jorgenson et al., 2001; Euskirchen et al., 2006; 2009; Lawrence et al., 2008). Due to its complex mosaic of soils, vegetation, permafrost extent and surface water bodies, the Tanana Flats lowland would be expected to respond dramatically and largely unpredictably to the degradation (and eventual total loss) of permafrost. Of particular concern is the potential loss or change of wetland habitats in the low gradient lowland locations if subsurface permafrost aquicludes or aquitards are lost.

Permafrost degradation leads to significant changes in peatland ecosystem carbon cycling (Camill, 1999; Camill et al., 2001; Turetsky et al., 2000; 2008; 2010; O’Donnell et al., 2012b; Wu, 2012). Ground subsidence following thaw typically results in surface inundation, an increase in hydrophilic taxa, and enhanced anaerobic decomposition, resulting in increased CH4 flux and decreased CO2 flux to the atmosphere (Bellisario et al., 1999; Turetsky et al., 2002; Wickland et al., 2006; Lee et al., 2012). In areas exposed to wind, ground subsidence can also foster increased snow depths and warmer winter soil temperatures (Zhang, 2005a). Following thermokarst initiation, the thermal capacity of the water then has the capability to continue to thaw the permafrost laterally (Camill, 2005; O’Donnell et al., 2012a). In locations where permafrost degradation leads to enhanced drainage and surface drying, greater oxidation reduces carbon accumulation rates (Robinson and Moore, 2000) and enhanced CO2 release (Waelbroeck, 1993; Frolking et al., 2006).

In upland locations where permafrost is present, partial or initial degradation of permafrost can result in wetter soil conditions where inundation occurs or where drainage is limited by permafrost or low hydraulic soil conductivity. These wetter conditions may increase hydrophilic vegetation such as Sphagnum and sedges. However, as permafrost degradation continues, surface microtopographic changes could yield greater surface run-off and drying from higher topographic regions (Schuur et al., 2007). This drying kills mosses and tussocks (Schuur et al., 2007; Osterkamp et al., 2009) and enhances net CO2 flux to the atmosphere.

3.5 Interior Alaska wetlands and hydrogeology

The boreal region contains a varied distribution of lakes that are stable, increasing, or decreasing. Lake stability is the result of heterogeneous permafrost, hydraulic gradients, and lake and catchment topography (Roach et al., 2011).Though they only cover ∼2% of interior Alaska by area, lakes are an important component of the area’s boreal ecosystems and their carbon cycle processes because of aquatic-terrestrial links between water bodies and the surrounding soil and vegetation. Nutrients, especially carbon and nitrogen, impact terrestrial ecosystems nearby, leading to enhanced lake productivity with increased run-off or permafrost thaw (Symstad et al., 2003; Ball et al., 2010).

The response of permafrost lakes to climate warming or other disturbance is not simple or uniform across the landscape (Roach et al., 2011; Rover et al., 2011). Time series remote sensing studies of the areal extent of thousands of lakes in arctic and sub-arctic Alaska have shown a general trend whereby lake area extents in the discontinuous permafrost zone tend to be decreasing, in some areas up to 31 percent of lake area coverage, while lake areal extents in the continuous permafrost zone are stable (Riordan et al., 2006). Lakes perched above continuous permafrost are believed to be more stable because they experience less vertical or horizontal drainage (Smith et al., 2005).

To determine the primary cause of lake drawdown in boreal Alaskan lakes, Roach et al. (2011) studied four lake regions and determined deeper lakes that formed from ice-rich thermokarst tended to be more persistent than shallow lakes (<1m) either absent of permafrost or perched on ice-poor permafrost. Terrestrialization was found to be the primary mechanism for lake drawdown (as opposed to subsurface talik drainage). In shallow, smaller volume basins, lake terrestrialization is thought to result from a combination of warmer water and greater proportional nutrient input that speeds vegetative productivity. Under a warmer climate this will increase transpiration rates (Roach et al., 2011), resulting in greater carbon uptake and vegetation growth (Keeling et al., 1996; Myneni et al., 1997).

In the Yukon Flats National Wildlife Refuge in Alaska, where stable isotope analyses revealed that ∼95% of the studied lakes sourced water from precipitation, river water, and groundwater, lake lowering is predominantly the result of evaporative losses exceeding supply as a result of warming temperatures and no net change in precipitation (Anderson et al., 2013).

Another driver of shrinking lakes in discontinuous permafrost is the development of thermokarst which leads to talik development and subsurface drainage (Marsh and Neumann, 2001; Yoshikawa and Hinzman, 2003). Fire or human activities that remove surface vegetation can cause ice-rich permafrost to thaw differentially and amplify irregular surface topography (Hinzman et al., 2003b; Yoshikawa and Hinzman, 2003). Where disturbance leads to talik formation sub-surface lake drainage, surface drying, and the lateral movement of surface water can occur.

Fen, bog, and marsh systems, often linked hydrologically to lakes and ponds, also play major roles in nutrient cycling between aquatic and terrestrial ecosystems (Fan et al., 2013; Rober et al., 2014). Numerous studies at the Bonanza Creek Long-Term Ecological Research in interior Alaska have confirmed these results (e.g., Kane et al., 2010; Wyatt et al., 2010; Jones et al, 2013). In the discontinuous permafrost zone permafrost degradation may result in the partial or full disappearance of the permafrost aquiclude which can increase or decrease local hydraulic gradients depending on a variety of factors (Britton, 1957; Kane and Slaughter, 1973; Billings and Peterson, 1980; Woo, 1986; Jorgenson et al., 2001; Yoshikawa and Hinzman, 2003). Fen systems are particularly sensitive to water table fluctuations with studies showing rapid (i.e., on the order of days) increased CH4 fluxes from fens when their water table elevations decrease (Roulet et al., 1992; Windsor et al., 1992). Flooded fens in interior Alaska were found to provide an increased CO2 sink, whereas droughts reduced their carbon sink capacity and even turned them into small carbon sources (Chivers et al., 2009).

Watersheds underlain by permafrost exhibit an intense seasonality of flow paths, discharge, and streamwater biogeochemistry (O’Donnell and Jones, 2006; Petrone et al., 2006; Walvoord and Striegl, 2007; Barker et al., 2014). Recent studies in the Chena River watershed, a non-glaciated river draining the boundary between the Tanana Flats lowlands and nearby uplands, have shown a strong correlation between discharge and dissolved organic carbon (DOC) and total dissolved nitrogen (Cai et al., 2008a, 2008b; Douglas et al., 2013). Of particular note, nutrient exports are rapidly flushed from surface soils during large discharge periods such as snow melt runoff or major summer precipitation events. This is consistent with results from studies in other subarctic and arctic rivers (Striegl et al., 2005, Finlay et al., 2006; Raymond et al., 2007; Walvoord and Striegl, 2007; McNamara et al., 2008; Frey and McClelland, 2009; Guo et al., 2012). The DOC mobilized during high flow events is generally believed to be of modern age (Guo and Macdonald, 2006; Neff et al., 2006) and part of the DOC transported during spring melt may be labile (Holmes et al., 2008). Where changes in permafrost extent are sufficient to alter flowpaths, the carbon balance of subarctic watersheds could shift from a net sink to a net source (McGuire et al., 2009; Grosse et al., 2011). Since snow melt occurs during the major seasonal transition between winter and spring/summer any shift in the timing of snow melt would alter the timing of the major nutrient fluxes out of northern watersheds. However, the potential effects of this change in seasonality are unknown.

4.1 Terrestrial changes as a response to a warming climate

High-latitude amplification of climate warming has caused Alaska’s boreal biome to warm 1.3°C during the last 50 years (Shulski and Wendler, 2007), twice the rate of the global high latitude average (Arctic Climate Impact Assessment 2005, Hinzman et al., 2005; Solomon et al., 2007; Chapin et al., 2010). Future climate scenarios project a 3–7°C increase in mean annual air temperatures for the Alaskan interior over the next 90 years (Chapman and Walsh, 2007; Walsh et al., 2008). In addition to this increase in mean annual temperatures, the occurrence of extreme low temperatures (characterized by temperatures <-40°C) has decreased on average from 14 to 8 days annually (Wendler and Shulski, 2009). A non-statistically significant 11% decrease in annual precipitation has occurred over the last 90 years near Fairbanks (Wendler and Shulski, 2009) with the strongest decreases in precipitation occurring in spring followed by winter.

Shoulder-season and seasonal transitions are occurring. More of the annual precipitation is arriving as rain (Liston and Hiemstra, 2011), and snow blanketed the boreal landscape 17 days fewer in 2009 compared to 1979. Fall snow arrival happens two days per decade later. The spring snow cover from 1972–2008 (Brown et al. 2010) and 1979–2009 (Liston and Hiemstra, 2011) has been disappearing 5 days earlier per decade. This reduction in snow cover primarily comes at a critical time when insolation is high as summer approaches (Chapin et al., 2005; Derksen and Brown 2012), resulting in a dramatically higher energy influx and a longer growing/drought season. The lack of a substantial increase in precipitation will do little to offset higher summer evapotranspiration rates (Scenarios Network for Alaska Planning. 2010), which may result in drier soils, lower lake levels, and increased fire potential.

We present historical data and projected future air temperature and precipitation information during summer and winter in interior Alaska in Figures 8 to 15. Temperature outputs (Figures 8, 10, 12, and 14) are taken from the Scenarios Network for Alaska and Arctic Planning (SNAP; http://www.snap.uaf.edu) program’s selected model results derived from climate model simulations from five climate models that perform well in Alaska, together with their multi-model average. These five models are: the Canadian Centre for Climate Modelling and Analysis General Circulation Model version 3.1, the Max Planck Institute European Centre Hamburg Model 5, the US National Oceanic and Atmospheric Administration Geophysical Fluid Dynamics Laboratory Coupled Climate Model 2.1, the United Kingdom Meteorological Office Coupled Model 3.0, and the Japan Center for Climate System Research Model for Interdisciplinary Research on Climate-medium resolution. The B1, A1B, and A2 scenarios (Intergovernmental Panel on Climate Change, 2000) represent, respectively, relatively lower, intermediate, and higher CO2 emission futures (Solomon et al., 2007) created for the World Climate Research Program’s Coupled Model Intercomparison Project, Phase 3 (CMIP3). Outputs from the five GCMs selected for SNAP and used here were statistically downscaled for the SNAP program from the GCM native scales to the sub-continental scales of Parameter-elevation Regressions on Independent Slopes Model (PRISM; http://prism.oregonstate.edu) data in Alaska (2km or 0.771km, depending on year) using a delta method with spline fitting on the anomalies in monthly temperature and precipitation as projected and compared to observations (see http://www.snap.uaf.edu/faq.php#faq_2). Precipitation data (Figures 9, 11, 13, and 15) are from the SNAP historical dataset using PRISM and data from the Climate Research Unit (CRU; http://www.cru.uea.ac.uk) datasets where PRISM data are unavailable.

Figure 8
Decadal summer (June-August) mean temperatures for Tanana Flats and surrounding areas of interior Alaska.

Major road systems (Figure 1) are black.

Figure 8
Decadal summer (June-August) mean temperatures for Tanana Flats and surrounding areas of interior Alaska.

Major road systems (Figure 1) are black.

Figure 9
Downscaled projected decadal summer (June-August) mean temperatures for interior Alaska.

Major road systems (Figure 1) are black. Data are from the Scenarios Network for Alaska and Arctic Planning projected dataset which is derived from downscaled climate model simulations averaged from five climate models that perform well in Alaska (General Circulation Model version 3.1, European Centre Hamburg Model 5, Coupled Climate Model 2.1, Coupled Model 3.0, and Model for Interdisciplinary Research on Climate-medium resolution). B1, A1B, and A2 scenarios represent relatively low, intermediate, and high CO2 emission futures (Solomon et al., 2007).

Figure 9
Downscaled projected decadal summer (June-August) mean temperatures for interior Alaska.

Major road systems (Figure 1) are black. Data are from the Scenarios Network for Alaska and Arctic Planning projected dataset which is derived from downscaled climate model simulations averaged from five climate models that perform well in Alaska (General Circulation Model version 3.1, European Centre Hamburg Model 5, Coupled Climate Model 2.1, Coupled Model 3.0, and Model for Interdisciplinary Research on Climate-medium resolution). B1, A1B, and A2 scenarios represent relatively low, intermediate, and high CO2 emission futures (Solomon et al., 2007).

Figure 10
Decadal summer (June-August) total precipitation for interior Alaska.

Major road systems (Figure 1) are black. Data are from the Scenarios Network for Alaska and Arctic Planning historical dataset derived from downscaled climate data from the Parameter-elevation Regressions on Independent Slopes Model (PRISM; http://prism.oregonstate.edu) and Climate Research Unit (CRU; http://www.cru.uea.ac.uk) datasets.

Figure 10
Decadal summer (June-August) total precipitation for interior Alaska.

Major road systems (Figure 1) are black. Data are from the Scenarios Network for Alaska and Arctic Planning historical dataset derived from downscaled climate data from the Parameter-elevation Regressions on Independent Slopes Model (PRISM; http://prism.oregonstate.edu) and Climate Research Unit (CRU; http://www.cru.uea.ac.uk) datasets.

Figure 11
Downscaled projected decadal summer (June-August) average total precipitation for interior Alaska.

a. Major road systems (Figure 1) are black.

Figure 11
Downscaled projected decadal summer (June-August) average total precipitation for interior Alaska.

a. Major road systems (Figure 1) are black.

Figure 12
Decadal winter (December-February) mean temperatures for interior Alaska.

Major road systems (Figure 1) are black.

Figure 12
Decadal winter (December-February) mean temperatures for interior Alaska.

Major road systems (Figure 1) are black.

Figure 13
Downscaled projected decadal winter (December-February) mean temperatures for interior Alaska.

Major road systems (Figure 1) are black.

Figure 13
Downscaled projected decadal winter (December-February) mean temperatures for interior Alaska.

Major road systems (Figure 1) are black.

Figure 14
Decadal winter (December-February) total precipitation for the Tanana Flats and surrounding areas of interior Alaska.

Major road systems (Figure 1) are black.

Figure 14
Decadal winter (December-February) total precipitation for the Tanana Flats and surrounding areas of interior Alaska.

Major road systems (Figure 1) are black.

Figure 15
Downscaled projected decadal winter (December-February) average total precipitation for the Tanana Flats and surrounding areas of interior Alaska.

Major road systems (Figure 1) are black.

Figure 15
Downscaled projected decadal winter (December-February) average total precipitation for the Tanana Flats and surrounding areas of interior Alaska.

Major road systems (Figure 1) are black.

Mean summer temperatures (June-August; Figure 8) show decadal variability and dramatic warming since 1990. Lowland regions have warmed more than the surrounding hills or mountains of the Alaska Range. Though some climate scenarios project greater rates of warming than others for interior Alaska (Figure 9) the trend of overall increasing air temperatures is projected to increase in the future.

There was a large amount of interdecadal variability in summer precipitation (1906–2006) (Figure 10) but no significant trends in interior Alaska (Wendler and Shulski, 2009). Modeling scenarios for future summer precipitation in the area (Figure 11) result in decadal variability in precipitation with the IPCC B1 and A1B scenarios projecting slightly wetter conditions and the IPCC A2 scenario projecting an overall drying. In general, the climate of interior Alaska is dry, receiving 29 cm of annual precipitation (Jorgenson et al., 2001; Wendler and Shulski, 2009). As a consequence, a few large convective summer storms can have a substantial impact on any season’s precipitation total.

Figure 12 depicts decadal mean winter temperatures for interior Alaska from 1910 to 2009. The results show a general trend of increasing winter temperatures since 1910. Wendler and Shulski (2009) report the greatest winter season warming in the months of December and January for their 1906 to 2006 records from a station near Fairbanks. The mean December-February warming since 2006 is 1.3°C (Wendler and Shulski, 2009). Future climate scenarios project an increase in December to February temperatures between the 2010 to 2019 and 2030–2039 decades with consistent agreement across three IPCC model scenarios (Figure 13). Winter season precipitation has decreased from 1910 to 2009 (Figure 14). The future winter season is expected to be shorter based on modeling efforts that simulate snow accumulation and melt over 30 years (1979–2009; Liston and Hiemstra, 2011) For interior Alaska, simulated results indicate a 4 cm decrease in precipitation and a 0.5°C drop in temperature during the snow-covered seasons. Forty percent of the annual precipitation arrived as snow in 2009 compared with forty-five percent in 1979. Over that 30 year period, snow arrived 8 days later and melted 9 days earlier, leading to a shortened snow season of ∼17 days. Projections of future precipitation in the area based on the IPCC model scenarios (Figure 15) suggest winter precipitation totals are expected to be slightly lower while summer precipitation is not expected to change much (Figure 10). This regional and local scale of change is in contrast to the fact that GCMs typically project warmer temperatures lead to greater moisture in the atmosphere and, as a consequence, increased precipitation rates (Räisänen, 2008; Walsh et al., 2008).

Increasing temperatures are expected to have pronounced effects on soil and water biogeochemical processes in regions underlain by discontinuous permafrost (Osterkamp and Romanovsky, 1999). As the climate warms in both summer and winter, permafrost will continue its current warming trend (Romanovsky et al., 2012), the active layer will become thicker, the lower boundary of permafrost will become shallower and the areal extent of permafrost will decrease. These structural changes will affect components of the surface water and energy balances. As the active layer thickens there is greater storage capacity for soil moisture and lags are introduced into the hydrologic response times to precipitation events. When permafrost is close to the surface, the stream and river discharge peaks are higher and base flow is lower (i.e., streams are “flashier”). If permafrost degrades, connectivity between surface and subsurface water flowpaths can increase (Walvoord et al., 2012). Depending on the hydraulic gradient, more infiltration of surface water or exfiltration of groundwater occurs when permafrost extent decreases. This has significant impacts at both large and small scales by reducing summer runoff and increasing the yearly proportion of winter runoff as deeper flowpaths become a larger component of subsurface flows (Douglas et al., 2013).

It is estimated that northern circumpolar permafrost soils contain 1,672 billion metric tons of organic carbon (Tarnocai et al., 2009) which accumulated slowly over thousands and tens of thousands of years. Roughly 60% of this carbon is believed to be located in the circumpolar permafrost zone and an estimated 277 billion metric tons of the northern soil carbon pool is associated with peatlands (Schuur et al., 2008). Discontinuous permafrost in interior Alaska is composed partially of “yedoma-type” permafrost which has a carbon content of 2–5%, up to 30 times greater than what is typically present in thawed mineral soils (Zimov et al., 2006a, 2006b; Douglas et al., 2011; Kanevskiy et al., 2011). This yedoma permafrost formed during the Pleistocene when windblown loess was repeatedly deposited to the soil surface to create ice-rich syngenetic permafrost (Shur and Jorgenson, 2007; Douglas et al., 2011). Yedoma deposits are believed to contain almost one quarter of the northern permafrost soil carbon pool (Tarnocai et al., 2009). Permafrost in interior Alaska has also been formed from alluvial and aeolian deposition (syngenetic type), peat accumulation (syngenetic and quasi-syngenetic types) and climate variations (epigenetic type).

Northern high latitude terrestrial soils have acted as carbon sinks over millennia as plants soak up CO2 from the atmosphere, partially decompose, and are buried in soils. However, a recent increase in high severity fires, an increase in permafrost thaw, and projected changes to the soil and vegetation composition due to climate warming will likely lessen this sequestering capability (Hayes et al., 2011). A larger concern is that permafrost thawing will yield increased carbon emissions from arctic and subarctic landscapes as permafrost degrades, providing a positive feedback to climate warming (Frey and Smith, 2005; McGuire et al., 2009; Koven et al., 2011; Schaefer et al., 2011; Schuur et al., 2013). Further, carbon emitted from thawing permafrost is several thousand years old, which essentially introduces a “new” carbon source to the atmosphere.

To identify the current inventory of carbon sinks in interior Alaska boreal ecosystems we separated the estimates available for carbon pools into aboveground, belowground, and soil C (Table 5). We also converted CH4 and CO2 fluxes reported in the literature into units of mg/m2/day to calculate fluxes from a variety of ecosystem types in interior Alaska for comparisons across studies (Table 6). To compare carbon fluxes between CO2 and CH4 the molar mass of each molecule was taken into account. Most of the CH4 fluxes are close to 0 given the large relative percent error for calculations based on flux measurements (Table 6). Bogs, moderately thawed soils, collapse-scar bogs/floating mats, and black spruce forest are all net carbon stores (see Table 6), but they can shift to sources on interannual timescales (Euskirchen et al., 2014). Alteration from one state to another with permafrost thaw (i.e., black spruce forest to collapse-scar bog) can lead to a short-term (years to centuries) net source of carbon in the form of mostly methane. However, these areas can switch to a net sink on the order of many centuries to millennia (O’Donnell et al., 2012b). The stability of these ecosystems is tied to preservation of permafrost and/or protection from disturbance. If permafrost thaws, bogs can turn to fens or they can dry as subsurface water drains away from the wetland. When permafrost under black spruce forest degrades or burns the capacity of the CO2 sink will likely decrease, but will eventually (over decades) increase as peat production increases following thaw and forest regeneration occurs following burn.

Table 5.
Carbon stocks in aboveground and belowground biomass for representative interior Alaska ecosystem types

Cells are left blank if data are not available for that particular parameter.

Table 6.
A summary of results from studies that present methane and carbon dioxide fluxes from common ecosystems present in interior Alaska

The carbon equivalents are given for the CH4 and CO2 fluxes. CO2 values represent net ecosystem exchange (NEE). Negative values indicate carbon uptake and positive values indicate flux to the atmosphere. Cells are left blank if data are not available for that particular parameter.

1NEE: net ecosystem exchange

Annual aboveground biomass production is lowest in black spruce ecosystems with less than one quarter of the annual biomass production of its deciduous counterparts and one third of the biomass production of white spruce ecosystems (Van Cleve and Viereck, 1983; Yarie and Billings, 2002). Consequently, CO2 uptake is low under slow-growing black spruce forests, but respiration (and thus CO2 release) also slows as organic carbon accumulates in the soil. Minimal release of CH4 occurs from these forests and in some cases they can act as a carbon sink (Euskirchen et al., 2010). Fire may alter this source-sink dynamic (Bond-Lamberty et al., 2007).

Despite the greater aboveground productivity of the deciduous tree species soil carbon storage is greater in black spruce forests. This is attributed to poor soil drainage which slows decomposition, inhibits sub-surface moisture run-off, accumulates more organic matter, and promotes the upward migration of the active layer during steady state climate conditions. Permafrost protects soil carbon from further decomposition and functions as a soil C stabilization mechanism. Carbon stocks in the top meter of black spruce lowland soils average 513 mg/ha, 85% more than what is found in deciduous forest soils (Table 5) and 75% more than white spruce soils. The amount of carbon in black spruce lowland soils is only rivaled by that of collapse-scar bogs, which average 751.7 mg C/ha (Jones et al., 2013). This suggests permafrost degradation, particularly in lowland terrains, can increase belowground carbon storage in the long term (centennial to millennial timescales) due to more anaerobic soil conditions that slow decomposition (O’Donnell et al., 2012b).

The amount of carbon uptake and biomass production in different ecosystem types depends on the complex interplay between successional stage, water balance, nutrient availability, and topography. For example, in black spruce forests where CO2 fluxes were measured, hot and dry summer conditions resulted in half the CO2 uptake of the previous growing season (Ueyama et al., 2006). Similar conditions in a black spruce forest in interior Alaska resulted in a net flux of CO2 to the atmosphere (Euskirchen et al., 2014). In a recent study of boreal black spruce soils, Wickland and Neff (2008) showed temperature and moisture are the main controls on carbon mineralization rates. Further, they found carbon mineralization rates were highest in soils from a well-drained site absent of permafrost. The extensive groundcover of mosses in the boreal forest contributes significantly to overall productivity, between an estimated 25% and 14% total aboveground net primary productivity in permafrost and permafrost-free upland forests (Turetsky et al., 2010). Flux estimates for moss dominated areas range from 24 to 77 grams carbon/m2/yr (Oechel and Van Cleve, 1986; Bisbee et al., 2001). This moss accumulation effectively buffers soils from atmospheric climate perturbations because of its high porosity, low thermal conductivity, and high water holding capacity (Rydin and MacDonald, 1985; O’Donnell et al., 2009; Turetsky et al., 2010; Turetsky et al., 2012). Topography can play a major role in carbon dynamics due to its role in controlling lateral and vertical movement (and, thus, availability) of water (Grant, 2004; Kane et al., 2005; 2010).

Thermokarst lakes are significant CH4 “hot spot” sources relative to other landform types with the highest CH4 fluxes occurring along actively thawing lake margins where highly labile carbon previously frozen in permafrost undergoes anaerobic decomposition and releases CH4 through diffusive or ebullition (bubbling) processes (Walter et al., 2007a; 2007b; 2008; Shakhova et al., 2010). Much of the carbon released from thermokarst lakes is old, having been locked in the permafrost since the Holocene or Pleistocene (Walter et al., 2006). Lakes can also emit a substantial amount of CO2 (Kling et al., 1991; Cole et al., 1994; Algesten et al., 2004). Although lakes are large CH4 and CO2 sources, they also have the potential to sequester carbon as lake sediments and peat accumulate with time (Jones et al., 2012; Walter-Anthony et al., 2014).

Evidence from Siberia suggests lake drainage can inhibit further lake expansion which may also limit carbon fluxes to the atmosphere from these systems (Van Huissteden and Dolman, 2012). Lake drainage can also result in peat accumulation within a matter of decades, and the accumulation of peat in these basins act as a carbon sink (Jones et al., 2012) and has the potential to offset losses of CH4 and CO2, but only after centuries to millenia. A recent data and modeling analysis of the radiative forcing of thermokarst lakes since the last deglaciation has revealed that despite being CH4 sources, thermokarst lakes as a whole became carbon sinks ∼5,000 years ago (Walter-Anthony et al., 2014). Thermokarst lakes are ephemeral on the landscape, lasting only 2,000–3,000 years before they drain or partially drain, which in both cases promotes peat accumulation (Jones et al., 2012; Walter-Anthony et al., 2014). This is because in the deep study lakes in yedoma permafrost, the water at the bottom is cold enough to limit decomposition of organic matter. The authors warn, however, that under a warmer climate that cannot support permafrost, these lakes will disappear where no aquitard is otherwise present. Then the lake sediments and soil carbon will decompose and be emitted to the atmosphere. In regions where lakes are expected to persist even in the absence of permafrost, warmer temperatures will result in lake terrestrialization, which will increase carbon uptake (Roach et al., 2011).

5.1 Potential changes in carbon sources and sinks with permafrost thaw

The amount of carbon released as a result of thawing permafrost depends partly on the quality and quantity of the organic substrate in the previously frozen soil (Schuur et al., 2008). Temperature, moisture content, nutrient availability, and oxygen availability are the primary controls on carbon transformation rates in soils. The primary control on the release of CH4 or CO2 to the atmosphere is the degree of soil saturation where anaerobic conditions result in CH4 production and dry aerobic conditions result in CO2 production.

In uplands, permafrost thaw can lead to the creation of topographic lows that can result in the formation of ponds, lakes or other standing water bodies. In well-drained uplands, permafrost thaw can lead to surface water drainage or soil drying (Schuur et al., 2008). Wetter conditions in uplands would favor carbon sequestration by Sphagnum mosses and sedges (Trumbore and Harden, 1997; Turetsky et al., 2000; Camill et al., 2001). Permafrost thaw in drier uplands would increase afforestation, lower albedo, and increase decomposition and the release of soil C (Goulden et al., 1998; Stieglitz et al., 2000; Lloyd et al., 2003; as cited in Camill, 2005).

In lowlands, surface inundation of organic-rich material leads to enhanced CH4 production. Complete permafrost degradation leads to drastic changes in surface moisture conditions as permafrost thaw triggers ground subsidence and surface inundation, which, in turn, results in a shift to completely anaerobic soil conditions. Depending on the degree of ice richness of the permafrost this ground subsidence can result in the formation of collapse-scar bogs, fens or thermokarst lakes. Talik formation can lead to significant groundwater input and the minerotrophic conditions that support fen vegetation. Methane emissions from collapse-scar bogs and fens increase as both a function of surface inundation from permafrost thaw and increased NPP, which is related to both temperature and nutrient availability (Bubier et al., 1995; 2005). In a metadata analysis of methane emissions from global wetlands, Turetsky et al. (2014) found that sites with permafrost emitted the least amount of methane, and that fens on average emit more than bogs. Plant species composition is a strong indicator of methane flux, with a greater presence of graminoid taxa linked to higher emissions (Bubier, 1995; Couwenberg et al., 2011).

Quantifying CH4 release from collapse-scar bogs and fen systems is difficult because of their unevenly distributed, episodic ebullition and diffusive fluxes. Nonetheless, CH4 emission from collapse-scar bogs and fens formed after permafrost thaw is typically an order of magnitude greater than pre-thaw (Table 6). Generally, they are only sources of carbon to the atmosphere for a few years to decades following thaw, but on longer decadal to centennial timescales they remain net carbon sinks. Their sink capacity can be weaker depending on interannual climate patterns (i.e., weaker during warm, dry conditions; O’Donnell et al., 2012b; Euskirchen et al., 2014). The modern age of the carbon released from these systems suggests CH4 production, for the most part, is not tied to the consumption of old, previously frozen carbon. More recently, however, emissions from thawing permafrost, thus older methane, accounted for up to 22% of flux from methane released during senescence in a collapse-scar bog in interior Alaska (Klapstein et al., 2014). Similarly, a permafrost peatland incubation study showed a major source of respired carbon (in the form of CO2) came from just above the active layer, which is an older carbon source (Dorrepaal et al., 2006). The permafrost thaw was not complete and only resulted in active layer thickening. Active layer thickening without complete collapse of deeper permafrost is thought to be a likely scenario for much of the discontinuous permafrost zone and is consistent with observations (Zhang et al., 2005a; 2005b). However, collapse-scar bogs usually expand laterally and represent a catastrophic collapse of the ground surface (Jorgenson and Osterkamp, 2005; O’Donnell et al., 2011). For both uplands and thermokarst lakes, permafrost thaw has resulted in elevated releases of CH4 (uplands) and CO2 (lakes) to the atmosphere (Walter et al., 2006; Zimov et al., 2006a; 2006b; Schuur et al., 2009). In both ecosystems a significant fraction of the carbon released to the atmosphere was found to be old permafrost carbon with radiocarbon ages ranging from the Holocene to the late Pleistocene. Thus, both collapse-scar bogs and thermokarst lakes initially pulse large quantities of carbon to the atmosphere, but over longer millennial timescales become carbon sinks as peat accumulates (O’Donnell et al., 2012b; Jones et al., 2013; Walter-Anthony et al., 2014).

In a well-drained upland tundra ecosystem in central Alaska, minimally thawed areas showed the lowest release of old carbon while deeper thaw resulted in a higher proportion of old carbon loss. Areas thawed over the last 15 years were net carbon sinks because increased plant growth offset carbon losses (Schuur et al., 2009). The release of old carbon was 78% higher in locations that had thawed decades earlier compared to sites where thaw had been minimal. This suggests long-term and continuous carbon loss from these upland areas as permafrost continues to thaw (Schuur et al., 2009). The greater carbon loss with increasing time since thaw in this upland site may reflect redistribution of water as the site thaws or, perhaps, more extensive vertical thawing (i.e., a deeper active layer). Ground subsidence due to melting of ground ice features may result in increased surface wetness initially but surface drying occurs as adjacent areas continue to thaw, followed by gully formation, which allows water to drain away. This process ultimately leads to the death of peat-forming mosses, an increase in shrub growth (Osterkamp et al., 2009), and the initiation of forest succession. Over time, the landscape is initially a carbon source (drying and thawing), then a slow sink as aboveground biomass accumulates (forest development).

5.2 Observed carbon fluxes from terrestrial ecosystems as a result of permafrost thaw

Arctic and sub-arctic watersheds export much of their carbon as DOC flushed out of surface vegetation and soils during spring melt (Macdonald and Yu, 2006; Raymond et al., 2007; Cai et al., 2008a). Roughly 60% of the annual DOC export from arctic and sub-arctic rivers occurs within the two months following snow melt and ice breakup (Finlay et al., 2006; Raymond et al., 2007), and this carbon is of modern age (Guo and Macdonald, 2006; Neff et al., 2006). Spring melt is also associated with what is typically the largest surface water loads of other nutrients (like nitrogen) and trace metals (Rember and Trefry, 2004) as melt water interacts with surface soil and vegetation (McNamara et al. 2008; Frey and McClelland, 2009). The timing of spring melt has been moving earlier (Stone et al., 2002) but the potential ramifications of an earlier melt for nutrient export fluxes or composition are largely unknown.

Changes in permafrost extent will likely have profound impacts on the export of DOC, dissolved inorganic carbon (DIC), and particulate organic carbon (POC) from terrestrial ecosystems into nearby waterways (Guo et al., 2007). The Arctic Ocean comprises only 1% of the global ocean volume but receives over 10% of the global river transported terrestrial dissolved organic matter (DOM) because arctic rivers contain extremely high DOM concentrations (Dittmar and Kattner, 2003). Striking relationships between DOC concentrations in permafrost versus permafrost-free watersheds are observed in western Siberia, where higher DOC concentrations occur in permafrost-free watersheds with large peatlands (Frey and Smith, 2005). If the watersheds become permafrost-free by 2100, DOC export in the West Siberia region could increase 46% (Frey et al., 2007). However, positive relationships between higher DOC concentrations and greater permafrost area have been reported elsewhere and this is ascribed to shallow flow paths of water through organic-rich soils where permafrost is present (Petrone et al., 2006). Meanwhile, other studies have shown decreased DOM export with increasing permafrost degradation after a brief increase in export initially following thaw in Alaska (MacLean et al., 1999; Striegl et al., 2005; 2007; Petrone et al., 2006; 2007), the Yukon Territory (Carey, 2003), and central Siberia (Kawahigashi et al., 2004; Prokushkin et al., 2007). This is thought to occur as a result of increased adsorption of DOM in newly exposed underlying mineral soils (Frey and McClelland, 2009).

As the climate warms, thawing of permafrost in peatland-rich regions could lead to the release of old carbon to streams and rivers (Smith et al., 2005) because subsurface flow would increase (Frey and McClelland, 2009). This would lead to an increase in the export of carbon from permafrost terrains. The hydraulic conductivity of peat can vary widely (Holden and Burt, 2003) but is generally considered low (Reynolds et al., 1992). Low hydraulic conductivities increase the residence time of water in peatlands and increase the potential for decomposition and leaching of DOM, but this also slows the rate of export from peatlands to nearby waterways. DOM production is low in slowly draining, anoxic peat soils because the anoxic conditions reduce decomposition rates (Moore and Knowles, 1989; Stutter et al., 2007). Another factor contributing to DOC production is the recalcitrance of the organic material in the peat, because vascular fen vegetation is more labile than Sphagnum moss.

Studies in West Siberian peatlands report that permafrost limits dissolved organic matter exports but DOC values rise in watersheds as peatland cover increases (Frey and Smith, 2005; Frey and McClelland, 2009). Striegl et al. (2007) report potential decreasing exports of dissolved organic matter from watersheds undergoing permafrost thaw due to the adsorption of DOM by newly exposed mineral soils. Thus, depending on permafrost extent and the percentage of mineral versus peat soil encountered by flows through watersheds, DOC export could either increase or decrease with permafrost thaw.

The lability of DOC appears to vary depending on time of the year. DOC lability is substantially higher during the spring freshet (Cooper et al. 2005; Guo and Macdonald, 2006; Holmes et al., 2008) and is largely recalcitrant during low-flow summer periods (Dittmar and Kattner, 2003; Holmes et al., 2008; Frey and McClelland, 2009). This can be explained by the short residence time, colder temperatures, and low microbial activity during spring melt. In contrast, increased thaw in the summer results in less surface water interaction with sediments and increased incorporation of slower flowing water at depth. Greater export of decaying terrestrial organic matter into the Arctic Ocean has been suggested to increase pCO2 levels, resulting in outgassing (Shakhova and Semiletov, 2007). This was observed along the Laptev and East Siberian Seas by Anderson et al. (2009) who determined an excess of DIC equal to 1012 g C is expected to outgas to the atmosphere due to terrestrial organic matter decay.

Although permafrost thaw results in the release of labile organic carbon, the age of DOC exported from arctic streams has been found to be of recent origin (Benner et al., 2004; Neff et al., 2006). Studies from the Kolyma River basin (Neff et al., 2006), the Yenisey, Lena, Ob’ in Siberia, and the Mackenzie and Yukon Rivers in Canada and Alaska (Raymond et al., 2007) show relatively young 14C ages of DOC during the spring freshet but DOC ages increase in the late summer as thaw depth increases. This suggests as climate warms and thaw depths increase the potential to export older carbon to the Arctic Ocean will increase.

5.3 Carbon sources and sinks in interior Alaska lands - a synthesis

We have synthesized the results from numerous studies on ecosystem dynamics and the carbon cycle in the boreal biome into three schematic diagrams to illustrate carbon sources, sinks and fluxes in interior Alaska boreal ecosystems (Figures 1618). The purpose of these diagrams is to provide information, based on the most current scientific studies, that can be used as a predictor of how lowland and upland ecosystem vegetation and soil processes are expected to respond to climate warming and their subsequent carbon cycle impacts.

Figure 16
Schematic diagram of present and likely future carbon cycle changes in lowland ecosystems with permafrost.

Climate Change impacts (as outlined in the text) are expected to include warmer summers, warmer winters, and a longer growing season. Natural system disturbances include permafrost thaw, fire, and vegetation changes while human caused disturbances include fire, clearing, impact cratering, vehicle/boat traffic, soil erosion, or the removal of the surface organic layer. References are as follows: a) Euskirchen et al. (2010) and Zhuang et al. (2003); b and c) Bellisario et al. (1998) and Wickland et al. (2006); d) Turetsky et al. (2008); e) Algesten et al. (2004), Walter et al. (2006), and Walter et al. (2008).

Figure 16
Schematic diagram of present and likely future carbon cycle changes in lowland ecosystems with permafrost.

Climate Change impacts (as outlined in the text) are expected to include warmer summers, warmer winters, and a longer growing season. Natural system disturbances include permafrost thaw, fire, and vegetation changes while human caused disturbances include fire, clearing, impact cratering, vehicle/boat traffic, soil erosion, or the removal of the surface organic layer. References are as follows: a) Euskirchen et al. (2010) and Zhuang et al. (2003); b and c) Bellisario et al. (1998) and Wickland et al. (2006); d) Turetsky et al. (2008); e) Algesten et al. (2004), Walter et al. (2006), and Walter et al. (2008).

Figure 17
Schematic diagram of present and likely future carbon cycle changes in lowland ecosystems without permafrost.

Climate Change impacts (as outlined in the text) are expected to include warmer summers, warmer winters, and a longer growing season. Natural system disturbances include fire and vegetation changes while human caused disturbances include fire, clearing, impact cratering, vehicle/boat traffic, soil erosion, or the removal of the surface organic layer. Details on vegetation types can be found in Jorgenson et al. (1999).

Figure 17
Schematic diagram of present and likely future carbon cycle changes in lowland ecosystems without permafrost.

Climate Change impacts (as outlined in the text) are expected to include warmer summers, warmer winters, and a longer growing season. Natural system disturbances include fire and vegetation changes while human caused disturbances include fire, clearing, impact cratering, vehicle/boat traffic, soil erosion, or the removal of the surface organic layer. Details on vegetation types can be found in Jorgenson et al. (1999).

Figure 18
Schematic diagram of present and likely future carbon cycle changes in upland ecosystems with and without permafrost.

Climate Change impacts (as outlined in the text) are expected to include warmer summers, warmer winters, and a longer growing season. Natural system disturbances include permafrost thaw, fire, and vegetation changes while human caused disturbances include fire, clearing, impact cratering, vehicle/boat traffic, soil erosion, or the removal of the surface organic layer. Fluxes and ecosystem dynamics are based on the work of Chapin et al., 2010; Euskirchen et al., 2010; and Wolken et al., 2011.

Figure 18
Schematic diagram of present and likely future carbon cycle changes in upland ecosystems with and without permafrost.

Climate Change impacts (as outlined in the text) are expected to include warmer summers, warmer winters, and a longer growing season. Natural system disturbances include permafrost thaw, fire, and vegetation changes while human caused disturbances include fire, clearing, impact cratering, vehicle/boat traffic, soil erosion, or the removal of the surface organic layer. Fluxes and ecosystem dynamics are based on the work of Chapin et al., 2010; Euskirchen et al., 2010; and Wolken et al., 2011.

Our synthesis assumes climate change in interior Alaska will lead to shorter, warmer winters and warmer summers with increased evapotranspiration and longer growing seasons. We first provide the initial response of the different ecotypes to these changes in the climatic regime. In the diagrams the long thin arrows denote a physical process that disturbs the system like forest degradation or aggradation and landscape flooding or drying. These include natural disturbances like permafrost degradation or the impact of fire on soil thermal properties and human- caused disturbances like vegetation clearing (Nicholas and Hinkel, 1996), prescribed burns, or altering landscape hydrology through infrastructure development. In each of the three diagrams the boxes denote ecosystem processes where we provide the best estimate for magnitude and directionality of CO2 or CH4 fluxes, which are represented by the short, bold arrows.

Mixed forests (Figure 16) slowly accumulate an organic layer as the forest develops toward a needle leaf forest (Euskirchen et al., 2010), resulting in a modest sink for both CO2 and CH4 as organic matter is stored belowground. Depending on the ice content of the permafrost, thaw can result in ponding that will turn the forest from a CO2 sink to a source and to a large CH4 source. Pond or lake drying in the lowlands can result in lake terrestrialization and eventually bog formation, which results in a drawdown of CO2 in the atmosphere through peat accumulation and provides a small CH4 source through anaerobic decomposition. These bogs can form through permafrost degradation or drying of lakes or ponds.

Needle leaf forests can undergo a shift to mixed forest when severe fire burns the surface organic material down to the mineral soil (Johnstone and Kasischke 2005; Barrett et al., 2011), or when fire return intervals are more rapid than the time required for succession to conifers. Severe fires are associated with major CO2 emissions and a major disruption in permafrost thermal stability. Permafrost degradation that causes surface wetting after major fires can transform a needle leaf forest into a bog which leads to CO2 storage and enhanced CH4 emission. If permafrost degradation increases groundwater input into a wetland, a fen can form, and fens are typically larger CH4 sources than bogs (Turetsky et al., 2014).

In lowland ecosystems without permafrost (Figure 17), like some areas in Tanana Flats, with a well-drained substrate, forests are usually deciduous or transitioning to mixed type in areas of more advanced stages of succession. Conversely, a mixed forest will emerge if a fen dries out and a drying lake system can turn into a fen, leading to enhanced CO2 uptake and CH4 release. Fire disturbance will result in a short-term (i.e., several decades) pulse of CO2 and CH4 to the atmosphere, but regeneration of the deciduous forest that emerges post-fire is a CO2 sink. Flooding inundates forests, leading to tree mortality, which can result in fen formation, a large source of CH4 and a moderate CO2 sink.

Upland ecosystems (Figure 18) are characterized by sloping terrain that tends to prevent the emergence of lakes, fens or bogs except in valley bottoms, which often comprise localized lowland type ecosystems. The boreal forest provides a small but steady CO2 sink whether or not permafrost is present and if there is no disturbance by fire, invasive species, or human activities. Immediately following disturbance, the uplands become a CO2 source, but with increasing time since disturbance the forests in upland regions become a substantial CO2 sink. This is the case for both non-permafrost and permafrost settings when the permafrost remains thaw stable. If the permafrost thaws, CO2 and CH4 can be emitted from the soils but if/when the forest continues to grow the area becomes a CO2 sink.

5.4 Challenges to interior Alaska land management due to climate warming

The hydrological and ecological shifts associated with thawing permafrost, particularly on the Tanana Flats lowlands, are expected to have large consequences for state and federal government land management and carbon itemization activities. Unfortunately, change will not come uniformly across the landscape. As a consequence, land and facilities planning will require a better knowledge of locations where change is expected to occur. As discussed previously, there are three major processes that can rapidly alter ecosystems: fire, altered nutrient dynamics due to either permafrost thaw or altered hydrogeology, and human and biological disturbance. These three processes are expected to respond to climate warming and in some cases will amplify the effects of climate warming at a variety of scales.

Increasingly, compounding disturbances are providing a challenge for landscape resource managers. An example: a significant fire occurs in a forest previously affected by physiological stress due to a decades-long drought from warmer summer air temperatures or decreased precipitation. The result of these compounding disturbances is a more severe fire than would be predicted or expected if the fire consumed a healthy forest. If drought conditions persisted during post fire succession then the recovery would take longer and soil development and forest succession would be limited. Another example of compounding disturbances would be that more severe fires, occurring more often, consume highly combustible drought-stressed white spruce and this would lead to yet larger and more expansive fires.

5.5 Anticipated impacts of a warming climate on interior Alaska ecosystems

Based on our assessment of the major carbon cycle processes that govern carbon sources and sinks in the boreal biome of interior Alaska, we have identified three main ecological responses to climate warming that will have the most profound influences on lands in interior Alaska over the next 100 years. These ecological changes will have a pronounced effect on the carbon cycle at a range of spatial and temporal scales.

First, climate-driven permafrost degradation will radically reorganize upland and lowland hydrology and vegetation by altering soil flow paths and changing subsurface flow through the partial degradation or wholesale loss of permafrost (Hinzman et al., 1991; 1998; Kane et al., 1991; Woo, 2000; Douglas et al., 2013). Changes in surface water extent due to permafrost degradation are linked to local ground subsidence either when near-surface ice-rich permafrost thaws (the area may still be underlain by continuous permafrost) or when permafrost is degraded and a connection develops between surface and sub-permafrost water. In the latter case, the local hydraulic gradients determine whether or not the area becomes drained or flooded. As a consequence, the permafrost extent also largely controls vegetation.

The ecosystem and its hydrologic and thermal regimes are dynamically coupled such that neither can be fully understood without considering the other (Harvey, 1988; Chahine, 1992; Hinzman et al., 1996; 2003b). Regions of discontinuous permafrost in interior Alaska and Canada have shown both increases and decreases in surface water spatial extent linked to permafrost degradation (Jorgenson et al., 2001; Romanovsky and Osterkamp, 2001; Osterkamp, 2005; Osterkamp and Jorgenson, 2006; Osterkamp, 2007). Near-surface soil moisture exerts a strong influence on the amount of heat transferred into soils, especially if the ground surface is covered by moss. Changes in the soil thermal regime will significantly alter the lateral and vertical distribution of permafrost. Groundwater recharge, runoff and water storage will be altered considerably as a result of permafrost thaw and this will increase the fraction of subsurface flow in annual river runoff. These changing hydrologic regimes will alter the seasonality and biogeochemistry of Alaskan river discharge by increasing the winter portion and, in some locations, may increase total discharge (Peterson et al., 2002).

Second, a longer growing season (Høye et al., 2007) will favor some species over others and this will alter vegetation composition, abundance, and productivity (Wolken et al., 2011), soils (Grosse et al., 2011) and their associated controls on the boreal carbon cycle. Due to the heterogeneous nature of soils, vegetation, and permafrost extent in interior Alaska and the potential feedbacks between a longer growing season and ecosystem respiration it is difficult to determine at a regional scale whether the longer growing season will result in a greater CO2 or CH4 sink (Euskirchen et al., 2006; Parmentier et al., 2011).

Third, decreased soil moisture and humidity will increase the number, areal extent, and severity of fires, which will alter permafrost stability and the fate of carbon. Fire is the single largest element of rapid change in boreal ecosystems. Carbon stored in aboveground and belowground biomass can be rapidly combusted and emitted to the atmosphere. Numerous studies have shown that the most severe fires burn vegetation down to the mineral soil and this can alter the post-fire ecological trajectory and permafrost thermal stability. Of equal importance is whether the fire-return interval is short enough to limit forest succession and soil development processes from sequestering carbon.

Forecasting ecological responses to climate warming is complicated by soil type, precipitation, surface and ground water hydrology, vegetation, slope, aspect, fire prevalence, and the thermal state of permafrost. Therefore, to reduce uncertainty in future projections and refine planning on lands in interior Alaska the ecosystem hydrologic and thermal regimes need to be linked. This will be accomplished by combining climate modeling with soil and snow pack thermal measurements and modeling with high-resolution digital elevation models and remote sensing measurements. From this, spatially explicit, high-resolution landscape change predictions can be developed. If/where they can be linked to ecosystem drivers such as hydrologic, soil, or vegetation processes a geospatial predictor of carbon source and sink processes can be developed. Synthesizing these measurements and model projections with remotely sensed tools will allow for broader application to a greater variety of terrains. This is of particular utility because much of the land areas in interior Alaska are roadless and remote. Anticipated ecosystem changes in interior Alaska will likely have severe ramifications for how and where State and Federal government agencies should direct resources to promote land use sustainability or ecological preservation or restoration or to address invasive plants.

Based on the information presented herein and guided by the results from numerous studies focused on the boreal biome, permafrost dynamics, climate warming impacts, and the carbon cycle, we provide the following recommendations for land management activities and future research needs. The purpose of these recommendations is to identify a series of measurements, management tools, and planning activities that we believe can help to integrate ecosystem processes into estimates of carbon sources and sinks in ecosystems of interior Alaska. With this information carbon management activities can be prioritized in terms of work effort and cost.

  1. We anticipate the following changes will occur to the ecosystems of interior Alaska with climate warming:

    • Flows in interior Alaska rivers are likely to change as the growing season expands, more precipitation falls as rain, and the timing of spring melt and major precipitation events are altered. Changing seasonality will alter the two major hydrologic transitions: spring melt and fall freeze-up. Since spring melt is a short period of time when large fluxes of DOC and other organic material are exported from watersheds, a shift to an earlier spring melt runoff may affect ecosystem processes. It is not expected that spring peak flows will increase substantially, but warmer summers and enhanced convective activity could yield localized intense storms and flooding. The U.S. Army Corps of Engineers’ Chena River Flood Control project and recreation area were designed with permafrost in mind so the facilities and infrastructure are not likely to be substantially affected by climate warming impacts to permafrost in the near future.

    • Based on projections for the Yukon River (Walvoord and Striegl, 2007), of which the Tanana River is the largest tributary, we would expect to see an increase in the groundwater contribution to streamflow. Summer season flows in the Tanana River are expected to become greater as run-off from increased glacier melt feeds headwater streams (Woo et al., 2008). This could lead to greater flooding risk for the city of Fairbanks and for other infrastructure in the Tanana River floodplain. The effect of an altered groundwater regime on watershed carbon dynamics is not well understood. An increase in low carbon glacial flows would be expected to lead to a decrease in glacial river carbon concentrations but may not change the total fluxes from terrestrial ecosystems.

    • The Tanana Flats (lowland) ecosystem is susceptible to major surface hydrological, vegetation and permafrost changes with climate warming. Due to the intricate feedbacks between permafrost, soils, hydrology, and climate change many of these processes and regime changes are difficult to predict and thus are not easily integrated into future planning scenarios. One consequence could be a continued decrease in the areal extent of permafrost as climate warms and fires become larger and deeper, resulting in greater localized groundwater upwelling zones. Many of the lowland deciduous forests could disappear and turn into a matrix of bogs and fens, depending on localized hydrogeology. The thermal state of permafrost and fire severity will play major roles in controlling whether this ecosystem is a carbon source or sink.

    • Locations like the southern Tanana flats lowland, where lands contain a mixture of upland and lowland ecosystems, will likely exhibit complicated and potentially unpredictable ecosystem responses to climate warming. Rocky moraines may be affected by increased fire frequency and this could lead to the reduction of white spruce forests while loess deposits are vulnerable to collapse due to the thawing of ice-rich permafrost (Toniolo et al., 2009). With warming, permafrost degradation, and the likely drying of the Tanana Flats fire will play a major role in controlling the carbon balance and vegetation succession of these ecosystems. This could seriously impact the infrastructure (roads, buildings, railway, and bridges) planned in this area by the Army, the Air Force, the Alaska Railroad, and the State of Alaska.

    • Upland ecosystems are expected to incur increased fire frequency (disturbance) and a general reduction in black and white spruce forests, drying of south-facing slopes, and a loss of permafrost on north-facing slopes as a response to climate warming (Woo, 2000). This will likely lead to enhanced near term CO2 emissions from this ecosystem that will be reduced over time with post-fire vegetation succession and transition to deciduous forest.

    • Spring melt runoff is the time of the year when more than half of the annual DOC exports leave arctic and sub-arctic watersheds, predominately as modern age carbon and as carbon associated with particulates, surface organic matter, other nutrients (like nitrogen), and trace metals. The spring melt event is expected to be shifted earlier in the spring but the nature and amount of this large pulse flow of aquatic carbon is not expected to change. Importantly, once snow cover departs the system has a longer drying period that both stresses trees and enhances chances of fire (Wolken et al. 2011).

  2. Estimated recovery times following disturbance (as required in Environmental Impact Statements, Environmental Assessments, and future planning scenarios) are likely to become less predictable in the future. Climate-driven changes in the timing of seasonal transitions, the amount of warming by season, and the fraction of wet versus dry precipitation will alter hydrologic, soil thermal, and vegetation regimes, making prediction of recovery difficult. Locations where a specific forest or soil type is currently climatically viable will be altered and this will limit our ability to predict the carbon cycle response(s) to disturbance. This will challenge the flexibility and adaptability of government agency land use and infrastructure development planning. When multiple disturbances are acting on a system at the same time (compounding disturbance) predicting the response becomes even more difficult.

  3. Fire management activities should be focused toward encouraging low to medium severity fires to maintain carbon storage in upper soils, preserve permafrost thermal stability, and prevent the loss of deep soil carbon during and immediately following larger, more severe fires. Though this is a difficult, expensive, and resource-intensive management action to undertake, it could have the greatest impact on preserving carbon in surface soils, promoting vegetation growth, and maintaining permafrost stability. In addition, by more actively managing fires, the potential for loss of infrastructure and lives will be greatly minimized because the potential for prescribed burns or natural fires to grow out of control will be minimized. A series of GIS-based decision support tools could be created and used to identify which locations are most susceptible to high severity fires and what management actions to take. This is explained in greater detail in item #8 below.

  4. Where possible, minimize loss of the insulative and organic rich surface soil and vegetation layer during disturbance (fire, road building, vertical or horizontal infrastructure development). This will help prevent permafrost degradation and will reduce the loss of soil carbon stores. In some select locations it may lead to permafrost formation or at least increased thaw stability (Shur and Jorgenson, 2007).

  5. Where activities such as airboat travel, road clearing, infrastructure development, or prescribed burns occur, we encourage minimizing the rapid draining or horizontal movement of water across the surface and shallow subsurface. Channeling, draining, or ponding water destroys the thermal balance of the surface and subsurface and this can lead to thermokarst and/or the liberation of surface and subsurface carbon. This is of particular concern in permafrost lowland areas where regional gradients are low (i.e., sub-meters of elevation change over kilometers of distance). At some of these sites permafrost provides a subsurface aquitard or aquiclude that maintains surface water. If/where this surface water leaves channels or encounters higher hydraulic conductivity soils the wetland areas supported by horizontal flows may be drained.

  6. We recommend establishing a series of long-term monitoring sites representing a varied amount of time since disturbance and diverse types of disturbances to follow landscape succession and soil processes over time. These sites would be located on lands overseen by the different landusers/stakeholders in the interior Alaska boreal forest. The frequency and types of measurements made would support the user’s specific needs with respect to land management requirements, ecosystem processes, and carbon itemization efforts. The most likely candidate locations for this work would be areas of similar landscape and ecosystem type that burned at different times/decades. This would allow a comparison across time since disturbance. Repeat imagery analysis and remote sensing tools could help identify and measure change over time at these long-term sites. Remote sensing could also be applied toward extrapolating measurements at one landscape type to similar regions in the interior Alaska boreal forest and elsewhere. There are a variety of remote sensing measurements that could be used to help manage fire and human disturbance and to predict where permafrost degradation could have the largest impacts on soil thermal and hydrologic stability and carbon cycle processes. The difficulty in accessing remote field sites and the wide variety of landscape and ecological processes occurring across interior Alaska make it likely that the most useful approach could be remote sensing and repeat imagery analyses verified with focused ground-truth efforts.

  7. More studies are needed during the two major seasonal transitions (winter-spring melt-summer and summer-fall-winter freeze up). The spring melt period is associated with what is typically the largest yearly redistribution of water (snow melt) across the landscape, and this occurs in a period of days to weeks. The fall to winter transition is a key part of soil and vegetation dynamics in the boreal biome as ground freezing is initiated and snow starts to cover the landscape. It is expected that the timing of these seasonal transitions will change in the future and it is likely that the soil, vegetation, and carbon cycle will respond to the changing seasonality and the longer growing season.

  8. We encourage the development of GIS-based decision support tools to help State and Federal government land managers decide where to construct or modify existing facilities and to identify what level of winter and summer vehicle traffic different terrains/landscapes can support as those landscapes change in response to climate warming. Soil characterization information, permafrost extent, wetlands delineation, and climate projection information could be used to identify locations where human activities or climate impacts are at most risk of altering permafrost thaw stability. High resolution digital elevation measurements (such as from ground-based or airborne light distance and ranging (LiDAR)) can be combined with wetland and soil moisture information to predict locations where thermokarst or subsidence is most likely to alter surface hydrology. This would be of particular utility in lowland regions where ground surface elevations may only change by a meter or two over many kilometers of distance. Fire is the single largest ecological disturbance in the boreal biome and provides the most rapid means of altering ecosystem source and sink compartments for carbon. A multi-faceted geospatial decision support tool could also provide guidance on where to apply focused management actions such as prescribed burns and where to site infrastructure or invasive species management strategies.

© 2014 Douglas et al. This is an open-access article distributed under the terms of the Creative Commons Attribution License, which permits unrestricted use, distribution, and reproduction in any medium, provided the original author and source are credited.

Much of the information and ideas presented in this paper were borne from discussions with a wide variety of colleagues representing numerous disciplines and we are thankful that such a collaborative community is working on these topics. Many of the topics and content presented were strengthened by hearty discussions with Torre Jorgenson. An earlier version of this manuscript benefitted greatly from the detailed and thoughtful reviews of Teresa Hollingsworth and Jon O’Donnell. We would also like to thank the anonymous reviewers who provided detailed constructive suggestions on how to improve the manuscript. GIS content was provided by collaborators at the U.S. Army Alaska Directorate of Public Works and Colorado State University.

Algesten
G
,
Sobek
S
,
Bergström
A
,
Ågren
A
,
Tranvik
LJ
, et al
2004
.
Role of lakes for organic carbon cycling in the boreal zone
.
Glob Change Biol
10
(
1
):
141
147
.
Allison
SD
,
Gartner
TB
,
Mack
MC
,
McGuire
K
,
Treseder
K
.
2010
.
Nitrogen alters carbon dynamics during early succession in boreal forest
.
Soil Biol Biochem
42
(
7
):
1157
1164
.
Allison
SD
,
Treseder
KK
.
2011
.
Climate change feedbacks to microbial decomposition in boreal soils
.
Fungal Ecol
4
(
6
):
362
374
.
Amiro
B
,
Barr
A
,
Barr
J
,
Black
T
,
Bracho
R
, et al
2010
.
Ecosystem carbon dioxide fluxes after disturbance in forests of North America
.
J Geophys Res Biogeosci
115
(
G4
).
Anderson
GS
.
1970
.
Hydrological reconnaissance of the Tanana Basin, central Alaska
.
U.S. Geological Survey
Hydrological Investigations Atlas, HA-319
.
Scale: 1:1,000,000
.
Anderson
L
,
Birks
J
,
Rover
J
,
Guldager
N
.
2013
.
Controls on recent Alaskan lake changes identified from water isotopes and remote sensing
.
Geophy Res Lett
40
:
3414
3418
.
Anderson
LG
,
Jutterström
S
,
Hjalmarsson
S
,
Wåhlström
I
,
Semiletov
I
.
2009
.
Out-gassing of CO2 from Siberian Shelf seas by terrestrial organic matter decomposition
.
Geophys Res Lett
36
(
20
).
Arctic Climate Impact Assessment (ACIA)
.
2005
.
Arctic Climate Impact Assessment
,
1042 pp.
,
Cambridge Univ. Press
,
Cambridge, U. K
.
Bagard
M
,
Chabaux
F
,
Pokrovsky
OS
,
Viers
J
,
Prokushkin
AS
, et al
2011
.
Seasonal variability of element fluxes in two Central Siberian rivers draining high latitude permafrost dominated areas
.
Geochim Cosmochim Ac
75
(
12
):
3335
3357
.
Ball
BA
,
Kominoski
JS
,
Adams
HE
,
Jones
SE
,
Kane
ES
, et al
2010
.
Direct and terrestrial vegetation-mediated effects of environmental change on aquatic ecosystem processes
.
BioScience
60
(
8
):
590
601
.
Balshi
M
,
McGuire
A
,
Duffy
P
,
Flannigan
M
,
Kicklighter
D
, et al
2009
.
Vulnerability of carbon storage in North American boreal forests to wildfires during the 21st century
.
Glob Change Biol
15
(
6
):
1491
1510
.
Barker
AJ
,
Douglas
TA
,
Jacobson
AD
,
McClelland
JW
,
Ilgen
AG
, et al
2014
.
Late season mobilization of trace metals in two small Alaskan Arctic watersheds as a proxy for landscape scale permafrost active layer dynamics
.
Chem Geol
381
:
180
193
.
Barrett
K
,
McGuire
A
,
Hoy
E
,
Kasischke
E
.
2011
.
Potential shifts in dominant forest cover in interior Alaska driven by variations in fire severity
.
Ecol Appl
21
(
7
):
2380
2396
.
Beck
PS
,
Goetz
SJ
,
Mack
MC
,
Alexander
HD
,
Jin
Y
, et al
2011
.
The impacts and implications of an intensifying fire regime on Alaskan boreal forest composition and albedo
.
Glob Change Biol
17
(
9
):
2853
2866
.
Bellisario
L
,
Bubier
J
,
Moore
T
,
Chanton
J
.
1999
.
Controls on CH4 emissions from a northern peatland
.
Global Biogeochem Cy
13
(
1
):
81
91
.
Bellisario
LM
,
Moore
TR
,
Bubier
J
.
1998
.
Net ecosystem CO∼ 2 exchange in a boreal peatland, northern Manitoba
.
Ecoscience-Quebec
(
5
):
534
541
.
Benner
R
,
Benitez-Nelson
B
,
Kaiser
K
,
Amon
RM
.
2004
.
Export of young terrigenous dissolved organic carbon from rivers to the Arctic Ocean
.
Geophys Res Lett
31
(
5
).
Berg
B
.
2000
.
Litter decomposition and organic matter turnover in northern forest soils
.
Forest Ecol Manag
133
(
1
):
13
22
.
Berg
EE
,
David
Henry J
,
Fastie
CL
,
De Volder
AD
,
Matsuoka
SM
.
2006
.
Spruce beetle outbreaks on the Kenai Peninsula, Alaska, and Kluane National Park and Reserve, Yukon Territory: relationship to summer temperatures and regional differences in disturbance regimes
.
Forest Ecol Manag
227
(
3
):
219
232
.
Bernhardt
EL
,
Hollingsworth
TN
,
Chapin
III FS
.
2011
.
Fire severity mediates climate-driven shifts in understorey community composition of black spruce stands of interior Alaska
.
J Veg Sci
22
(
1
):
32
44
.
Billings
W
,
Peterson
K
.
1980
.
Vegetational change and ice-wedge polygons through the thaw-lake cycle in Arctic Alaska
.
Arctic Alpine Res
:
413
432
.
Bisbee
KE
,
Gower
ST
,
Norman
JM
,
Nordheim
EV
.
2001
.
Environmental controls on ground cover species composition and productivity in a boreal black spruce forest
.
Oecologia
129
(
2
):
261
270
.
Bonan
GB
.
2008
.
Forests and climate change: forcings, feedbacks, and the climate benefits of forests
.
Science
320
(
5882
):
1444
1449
.
Bond-Lamberty
B
,
Peckham
SD
,
Ahl
DE
,
Gower
ST
.
2007
.
Fire as the dominant driver of central Canadian boreal forest carbon balance
.
Nature
450
(
7166
):
89
92
.
Brabets
TP
,
Wang
B
,
Meade
RH
.
2000
.
Environmental and hydrologic overview of the Yukon River Basin, Alaska and Canada
.
U.S. Department of the interior
,
US Geological Survey Report 99–4204
.
Britton
ME
.
1957
.
Vegetation of the arctic tundra
.
Oregon State University Press
.
Brown
R
,
Derksen
C
,
Wang
L
.
2010
.
A multi-data set analysis of variability and change in Arctic spring snow cover extent, 1967–2008
.
J Geophys Res-Atmos
115
:
D16111
. doi:
Bubier
J
,
Moore
T
,
Savage
K
,
Crill
P
.
2005
.
A comparison of methane flux in a boreal landscape between a dry and a wet year
.
Global Biogeochem Cy
19
(
1
).
Bubier
JL
,
Moore
TR
,
Bellisario
L
,
Comer
NT
,
Crill
PM
.
1995
.
Ecological controls on methane emissions from a northern peatland complex in the zone of discontinuous permafrost, Manitoba, Canada
.
Global Biogeochem Cy
9
(
4
):
455
470
.
Burn
C
.
1998
.
The response (1958–1997) of permafrost and near-surface ground temperatures to forest fire, Takhini River valley, southern Yukon Territory
.
Can J Earth Sci
35
(
2
):
184
199
.
Cai
Y
,
Guo
L
,
Douglas
TA
,
Whitledge
TE
.
2008b
.
Seasonal variations in nutrient concentrations and speciation in the Chena River, Alaska
.
J Geophys Res
113
(
G3
).
Cai
Y
,
Guo
L
,
Douglas
TA
.
2008a
.
Temporal variations in organic carbon species and fluxes from the Chena River, Alaska
.
Limnol Oceanogr
53
(
4
):
1408
.
Camill
P
,
Lynch
JA
,
Clark
JS
,
Adams
JB
,
Jordan
B
.
2001
.
Changes in biomass, aboveground net primary production, and peat accumulation following permafrost thaw in the boreal peatlands of Manitoba, Canada
.
Ecosystems
4
(
5
):
461
478
.
Camill
P
.
1999
.
Patterns of boreal permafrost peatland vegetation across environmental gradients sensitive to climate warming
.
Can J Bot
77
(
5
):
721
733
.
Camill
P
.
2005
.
Permafrost thaw accelerates in boreal peatlands during late-20th century climate warming
.
Climatic Change
68
(
1–2
):
135
152
.
Carey
S
,
Quinton
W
.
2005
.
Evaluating runoff generation during summer using hydrometric, stable isotope and hydrochemical methods in a discontinuous permafrost alpine catchment
.
Hydrol Process
19
(
1
):
95
114
.
Carey
SK
.
2003
.
Dissolved organic carbon fluxes in a discontinuous permafrost subarctic alpine catchment
.
Permafrost Periglac
14
(
2
):
161
171
.
Chacho
E
,
Arcone
S
,
Delaney
A
.
1995
.
Blair Lakes target facility permafrost and groundwater study
.
US Army Cold Regions Research and Engineering Laboratory
,
Hanover, NH
.
30 p
.
Chahine
MT
.
1992
.
The hydrological cycle and its influence on climate
.
Nature
359
(
6394
):
373
380
.
Chapin
III FS
,
Hollingsworth
TN
,
Murray
DF
,
Viereck
LA
,
Walker
MD
.
2006
.
Floristic diversity and vegetation distribution in the Alaskan boreal forest
.
New York
:
Oxford University Press
.
Chapin
III FS
,
McGuire
A
,
Randerson
J
,
Pielke
R
,
Baldocchi
D
, et al
2000
.
Arctic and boreal ecosystems of western North America as components of the climate system
.
Glob Change Biol
6
(
S1
):
211
223
.
Chapin
III FS
,
McGuire
AD
,
Ruess
R
,
Hollingsworth
T
,
Mack
M
, et al
2010
.
Resilience of Alaska’s boreal forest to climatic change
.
Can J Forest Res
40
(
7
):
1360
1370
.
Chapin
III FS
,
Sturm
M
,
Serreze
MC
,
McFadden
JP
,
Key
JR
, et al
2005
.
Role of land-surface changes in Arctic summer warming
.
Science
310
:
657
660
.
Chapin
III FS
,
Trainor
SF
,
Huntington
O
,
Lovecraft
AL
,
Zavaleta
E
, et al
2008
.
Increasing wildfire in Alaska’s boreal forest: pathways to potential solutions of a wicked problem
.
Bioscience
58
(
6
):
531
540
.
Chapman
WL
,
Walsh
JE
.
2007
.
Simulations of Arctic temperature and pressure by global coupled models
.
J Climate
20
(
4
).
Chivers
MR
,
Turetsky
MR
,
Waddington
JM
,
Harden
JW
,
McGuire
AD
.
2009
.
Effects of experimental water table and temperature manipulations on ecosystem CO2 fluxes in an Alaskan rich fen
.
Ecosystems
12
(
8
):
1329
1342
.
Cleve
KV
,
Dyrness
C
,
Marion
G
,
Erickson
R
.
1993
.
Control of soil development on the Tanana River floodplain, interior Alaska
.
Can J Forest Res
23
(
5
):
941
955
.
Cole
JJ
,
Caraco
NF
,
Kling
GW
,
Kratz
TK
.
1994
.
Carbon dioxide supersaturation in the surface waters of lakes
.
Science
265
(
5178
):
1568
.
Conn
JS
,
Werdin-Pfisterer
NR
,
Beattie
KL
,
Densmore
RV
.
2011
.
Ecology of invasive Melilotus albus on Alaskan glacial river floodplains
.
Arct Antarct Alp Res
43
(
3
):
343
354
.
Cooper
LW
,
Benner
R
,
McClelland
JW
,
Peterson
BJ
,
Holmes
RM
, et al
2005
.
Linkages among runoff, dissolved organic carbon, and the stable oxygen isotope composition of seawater and other water mass indicators in the Arctic Ocean
.
J Geophys Res Biogeosci
110
(
G2
).
Cortés-Burns
H
,
Lapina
I
,
Klein
S
,
Carlson
M
,
Flagstad
L
.
2007
.
Invasive plant species monitoring and control-areas impacted by 2004 and 2005 fires in interior Alaska: A survey of Alaska BLM lands along the Dalton, Steese, and Taylor Highways
.
Report funded by the Alaska State Office, Bureau of Land Management, US Department of the interior. Anchorage, AK
.
Couwenberg
J
,
Thiele
A
,
Tanneberger
F
,
Augustin
J
,
Bärisch
S
, et al
2011
.
Assessing greenhouse gas emissions from peatlands using vegetation as a proxy
.
Hydrobiologia
674
(
1
):
67
89
.
Davidson
EA
,
Janssens
IA
.
2006
.
Temperature sensitivity of soil carbon decomposition and feedbacks to climate change
.
Nature
440
(
7081
):
165
173
.
Derksen
C
,
Brown
R
.
2012
:
Spring snow cover extent reductions in the 2008–2012 period exceeding climate model projections
.
Geophys Res Lett
39
:
L19504
.
Dittmar
T
,
Kattner
G
.
2003
.
The biogeochemistry of the river and shelf ecosystem of the Arctic Ocean: A review
.
Mar Chem
83
(
3
):
103
120
.
Dorrepaal
E
,
Aerts
R
,
Cornelissen
J
,
Van
Logtestijn R
,
Callaghan
T
.
2006
.
Sphagnum modifies climate-change impacts on subarctic vascular bog plants
.
Funct Ecol
20
(
1
):
31
41
.
Douglas
T
,
Jorgenson
MT
,
Kanevskiy
M
,
Romanovsky
, et al
2008
,
Permafrost dynamics at the Fairbanks Permafrost Experimental Station near Fairbanks, Alaska
,
Proceedings of the Ninth International Conference on Permafrost
,
edited by D. Kane, and K. Hinkle, pp. 373
.
Douglas
TA
,
Blum
JD
,
Guo
L
,
Keller
K
,
Gleason
JD
.
2013
.
Hydrogeochemistry of seasonal flow regimes in the Chena River, a subarctic watershed draining discontinuous permafrost in interior Alaska (USA)
.
Chem Geol
335
:
48
62
.
Douglas
TA
,
Fortier
D
,
Shur
YL
,
Kanevskiy
MZ
,
Guo
L
, et al
2011
.
Biogeochemical and geocryological characteristics of wedge and thermokarst-cave ice in the CRREL permafrost tunnel, Alaska
.
Permafrost Periglac
22
(
2
):
120
128
.
Duffy
PA
,
Walsh
JE
,
Graham
JM
,
Mann
DH
,
Rupp
TS
.
2005
.
Impacts of large-scale atmospheric-ocean variability on Alaskan fire season severity
.
Ecol Appl
15
(
4
):
1317
1330
.
Dyrness
C
,
Viereck
L
,
Van Cleve
K
.
1986
.
Fire in taiga communities of interior Alaska, in, Forest ecosystems in the Alaskan taiga
.
Springer
: pp.
74
86
.
Euskirchen
E
,
McGuire
A
,
Chapin
FS
,
Rupp
T
.
2010
.
The changing effects of Alaska’s boreal forests on the climate system
.
Can J Forest Res
40
(
7
):
1336
1346
.
Euskirchen
E
,
McGuire
A
,
Chapin
III F
,
Yi
S
,
Thompson
C
.
2009
.
Changes in vegetation in northern Alaska under scenarios of climate change, 2003–2100: Implications for climate feedbacks
.
Ecol Appl
19
(
4
):
1022
1043
.
Euskirchen
E
,
McGuire
AD
,
Kicklighter
DW
,
Zhuang
Q
,
Clein
JS
, et al
2006
.
Importance of recent shifts in soil thermal dynamics on growing season length, productivity, and carbon sequestration in terrestrial high-latitude ecosystems
.
Glob Change Biol
12
(
4
):
731
750
.
Euskirchen
ES
,
Carman
TB
,
McGuire
AD
.
2014
.
Changes in the structure and function of northern Alaskan ecosystems when considering variable leaf-out times across groupings of species in a dynamic vegetation model
.
Global Change Biology
20
(
3
):
963
978
.
Fan
Z
,
McGuire
AD
,
Turetsky
MR
,
Harden
JW
,
Waddington
JM
, et al
2013
.
The response of soil organic carbon of a rich fen peatland in interior Alaska to projected climate change
.
Glob Change Biol
19
(
2
):
604
620
.
Field
CB
,
Lobell
DB
,
Peters
HA
,
Chiariello
NR
.
2007
.
Feedbacks of Terrestrial Ecosystems to Climate Change
.
Annual Rev Environ Resour
32
:
1
29
.
Finlay
J
,
Neff
J
,
Zimov
S
,
Davydova
A
,
Davydov
S
.
2006
.
Snowmelt dominance of dissolved organic carbon in high-latitude watersheds: Implications for characterization and flux of river DOC
.
Geophys Res Lett
33
(
10
).
Fleming
RA
.
2000
.
Climate change and insect disturbance regimes in Canada’s boreal forests
.
World Resour Rev
12
(
3
):
521
548
.
Frey
KE
,
McClelland
JW
.
2009
.
Impacts of permafrost degradation on arctic river biogeochemistry
.
Hydrol Process
23
(
1
):
169
182
.
Frey
KE
,
Siegel
DI
,
Smith
LC
.
2007
.
Geochemistry of west Siberian streams and their potential response to permafrost degradation
.
Water Resour Res
43
(
3
).
Frey
KE
,
Smith
LC
.
2005
.
Amplified carbon release from vast West Siberian peatlands by 2100
.
Geophys Res Lett
32
(
9
).
Frolking
S
,
Roulet
N
,
Fuglestvedt
J
.
2006
.
How northern peatlands influence the Earth’s radiative budget: Sustained methane emission versus sustained carbon sequestration
.
J Geophys Res Biogeosci
111
(
G1
).
Genet
H
,
McGuire
A
,
Barrett
K
,
Breen
A
,
Euskirchen
E
, et al
2013
.
Modeling the effects of fire severity and climate warming on active layer thickness and soil carbon storage of black spruce forests across the landscape in interior Alaska
.
Environ Res Lett
8
(
4
):
045016
.
Goodale
CL
,
Apps
MJ
,
Birdsey
RA
,
Field
CB
,
Heath
LS
, et al
2002
.
Forest carbon sinks in the Northern Hemisphere
.
Ecol Appl
12
(
3
):
891
899
.
Goulden
ML
,
McMillan
A
,
Winston
G
,
Rocha
A
,
Manies
K
, et al
2011
.
Patterns of NPP, GPP, respiration, and NEP during boreal forest succession
.
Glob Change Biol
17
(
2
):
855
871
.
Goulden
ML
,
Wofsy
SC
,
Harden
JW
,
Trumbore
SE
,
Crill
PM
, et al
1998
.
Sensitivity of boreal forest carbon balance to soil thaw
.
Science
279
(
5348
):
214
217
.
Gower
S
,
Krankina
O
,
Olson
R
,
Apps
M
,
Linder
S
, et al
2001
.
Net primary production and carbon allocation patterns of boreal forest ecosystems
.
Ecol Appl
11
(
5
):
1395
1411
.
Grant
RF
.
2004
.
Modeling topographic effects on net ecosystem productivity of boreal black spruce forests
.
Tree Physiol
24
(
1
):
1
18
.
Grosse
G
,
Harden
J
,
Turetsky
M
,
McGuire
AD
,
Camill
P
, et al
2011
.
Vulnerability of high-latitude soil organic carbon in North America to disturbance
.
J Geophys Res Biogeosci
116
(
G4
).
Guo
L
,
Cai
Y
,
Belzile
C
,
Macdonald
RW
.
2012
.
Sources and export fluxes of inorganic and organic carbon and nutrient species from the seasonally ice-covered Yukon River
.
Biogeochemistry
107
(
1–3
):
187
206
.
Guo
L
,
Macdonald
RW
.
2006
.
Source and transport of terrigenous organic matter in the upper Yukon River: evidence from isotope (δ13C, Δ14C, and δ15N) composition of dissolved, colloidal, and particulate phases
.
Global Biogeochem Cy
20
(
2
).
Guo
L
,
Ping
C
,
Macdonald
RW
.
2007
.
Mobilization pathways of organic carbon from permafrost to arctic rivers in a changing climate
.
Geophys Res Lett
34
(
13
).
Güsewell
S
,
Koerselman
W
,
Verhoeven
JT
.
2002
.
Time-dependent effects of fertilization on plant biomass in floating fens
.
J Veg Sci
13
(
5
):
705
718
.
Gutsell
S
,
Johnson
EA
.
2002
.
Accurately ageing trees and examining their height-growth rates: implications for interpreting forest dynamics
.
J Ecol
90
:
153
166
.
Hamilton
TD
,
Ager
TA
,
Robinson
SW
.
1983
.
Late Holocene ice wedges near Fairbanks, Alaska, USA: environmental setting and history of growth
.
Arctic Alpine Res
157
168
.
Han
E
,
Bauce
E
,
Trempe-Bertrand
F
.
2000
.
Development of the first-instar spruce budworm (Lepidoptera: Tortricidae)
.
Ann Entomol Soc Am
93
(
3
):
536
540
.
Harden
J
,
Trumbore
S
,
Stocks
B
,
Hirsch
A
,
Gower
S
, et al.
2000
.
The role of fire in the boreal carbon budget
.
Glob Change Biol
6
(
S1
):
174
184
.
Harden
JW
,
Koven
CD
,
Ping
C
,
Hugelius
G
,
David McGuire
A
, et al
2012
.
Field information links permafrost carbon to physical vulnerabilities of thawing
.
Geophys Res Lett
39
(
15
).
Harden
JW
,
Manies
KL
,
O’Donnell
J
,
Johnson
K
,
Frolking
S
, et al
2012
.
Spatiotemporal analysis of black spruce forest soils and implications for the fate of C
.
J Geophys Res Biogeosci
117
(
G1
).
Harden
JW
,
Manies
KL
,
Turetsky
MR
,
Neff
JC
.
2006
.
Effects of wildfire and permafrost on soil organic matter and soil climate in interior Alaska
.
Glob Change Biol
12
(
12
):
2391
2403
.
Harden
JW
,
Meier
R
,
Silapaswan
C
,
Swanson
DK
,
McGuire
AD
.
2001
.
Soil drainage and its potential for influencing wildfires in Alaska
.
Studies by the US Geological Survey in Alaska
.
pp. 139–144
.
Hartley
IP
,
Garnett
MH
,
Sommerkorn
M
,
Hopkins
DW
,
Fletcher
BJ
, et al
2012
.
A potential loss of carbon associated with greater plant growth in the European Arctic
.
Nature Climate Change
2
(
12
):
875
879
.
Harvey
LD
.
1988
.
On the role of high latitude ice, snow, and vegetation feedbacks in the climatic response to external forcing changes
.
Climatic Change
13
(
2
):
191
224
.
Hayes
DJ
,
McGuire
AD
,
Kicklighter
DW
,
Gurney
KR
,
Burnside
T
, et al
2011
.
Is the northern high-latitude land-based CO2 sink weakening?
Global Biogeochem Cy
25
(
3
).
Hicke
JA
,
Allen
CD
,
Desai
AR
,
Dietze
MC
,
Hall
RJ
, et al
2012
.
Effects of biotic disturbances on forest carbon cycling in the United States and Canada
.
Glob Change Biol
18
(
1
):
7
34
.
Higgins
PA
,
Harte
J
.
2012
.
Carbon Cycle Uncertainty Increases Climate Change Risks and Mitigation Challenges
.
J Clim
25
(
21
).
Hinzman
L
,
Kane
D
,
Benson
C
,
Everett
K
.
1996
.
Energy balance and hydrological processes in an arctic watershed, in, Landscape Function and Disturbance in Arctic Tundra
.
Springer
: pp.
131
154
.
Hinzman
L
,
Kane
D
,
Gieck
R
,
Everett
K
.
1991
.
Hydrologic and thermal properties of the active layer in the Alaskan Arctic
.
Cold Reg Sci Technol
19
(
2
):
95
110
.
Hinzman
L
,
Kane
D
,
Yoshikawa
K
,
Carr
A
,
Bolton
W
, et al
2003b
.
Hydrological variations among watersheds with varying degrees of permafrost
,
Proceedings of the Eighth International Conference on Permafrost
,
pp. 21
.
Hinzman
LD
,
Bettez
ND
,
Bolton
WR
,
Chapin
FS
,
Dyurgerov
MB
, et al
2005
.
Evidence and implications of recent climate change in northern Alaska and other arctic regions
.
Climatic Change
72
(
3
):
251
298
.
Hinzman
LD
,
Fukuda
M
,
Sandberg
DV
,
Chapin
FS
,
Dash
D
.
2003a
.
FROSTFIRE: An experimental approach to predicting the climate feedbacks from the changing boreal fire regime
.
J Geophys Res Atmos
108
(
D1
): FFR 9–1-FFR 9–6.
Hinzman
LD
,
Goering
DJ
,
Kane
DL
.
1998
.
A distributed thermal model for calculating soil temperature profiles and depth of thaw in permafrost regions
.
J Geophys Res Atmos
103
(
D22
):
28975
28991
.
Hobbie
SE
,
Nadelhoffer
KJ
,
Högberg
P
.
2002
.
A synthesis: the role of nutrients as constraints on carbon balances in boreal and arctic regions
.
Plant Soil
242
(
1
):
163
170
.
Holden
J
,
Burt
T
.
2003
.
Hydraulic conductivity in upland blanket peat: measurement and variability
.
Hydrol Process
17
(
6
):
1227
1237
.
Hollingsworth
TN
,
Lloyd
AH
,
Nossov
DR
,
Ruess
RW
,
Charlton
BA
, et al
2010
.
Twenty-five years of vegetation change along a putative successional chronosequence on the Tanana River, Alaska
.
Can J Forest Res
40
(
7
):
1273
1287
.
Hollingsworth
TN
,
Schuur
EAG
,
Chapin
III FS
,
Walker
MD
.
2008
.
Plant community composition as a predictor of regional soil carbon storage in Alaskan boreal black spruce ecosystems
.
Ecosystems
11
(
4
):
629
642
.
Hollingsworth
TN
,
Walker
MD
,
Chapin
III FS
,
Parsons
AL
.
2006
.
Scale-dependent environmental controls over species composition in Alaskan black spruce communities
.
Can J For Res
36
(
7
):
1781
1796
.
Holmes
RM
,
McClelland
JW
,
Raymond
PA
,
Frazer
BB
,
Peterson
BJ
, et al
2008
.
Lability of DOC transported by Alaskan rivers to the Arctic Ocean
.
Geophys Res Lett
35
(
3
).
Homer
C
,
Dewitz
J
,
Fry
J
,
Coan
M
,
Hossain
N
, et al
2007
.
Completion of the 2001 National Land Cover Database for the Counterminous United States
.
Photogramm Eng Rem S
73
(
4
):
337
.
Hopkins
DM
,
Karlstrom
T
,
Black
R
,
Williams
J
,
Pewe
T
, et al
1955
.
Permafrost and ground water in Alaska
.
Geol Surv Prof Paper
:
113
146
.
Høye
TT
,
Post
E
,
Meltofte
H
,
Schmidt
NM
,
Forchhammer
MC
.
2007
.
Rapid advancement of spring in the High Arctic
.
Curr Biol
17
(
12
):
R449
R451
.
Intergovernmental Panel on Climate Change (IPCC)
.
2000
.
Special Report on Emissions Scenarios
.
N. Nakicenovic and R. Swart (eds.)
Cambridge University Press
.
Iwata
H
,
Ueyama
M
,
Harazono
Y
,
Tsuyuzaki
S
,
Kondo
M
, et al
2011
.
Quick recovery of carbon dioxide exchanges in a burned black spruce forest in interior Alaska
.
Sola
(
7
):
105
108
.
Jafarov
E
,
Romanovsky
V
,
Genet
H
,
McGuire
A
,
Marchenko
S
.
2013
.
The effects of fire on the thermal stability of permafrost in lowland and upland black spruce forests of interior Alaska in a changing climate
.
Environ Res Lett
8
(
3
):
035030
.
Johnson
KD
,
Harden
J
,
McGuire
AD
,
Bliss
NB
,
Bockheim
JG
, et al
2011
.
Soil carbon distribution in Alaska in relation to soil-forming factors
.
Geoderma
16
:
771
84
.
Johnstone
JF
,
Chapin
FS III
.
2003
.
Non-equilibrium succession dynamics indicate continued northern migration of lodgepole pine
.
Glob Change Biol
9
(
10
):
1401
1409
.
Johnstone
JF
,
Chapin
III FS
.
2006
.
Effects of soil burn severity on post-fire tree recruitment in boreal forest
.
Ecosystems
9
(
1
):
14
31
.
Johnstone
JF
,
Hollingsworth
TN
,
Chapin
FS
,
Mack
MC
.
2010
.
Changes in fire regime break the legacy lock on successional trajectories in Alaskan boreal forest
.
Glob Change Biol
16
(
4
):
1281
1295
.
Johnstone
JF
,
Kasischke
ES
.
2005
.
Stand-level effects of soil burn severity on postfire regeneration in a recently burned black spruce forest
.
Can J Forest Res
35
(
9
):
2151
2163
.
Jones
MC
,
Booth
RK
,
Yu
Z
,
Ferry
P
.
2013
.
A 2200-year record of permafrost dynamics and carbon cycling in a collapse-scar bog, interior Alaska
.
Ecosystems
16
(
1
):
1
19
.
Jones
MC
,
Grosse
G
,
Jones
BM
,
Walter Anthony
K
.
2012
.
Peat accumulation in drained thermokarst lake basins in continuous, ice-rich permafrost, northern Seward Peninsula, Alaska
.
J Geophys Res Biogeosci
117
(
G2
).
Jorgenson
M
,
Osterkamp
T
.
2005
.
Response of boreal ecosystems to varying modes of permafrost degradation
.
Can J Forest Res
35
(
9
):
2100
2111
.
Jorgenson
MT
,
Racine
CH
,
Walters
JC
,
Osterkamp
TE
.
2001
.
Permafrost degradation and ecological changes associated with a warming climate in central Alaska
.
Climatic Change
48
(
4
):
551
579
.
Jorgenson
MT
,
Romanovsky
V
,
Harden
J
,
Shur
Y
,
O’Donnell
J
, et al
2010
.
Resilience and vulnerability of permafrost to climate change
.
Can J Forest Res
40
(
7
):
1219
1236
.
Jorgenson
MT
,
Roth
JE
,
Raynolds
MK
,
Smith
MD
,
Lentz
W
.
1999
.
An ecological land survey for Fort Wainwright, Alaska
.
U.S. Army Cold Regions Research and Engineering Laboratory Report CRREL-99–9
.
Jorgenson
MT
,
Shur
Y
.
2007
.
Evolution of lakes and basins in northern Alaska and discussion of the thaw lake cycle
.
J Geophys Res Earth Surf
112
(
F2
).
Jorgenson
MT
,
Yoshikawa
K
,
Kanevskiy
M
,
Shur
Y
,
Romanovsky
V
, et al
2008
,
Permafrost characteristics of Alaska
,
Institute of Northern Engineering
,
University of Alaska Fairbanks
.
Juutinen
S
,
Bubier
JL
,
Moore
TR
.
2010
.
Responses of vegetation and ecosystem CO2 exchange to 9 years of nutrient addition at Mer Bleue bog
.
Ecosystems
13
(
6
):
874
887
.
Kane
DL
,
Hinzman
LD
,
Zarling
JP
.
1991
.
Thermal response of the active layer to climatic warming in a permafrost environment
.
Cold Reg Sci Technol
19
(
2
):
111
122
.
Kane
E
,
Valentine
D
,
Schuur
EA
,
Dutta
K
.
2005
.
Soil carbon stabilization along climate and stand productivity gradients in black spruce forests of interior Alaska
.
Can J Forest Res
35
(
9
):
2118
2129
.
Kane
EZ
,
Turetsky
MR
,
Harden
JW
,
McGuire
AD
,
Waddington
JM
.
2010
.
Seasonal ice and hydrologic controls on dissolved organic carbon and nitrogen concentrations in a boreal-rich fen
.
J Geophys Res Biogeosci
115
(
G04012
). doi:
Kane
K
,
Slaughter
C
.
1973
.
Recharge of a central Alaska lake by subpermafrost groundwater
.
International Conference on Permafrost, 2d, Yakutsk, Siberia. Papers
.
Kanevskiy
M
,
Shur
Y
,
Fortier
D
,
Jorgenson
M
,
Stephani
E
.
2011
.
Cryostratigraphy of late Pleistocene syngenetic permafrost (yedoma) in northern Alaska, Itkillik River exposure
.
Quaternary Res
75
(
3
):
584
596
.
Karl
TR
,
Trenberth
KE
.
2003
.
Modern global climate change
.
Science
302
(
5651
):
1719
1723
. doi:
Kasischke
ES
,
French
NH
,
O’Neill
KP
,
Richter
DD
,
Bourgeau-Chavez
LL
, et al
2000a
.
Influence of fire on long-term patterns of forest succession in Alaskan boreal forests, in, Fire, Climate Change, and Carbon Cycling in the Boreal Forest
.
Springer
: pp.
214
235
.
Kasischke
ES
,
Johnstone
JF
.
2005
.
Variation in postfire organic layer thickness in a black spruce forest complex in interior Alaska and its effects on soil temperature and moisture
.
Can J Forest Res
35
(
9
):
2164
2177
.
Kasischke
ES
,
O’Neill
KP
,
French
NH
,
Bourgeau-Chavez
LL
.
2000b
.
Controls on patterns of biomass burning in Alaskan boreal forests, in, Fire, climate change, and carbon cycling in the boreal forest
.
Springer
: pp.
173
196
.
Kasischke
ES
,
Turetsky
MR
.
2006
.
Recent changes in the fire regime across the North American boreal region-Spatial and temporal patterns of burning across Canada and Alaska
.
Geophys Res Lett
33
(
9
).
Kasischke
ES
,
Verbyla
DL
,
Rupp
TS
,
McGuire
AD
,
Murphy
KA
, et al
2010
.
Alaska’s changing fire regime-implications for the vulnerability of its boreal forests
.
Can J Forest Res
40
(
7
):
1313
1324
.
Kawahigashi
M
,
Kaiser
K
,
Kalbitz
K
,
Rodionov
A
,
Guggenberger
G
.
2004
.
Dissolved organic matter in small streams along a gradient from discontinuous to continuous permafrost
.
Glob Change Biol
10
(
9
):
1576
1586
.
Keeling
CD
,
Chin
JFS
,
Whorf
TP
.
1996
.
Increased activity of northern vegetation inferred from atmospheric CO2 measurements
.
Nature
382
:
146
149
.
Klapstein
SJ
,
Turetsky
MR
,
McGuire
AD
,
Harden
JW
,
Czimczik
CI
, et al
2014
.
Controls on methane released through ebullition in peatlands affected by permafrost degradation
.
J Geophys Res Biogeosci
119
(
3
):
418
431
. doi:
Kling
GW
,
Kipphut
GW
,
Miller
MC
.
1991
.
Arctic lakes and streams as gas conduits to the atmosphere: implications for tundra carbon budgets
.
Science
251
(
4991
):
298
301
. doi:
Koven
CD
,
Ringeval
B
,
Friedlingstein
P
,
Ciais
P
,
Cadule
P
, et al
2011
.
Permafrost carbon-climate feedbacks accelerate global warming
.
P Natl Acad Sci USA
108
(
36
):
14769
14774
. doi:
Krankina
ON
,
Harmon
ME
,
Schnekenburger
F
,
Sierra
CA
.
2012
.
Carbon balance on federal forest lands of Western Oregon and Washington: the impact of the Northwest Forest Plan
.
Forest Ecol Manag
(
286
)
171
182
.
Laganière
J
,
Angers
DA
,
Paré
D
,
Bergeron
Y
,
Chen
HYH
.
2011
.
Black spruce soils accumulate more uncomplexed organic matter than aspen soils
.
Soil Sci Soc Amer J
75
(
3
):
1125
1132
.
Lal
R
.
2005
.
Forest soils and carbon sequestration
.
Forest Ecol Manag
220
(
1
):
242
258
.
Lavoie
M
,
Paré
D
,
Fenton
N
,
Groot
A
,
Taylor
K
.
2005
.
Paludification and management of forested peatlands in Canada: a literature review
.
Env Rev
13
(
2
):
21
50
.
Lawrence
DM
,
Slater
AG
,
Romanovsky
VE
,
Nicolsky
DJ
.
2008
.
Sensitivity of a model projection of near-surface permafrost degradation to soil column depth and representation of soil organic matter
.
J Geophys Res Earth Surf
113
(
F2
).
Lee
H
,
Schuur
EA
,
Inglett
KS
,
Lavoie
M
,
Chanton
JP
.
2012
.
The rate of permafrost carbon release under aerobic and anaerobic conditions and its potential effects on climate
.
Glob Change Biol
18
(
2
):
515
527
.
Limpens
J
,
Heijmans
MM
.
2008
.
Swift recovery of Sphagnum nutrient concentrations after excess supply
.
Oecologia
157
(
1
):
153
161
.
Liston
GE
,
Hiemstra
CA
.
2011
.
The Changing Cryosphere: Pan-Arctic Snow Trends (1979–2009)
.
J Clim
24
(
21
).
Litvak
M
,
Miller
S
,
Wofsy
SC
,
Goulden
M
.
2003
.
Effect of stand age on whole ecosystem CO2 exchange in the Canadian boreal forest
.
J Geophys Res-Atmos
108
(
D3
).
Liu
H
,
Randerson
JT
.
2008
.
Interannual variability of surface energy exchange depends on stand age in a boreal forest fire chronosequence
.
J Geophys Res Biogeosci
113
(
G1
).
Lloyd
AH
,
Yoshikawa
K
,
Fastie
CL
,
Hinzman
L
,
Fraver
M
.
2003
.
Effects of permafrost degradation on woody vegetation at arctic treeline on the Seward Peninsula, Alaska
.
Permafrost Periglac
14
(
2
):
93
101
.
Macdonald
RW
,
Yu
Y
.
2006
.
The Mackenzie estuary of the Arctic Ocean, in Estuaries
.
Berlin Heidelberg: Springer
: pp.
91
120
.
MacLean
R
,
Oswood
MW
,
Irons
III JG
,
McDowell
WH
.
1999
.
The effect of permafrost on stream biogeochemistry: a case study of two streams in the Alaskan (USA) taiga
.
Biogeochemistry
47
(
3
):
239
267
.
Mann
D
,
Fastie
C
,
Rowland
E
,
Bigelow
N
.
1995
.
Spruce succession, disturbance, and geomorphology on the Tanana River floodplain, Alaska
.
Ecoscience
2
(
2
):
184
199
.
Manninen
O
,
Stark
S
,
Kytöviita
M
,
Lampinen
L
,
Tolvanen
A
.
2009
.
Understorey plant and soil responses to disturbance and increased nitrogen in boreal forests
.
J Veg Sci
20
(
2
):
311
322
.
Marchenko
S
,
Romanovsky
V
,
Tipenko
G
.
2008
.
Numerical modeling of spatial permafrost dynamics in Alaska
.
Proceedings of the ninth international conference on permafrost
.
Institute of Northern Engineering
,
University of Alaska Fairbanks
,
pp. 1125
.
Marsh
P
,
Neumann
NN
.
2001
.
Processes controlling the rapid drainage of two ice-rich permafrost-dammed lakes in NW Canada
.
Hydrol Process
15
(
18
):
3433
3446
.
McGuire
AD
,
Anderson
LG
,
Christensen
TR
,
Dallimore
S
,
Guo
L
, et al
2009
.
Sensitivity of the carbon cycle in the Arctic to climate change
.
Ecol Monogr
79
(
4
):
523
555
.
McGuire
AD
,
Chapin
III F
,
Ruess
R
.
2010
.
The dynamics of change in Alaska’s boreal forests: resilience and vulnerability in response to climate warming
.
Can J Forest Res
40
(
7
).
McNamara
JP
,
Kane
DL
,
Hobbie
JE
,
Kling
GW
.
2008
.
Hydrologic and biogeochemical controls on the spatial and temporal patterns of nitrogen and phosphorus in the Kuparuk River, arctic Alaska
.
Hydrol Process
22
(
17
):
3294
3309
.
Moore
T
,
Knowles
R
.
1989
.
The influence of water table levels on methane and carbon dioxide emissions from peatland soils
.
Can J Soil Sci
69
(
1
):
33
38
.
Myers-Smith
I
,
Harden
J
,
Wilmking
M
,
Fuller
C
,
McGuire
A
, et al
2008
.
Wetland succession in a permafrost collapse: interactions between fire and thermokarst
.
Biogeosci
5
(
5
).
Myneni
RB
,
Keeling
CD
,
Tucker
CJ
,
Asrar
G
,
Nemani
RR
.
1997
.
Increased plant growth in the northern high latitudes from 1981–1991
.
Nature
386
:
698
702
.
Nams
VO
,
Folkard
NF
,
Smith
JN
.
1993
.
Effects of nitrogen fertilization on several woody and nonwoody boreal forest species
.
Can J Bot
71
(
1
):
93
97
.
Neff
J
,
Finlay
J
,
Zimov
S
,
Davydov
S
,
Carrasco
J
, et al
2006
.
Seasonal changes in the age and structure of dissolved organic carbon in Siberian rivers and streams
.
Geophys Res Lett
33
(
23
).
Nicholas
JR
,
Hinkel
KM
.
1996
.
Concurrent permafrost aggradation and degradation induced by forest clearing, central Alaska, USA
.
Arctic Alpine Res
28
.
Nossov
DR
,
Hollingsworth
TN
,
Ruess
RW
,
Kielland
K
.
2011
.
Development of Alnus tenuifolia stands on an Alaskan floodplain: patterns of recruitment, disease and succession
.
J Ecol
99
(
2
):
621
633
.
Nossov
DR
,
Jorgenson
MT
,
Kielland
K
,
Kanevskiy
MZ
.
2013
.
Edaphic and microclimatic controls over permafrost response to fire in interior Alaska
.
Environ Res Lett
8
(
3
):
035013
.
O’Donnell
JA
,
Aiken
GR
,
Walvoord
MA
,
Butler
KD
.
2012a
.
Dissolved organic matter composition of winter flow in the Yukon River basin: Implications of permafrost thaw and increased groundwater discharge
.
Global Biogeochem Cy
26
(
4
).
O’Donnell
JA
,
Harden
JW
,
McGuire
AD
,
Kanevskiy
MZ
,
Jorgenson
MT
, et al
2011
.
The effect of fire and permafrost interactions on soil carbon accumulation in an upland black spruce ecosystem of interior Alaska: implications for post-thaw carbon loss
.
Glob Change Biol
17
(
3
):
1461
1474
.
O’Donnell
JA
,
Jones
JB
.
2006
.
Nitrogen retention in the riparian zone of catchments underlain by discontinuous permafrost
.
Freshwater Biol
51
(
5
):
854
864
.
O’Neill
KP
,
Kasischke
ES
,
Richter
DD
.
2002
.
Environmental controls on soil CO2 flux following fire in black spruce, white spruce, and aspen stands of interior Alaska
.
Can J Forest Res
32
(
9
):
1525
1541
.
O’Neill
KP
,
Kasischke
ES
,
Richter
DD
.
2003
.
Seasonal and decadal patterns of soil carbon uptake and emission along an age sequence of burned black spruce stands in interior Alaska
.
J Geophys Res-Atmos
108
(
D1
).
O’Donnell
JA
,
Jorgenson
MT
,
Harden
JW
,
McGuire
AD
,
Kanevskiy
MZ
, et al
2012b
.
The effects of permafrost thaw on soil hydrologic, thermal, and carbon dynamics in an Alaskan peatland
.
Ecosystems
15
(
2
):
213
229
.
O’Donnell
JA
,
Romanofsky
VE
,
Harden
JW
,
McGuire
AD
.
2009
The effect of moisture content on the thermal conductivity of moss and organic soil horizons from black spruce ecosystems in interior Alaska
.
Soil Sci
174
:
646
651
.
O’Donnell
JA
,
Turetsky
MR
,
Harden
JW
,
Manies
KL
,
Pruett
LE
, et al
2009
.
Interactive effects of fire, soil climate, and moss on CO2 fluxes in black spruce ecosystems of interior Alaska
.
Ecosystems
12
(
1
):
57
72
.
Oechel
W
,
Van Cleve
K
.
1986
.
The role of bryophytes in nutrient cycling in the taiga, in, Forest ecosystems in the Alaskan taiga.
Springer
: pp.
121
137
.
Olefeldt
D
,
Turetsky
MR
,
Crill
PM
,
McGuire
AD
.
2013
.
Environmental and physical controls on northern terrestrial methane emissions across permafrost zones
.
Glob Change Biol
19
(
2
):
589
603
.
Osterkamp
T
,
Jorgenson
J
.
2006
.
Warming of permafrost in the Arctic National Wildlife Refuge, Alaska
.
Permafrost Periglac
17
(
1
):
65
69
.
Osterkamp
T
,
Jorgenson
M
,
Schuur
E
,
Shur
Y
,
Kanevskiy
M
, et al
2009
.
Physical and ecological changes associated with warming permafrost and thermokarst in interior Alaska
.
Permafrost Periglac
20
(
3
):
235
256
.
Osterkamp
T
,
Romanovsky
V
.
1999
.
Evidence for warming and thawing of discontinuous permafrost in Alaska
.
Permafrost Periglac
10
(
1
):
17
37
.
Osterkamp
T
,
Viereck
L
,
Shur
Y
,
Jorgenson
M
,
Racine
C
, et al
2000
.
Observations of thermokarst and its impact on boreal forests in Alaska, USA
.
Arct Antarct Alp Res
32
(
3
):
303
315
.
Osterkamp
T
.
2005
.
The recent warming of permafrost in Alaska
.
Global Planet Change
49
(
3
):
187
202
.
Osterkamp
T
.
2007
.
Characteristics of the recent warming of permafrost in Alaska
.
J Geophys Res Earth Surf
112
(
F2
).
Parmentier
F
,
Van Der Molen
M
,
Van Huissteden
J
,
Karsanaev
S
,
Kononov
A
, et al
2011
.
Longer growing seasons do not increase net carbon uptake in the northeastern Siberian tundra
.
J Geophys Res Biogeosci
116
(
G4
).
Peterson
BJ
,
Holmes
RM
,
McClelland
JW
,
Vorosmarty
CJ
,
Lammers
RB
, et al
2002
.
Increasing river discharge to the Arctic Ocean
.
Science
298
(
5601
):
2171
2173
. doi: .
Petrone
K
,
Hinzman
L
,
Shibata
H
,
Jones
J
,
Boone
R
.
2007
.
The influence of fire and permafrost on sub-arctic stream chemistry during storms
.
Hydrol Process
21
(
4
):
423
434
.
Petrone
KC
,
Jones
JB
,
Hinzman
LD
,
Boone
RD
.
2006
.
Seasonal export of carbon, nitrogen, and major solutes from Alaskan catchments with discontinuous permafrost
.
J Geophys Res Biogeosci
111
(
G2
).
Pielke
RA
,
Adegoke
J
,
Beltran-Przekurat
A
,
Hiemstra
CA
,
Lin
J
, et al
2007
.
An overview of regional land-use and land-cover impacts on rainfall
.
Tellus B
59
(
3
):
587
601
.
Ping
CL
,
Michaelson
GJ
,
Kimble
JM
.
1997
.
Carbon storage along a latitudinal transect in Alaska
.
Nutr Cycl Agroecosystems
49
(
1–3
):
235
242
.
Potter
C
.
2004
.
Predicting climate change effects on vegetation, soil thermal dynamics, and carbon cycling in ecosystems of interior Alaska
.
Ecol Model
175
(
1
):
1
24
.
Prokushkin
A
,
Gleixner
G
,
McDowell
W
,
Ruehlow
S
,
Schulze
E
.
2007
.
Source-and substrate-specific export of dissolved organic matter from permafrost-dominated forested watershed in central Siberia
.
Global Biogeochem Cy
21
(
4
).
Racine
CH
,
Walters
JC
.
1994
.
Groundwater-discharge fens in the Tanana Lowlands, interior Alaska, USA
.
Arctic Alpine Res
:
418
426
.
Räisänen
J
.
2008
.
Warmer climate: less or more snow?
Clim Dynam
30
(
2–3
):
307
319
.
Ramanathan
V
.
1988
.
The greenhouse theory of climate change: a test by an inadvertent global experiment
.
Science
240
(
4850
):
293
299
. doi:
Randerson
JT
,
Liu
H
,
Flanner
MG
,
Chambers
SD
,
Jin
Y
, et al
2006
.
The impact of boreal forest fire on climate warming
.
Science
314
(
5802
):
1130
1132
. doi:
Raymond
PA
,
McClelland
J
,
Holmes
R
,
Zhulidov
A
,
Mull
K
, et al
2007
.
Flux and age of dissolved organic carbon exported to the Arctic Ocean: A carbon isotopic study of the five largest arctic rivers
.
Global Biogeochem Cy
21
(
4
):
1
B4011
.
Rember
RD
,
Trefry
JH
.
2004
.
Increased concentrations of dissolved trace metals and organic carbon during snowmelt in rivers of the Alaskan Arctic
.
Geochim Cosmochim Ac
68
(
3
):
477
489
.
Reynolds
W
,
Brown
DA
,
Mathur
S
,
Overend
R
.
1992
.
Effect of in-situ gas accumulation on the hydraulic conductivity of peat
.
Soil Sci
153
(
5
):
397
408
.
Riordan
B
,
Verbyla
D
,
McGuire
AD
.
2006
.
Shrinking ponds in subarctic Alaska based on 1950–2002 remotely sensed images
.
J Geophys Res Biogeosci
111
(
G4
).
Roach
J
,
Griffith
B
,
Verbyla
D
,
Jones
J
.
2011
.
Mechanisms influencing changes in lake area in Alaskan boreal forest
.
Glob Change Biol
17
(
8
):
2567
2583
.
Rober
AR
,
Wyatt
KH
,
Stevenson
RJ
,
Turetsky
MR
.
2014
.
Spatial and temporal variability of algal community dynamics and productivity in floodplain wetlands along the Tanana River, Alaska
.
Freshwater Science
33
(
3
):
765
777
.
Robinson
SD
,
Moore
TR
.
2000
.
The influence of permafrost and fire upon carbon accumulation in high boreal peatlands, Northwest Territories, Canada
.
Arctic, Antarctic, and Alpine Research
32
(
2
):
155
166
.
Romanovsky
V
,
Osterkamp
T
.
2001
.
Permafrost: changes and impacts, in, Permafrost response on economic development, environmental security and natural resources
.
Springer
: pp.
297
315
.
Romanovsky
VE
,
Smith
SL
,
Christiansen
HH
,
Shiklomanov
NI
,
Streletskiy
DA
, et al
2012
.
Permafrost
.
Arctic Report Card: Update for 2012
. http://www.arctic.noaa.gov/reportcard
Roulet
N
,
Moore
TIM
,
Bubier
J
,
Lafleur
P
.
1992
.
Northern fens: methane flux and climatic change
.
Tellus B
44
(
2
):
100
105
.
Rover
J
,
Ji
L
,
Wylie
BK
,
Tieszen
LL
.
2011
.
Establishing water body areal extent trends in interior Alaska from multitemporal Landsat data
.
Remote Sensing Letters
3
(
7
):
595
604
.
Rydin
H
,
McDonald
A
.
1985
.
Tolerance of Sphagnum to water level
.
J Bryol
13
(
4
):
571
578
.
Scenarios Network for Alaska Planning
.
2010
.
Climate change impacts on water availability in Alaska
. Available from http://www.snap.uaf.edu
Schaefer
K
,
Lantuit
H
,
Romanovsky
VE
,
Schuur
EAG
,
Witt
R
.
2014
.
The impact of the permafrost carbon feedback on global climate
.
Environmental Research Letters
9
(
8
):
085003
.
Schaefer
K
,
Zhang
T
,
Bruhwiler
L
,
Barrett
AP
.
2011
.
Amount and timing of permafrost carbon release in response to climate warming
.
Tellus B
63
(
2
):
165
180
.
Schlesinger
WH
.
1997
.
Biogeochemistry: An Analysis of Global Change.
Schuur
EA
,
Bockheim
J
,
Canadell
JG
,
Euskirchen
E
,
Field
CB
, et al
2008
.
Vulnerability of permafrost carbon to climate change: Implications for the global carbon cycle
.
Bioscience
58
(
8
):
701
714
.
Schuur
EA
,
Vogel
JG
,
Crummer
KG
,
Lee
H
,
Sickman
JO
, et al
2009
.
The effect of permafrost thaw on old carbon release and net carbon exchange from tundra
.
Nature
459
(
7246
):
556
559
.
Schuur
EAG
,
Abbott
B
.
2011
.
The Permafrost Carbon Network
.
Nature
480
:
32
33
.
Schuur
EAG
,
Abbott
BW
,
Bowden
WB
,
Brovkin
V
,
Camill
P
, et al
2013
.
Expert assessment of vulnerability of permafrost carbon to climate change
.
Climatic Change
119
(
2
):
359
374
.
Schuur
EAG
,
Crummer
KG
,
Vogel
JG
,
Mack
MC
.
2007
.
Plant species composition and productivity following permafrost thaw and thermokarst in Alaskan tundra
.
Ecosystems
10
(
2
):
280
292
.
Shakhova
N
,
Semiletov
I
,
Salyuk
A
,
Yusupov
V
,
Kosmach
D
, et al.
2010
.
Extensive methane venting to the atmosphere from sediments of the East Siberian Arctic Shelf
.
Science
327
(
5970
):
1246
1250
. doi:
Shakhova
N
,
Semiletov
I
.
2007
.
Methane release and coastal environment in the East Siberian Arctic shelf
.
J Marine Syst
66
(
1
):
227
243
.
Shenoy
A
,
Johnstone
JF
,
Kasischke
ES
,
Kielland
K
.
2011
.
Persistent effects of fire severity on early successional forests in interior Alaska
.
Forest Ecol Manag
261
(
3
):
381
390
.
Shulski
M
,
Wendler
G
.
2007
.
The Climate of Alaska
.
University of Alaska Press
.
Shur
Y
,
Jorgenson
M
.
2007
.
Patterns of permafrost formation and degradation in relation to climate and ecosystems
.
Permafrost Periglac
18
(
1
):
7
19
.
Sierra
C
,
Harmon
M
,
Thomann
E
,
Perakis
S
,
Loescher
H
.
2010
.
Amplification and dampening of soil respiration by changes in temperature variability
.
Biogeosci Discuss
7
(
6
).
Smith
LC
,
Sheng
Y
,
MacDonald
GM
,
Hinzman
LD
.
2005
.
Disappearing Arctic lakes
.
Science
308
(
5727
):
1429
. doi:
Solomon
S
,
Qin
D
,
Manning
M
,
Chen
Z
,
Marquis
M
, et al.
2007
.
Climate Change 2007: The physical science basis. Contribution of Working Group I to the fourth assessment report of the Intergovernmental Panel on Climate Change
.
Cambridge University Press
.
Sparrow
SD
,
Cochran
VL
,
Sparrow
EB
.
1993
.
Herbage yield and nitrogen accumulation by seven legume crops on acid and neutral soils in a subarctic environment
.
Can J Plant Sci
73
(
4
):
1037
1045
.
Sparrow
SD
,
Cochran
VL
,
Sparrow
EB
.
1995
.
Dinitrogen fixation by seven legume crops in Alaska
.
Agron J
87
(
1
):
34
41
.
Stieglitz
M
,
Giblin
A
,
Hobbie
J
,
Williams
M
,
Kling
G
.
2000
.
Simulating the effects of climate change and climate variability on carbon dynamics in Arctic tundra
.
Global Biogeochem Cy
14
(
4
):
1123
1136
.
Stone
RS
,
Dutton
EG
,
Harris
JM
,
Longenecker
D
.
2002
.
Earlier spring snowmelt in northern Alaska as an indicator of climate change
.
J Geophys Res: Atmospheres (1984–2012)
107
(
D10
): ACL-10.
Striegl
RG
,
Aiken
GR
,
Dornblaser
MM
,
Raymond
PA
,
Wickland
KP
.
2005
.
A decrease in discharge-normalized DOC export by the Yukon River during summer through autumn
.
Geophys Res Lett
32
(
21
).
Striegl
RG
,
Dornblaser
MM
,
Aiken
GR
,
Wickland
KP
,
Raymond
PA
.
2007
.
Carbon export and cycling by the Yukon, Tanana, and Porcupine rivers, Alaska, 2001–2005
.
Water Resour Res
43
(
2
).
Stutter
M
,
Lumsdon
D
,
Thoss
V
.
2007
.
Physico-chemical and biological controls on dissolved organic matter in peat aggregate columns
.
Eur J Soil Sci
58
(
3
):
646
657
.
Symstad
AJ
,
Chapin
FS III
,
Wall
DH
,
Gross
KL
,
Huenneke
LF
, et al
2003
.
Long-term and large-scale perspectives on the relationship between biodiversity and ecosystem functioning
.
Bioscience
53
(
1
):
89
98
.
Takeuchi
N
.
2009
.
Temporal and spatial variations in spectral reflectance and characteristics of surface dust on Gulkana Glacier, Alaska Range
.
J Glaciol
55
(
192
):
701
709
.
Tarnocai
C
,
Campbell
I
.
2002
.
Soils of the polar regions, in, Encyclopaedia of Soil Science.
New York
:
Marcel Dekker
: pp.
1018
1021
.
Tarnocai
C
,
Canadell
J
,
Schuur
E
,
Kuhry
P
,
Mazhitova
G
, et al
2009
.
Soil organic carbon pools in the northern circumpolar permafrost region
.
Global Biogeochem Cy
23
(
2
).
Toniolo
H
,
Kodial
P
,
Hinzman
LD
,
Yoshikawa
K
.
2009
.
Spatio-temporal evolution of a thermokarst in interior Alaska
.
Cold Reg Sci Technol
56
(
1
):
39
49
.
Trumbore
SE
,
Harden
J
.
1997
.
Accumulation and turnover of carbon in organic and mineral soils of the BOREAS northern study area
.
J Geophys Res
102
(
24
):
28817
28830
.
Turetsky
M
,
Treat
C
,
Waldrop
M
,
Waddington
J
,
Harden
J
, et al
2008
.
Short-term response of methane fluxes and methanogen activity to water table and soil warming manipulations in an Alaskan peatland
.
J Geophys Res Biogeosci
113
(
G3
).
Turetsky
M
,
Wieder
R
,
Williams
C
,
Vitt
D
.
2000
.
Organic matter accumulation, peat chemistry, and permafrost melting in peatlands of boreal Alberta
.
Ecoscience
7
(
3
):
379
392
.
Turetsky
MR
,
Bond-Lamberty
B
,
Euskirchen
E
,
Talbot
J
,
Frolking
S
, et al
2012
.
The resilience and functional role of moss in boreal and arctic ecosystems
.
New Phytol
196
(
1
):
49
67
.
Turetsky
MR
, et al
2014
.
A synthesis of methane emissions from 71 northern, temperate, and subtroptical wetlands
.
Glob Change Biol
29
(
7
):
2183
2197
.
Turetsky
MR
,
Kane
ES
,
Harden
JW
,
Ottmar
RD
,
Manies
KL
, et al
2011
.
Recent acceleration of biomass burning and carbon losses in Alaskan forests and peatlands
.
Nature Geosci
4
(
1
):
27
31
.
Turetsky
MR
,
Mack
MC
,
Hollingsworth
TN
,
Harden
JW
.
2010
.
The role of mosses in ecosystem succession and function in Alaska’s boreal forest
.
Can J Forest Res
40
(
7
):
1237
1264
.
Turetsky
MR
,
Wieder
RK
,
Vitt
DH
.
2002
.
Boreal peatland C fluxes under varying permafrost regimes
.
Soil Biol Biochem
34
(
7
):
907
912
.
Turkington
R
,
John
E
,
Krebs
C
,
Dale
M
,
Nams
V
, et al
1998
.
The effects of NPK fertilization for nine years on boreal forest vegetation in northwestern Canada
.
J Veg Sci
9
(
3
):
333
346
.
Ueyama
M
,
Harazono
Y
,
Ohtaki
E
,
Miyata
A
.
2006
.
Controlling factors on the interannual CO2 budget at a subarctic black spruce forest in interior Alaska
.
Tellus B
58
(
5
):
491
501
.
Van Cleve
K
,
Viereck
L
.
1983
.
A comparison of successional sequences following fire on permafrost-dominated and permafrost-free sites in interior Alaska
.
Permafrost: Fourth International Conference, Proceedings
,
pp. 1286
.
Van Cleve
K
,
Viereck
LA
,
Schlentner
RL
.
1971
.
Accumulation of nitrogen in alder (Alnus) ecosystems near Fairbanks, Alaska
.
Arctic Alpine Res
101
114
.
Van
Huissteden J
,
Dolman
A
.
2012
.
Soil carbon in the Arctic and the permafrost carbon feedback
.
Curr Opinion Environ Sustain
4
(
5
):
545
551
.
Viereck
L
,
Dyrness
C
,
Foote
M
.
1993
.
An overview of the vegetation and soils of the floodplain ecosystems of the Tanana River, interior Alaska
.
Can J Forest Res
23
(
5
):
889
898
.
Viereck
LA
,
Dyrness
CT
,
Van
Cleve K
,
Foote
MJ
.
1983
.
Vegetation, soils, and forest productivity in selected forest types in interior Alaska
.
Can J Forest Res
13
(
5
):
703
720
.
Viereck
LA
.
1973a
.
Ecological effects of river flooding and forest fires on permafrost in the taiga of Alaska
.
International Conference on Permafrost, 2d, Yakutsk, Siberia
.
Viereck
LA
.
1973b
.
Wildfire in the taiga of Alaska
.
Quaternary Res
3
:
465
495
.
Viereck
LA
.
1982
.
Effects of fire and firelines on active layer thickness and soil temperatures in interior Alaska
.
Proceedings of the Fourth Canadian Permafrost Conference
.
National Research Council of Canada Calgary
:
Ottawa, Ontario
:
pp. 123–135
.
Volney
WJA
,
Fleming
RA
.
2000
.
Climate change and impacts of boreal forest insects
.
Agr Ecosyst Environ
82
(
1
):
283
294
.
Waelbroeck
C
.
1993
.
Climate-soil processes in the presence of permafrost: A systems modelling approach
.
Ecol Model
69
(
3
):
185
225
.
Wagner
D
,
DeFoliart
L
,
Doak
P
,
Schneiderheinze
J
.
2008
.
Impact of epidermal leaf mining by the aspen leaf miner (Phyllocnistis populiella) on the growth, physiology, and leaf longevity of quaking aspen
.
Oecologia
157
(
2
):
259
267
.
Walker
MD
,
Wahren
CH
,
Hollister
RD
,
Henry
GH
,
Ahlquist
LE
, et al
2006
.
Plant community responses to experimental warming across the tundra biome
.
P Natl Acad Sci USA
103
(
5
):
1342
1346
. doi:
Walsh
JE
,
Chapman
WL
,
Romanovsky
V
,
Christensen
JH
,
Stendel
M
.
2008
.
Global climate model performance over Alaska and Greenland
.
J Clim
21
(
23
).
Walter
K
,
Chanton
J
,
Chapin
F
,
Schuur
E
,
Zimov
S
.
2008
.
Methane production and bubble emissions from arctic lakes: Isotopic implications for source pathways and ages
.
J Geophys Res Biogeosci
113
(
G3
).
Walter
K
,
Zimov
S
,
Chanton
JP
,
Verbyla
D
,
Chapin
FS
.
2006
.
Methane bubbling from Siberian thaw lakes as a positive feedback to climate warming
.
Nature
443
(
7107
):
71
75
.
Walter
KM
,
Edwards
ME
,
Grosse
G
,
Zimov
SA
,
Chapin
FS III
.
2007a
.
Thermokarst lakes as a source of atmospheric CH4 during the last deglaciation
.
Science
318
(
5850
):
633
636
. doi:
Walter
KM
,
Smith
LC
,
Chapin
FS III
.
2007b
.
Methane bubbling from northern lakes: present and future contributions to the global methane budget
.
Philos Trans A Math Phys Eng Sci
365
(
1856
):
1657
1676
. doi:
Walter-Anthony
KM
,
Zimov
SA
,
Grosse
G
,
Jones
MC
,
Anthony
PM
, et al.
2014
.
A shift of thermokarst lakes from carbon sources to sinks during the Holocene epoch
.
Nature
511
(
7510
). doi:
Walters
JC
,
Racine
CH
,
Jorgenson
MT
.
1998
.
Characteristics of permafrost in the Tanana Flats
,
interior Alaska
,
in Permafrost: Seventh International Conference
, June, pp.
23
27
.
Walvoord
MA
,
Striegl
RG
.
2007
.
Increased groundwater to stream discharge from permafrost thawing in the Yukon River basin: Potential impacts on lateral export of carbon and nitrogen
.
Geophys Res Lett
34
(
12
).
Walvoord
MA
,
Voss
CI
,
Wellman
TP
.
2012
.
Influence of permafrost distribution on groundwater flow in the context of climate-driven permafrost thaw: Example from Yukon Flats Basin, Alaska, United States
.
Water Resour Res
48
(
7
).
Welp
L
,
Randerson
J
,
Liu
H
.
2006
.
Seasonal exchange of CO2 and δ18O-CO2 varies with postfire succession in boreal forest ecosystems
.
J Geophys Res Biogeosci
111
(
G3
).
Wendler
G
,
Conner
J
,
Moore
B
,
Shulski
M
,
Stuefer
M
.
2011
.
Climatology of Alaskan wildfires with special emphasis on the extreme year of 2004
.
Theor Appl Climatol
104
(
3–4
):
459
472
.
Wendler
G
,
Shulski
M
.
2009
.
A century of climate change for Fairbanks, Alaska
.
Arctic
62
(
3
).
White
D
,
Hinzman
L
,
Alessa
L
,
Cassano
J
,
Chambers
M
, et al
2007
.
The arctic freshwater system: Changes and impacts
.
J Geophys Res Biogeosci
112
(
G4
).
Wickland
KP
,
Neff
JC
.
2008
.
Decomposition of soil organic matter from boreal black spruce forest: environmental and chemical controls
.
Biogeochem
87
(
1
):
29
47
.
Wickland
KP
,
Striegl
RG
,
Neff
JC
,
Sachs
T
.
2006
.
Effects of permafrost melting on CO2 and CH4 exchange of a poorly drained black spruce lowland
.
J Geophys Res Biogeosci
111
(
G2
).
Windsor
J
,
Moore
TR
,
Roulet
NT
.
1992
.
Episodic fluxes of methane from subarctic fens
.
Can J Soil Sci
72
(
4
):
441
452
.
Wolken
JM
,
Hollingsworth
TN
,
Rupp
TS
,
Chapin
III FS
,
Trainor
SF
, et al
2011
.
Evidence and implications of recent and projected climate change in Alaska’s forest ecosystems
.
Ecosphere
2
(
11
):
art124
.
Woo
M
,
Thorne
R
,
Szeto
K
,
Yang
D
.
2008
.
Streamflow hydrology in the boreal region under the influences of climate and human interference
.
Phil Trans Royal Soc B: Bio Sci
363
(
1501
):
2249
2258
.
Woo
M
.
1986
.
Permafrost hydrology in North America
.
Atmosphere-Ocean
24
(
3
):
201
234
.
Woo
M
.
2000
. Permafrost and hydrology, in,
The Arctic: Environment, People, Policy
.
New Jersey
: pp.
57
96
.
Wu
J
.
2012
.
Response of peatland development and carbon cycling to climate change: a dynamic system modeling approach
.
Environ Earth Sci
65
(
1
):
141
151
.
Wurtz
TL
,
Macander
MJ
,
Spellman
BT
.
2010
. Spread of invasive plants from roads to river systems in Alaska: a network model, in
Pye
JM
, et al, eds.,
Advances in threat assessment and their application to forest and rangeland management
.
Wurtz
TL
,
Ott
RA
,
Maisch
JC
.
2006
.
Timber harvest in interior Alaska, in, Alaska’s Changing Boreal Forest
.
Oxford University Press
: pp.
302
308
.
Wyatt
KH
,
Stevenson
R
,
Turetsky
MR
.
2010
.
The importance of nutrient co-limitation in regulating algal community composition, productivity and algal-derived DOC in an oligotrophic marsh in interior Alaska
.
Freshwater Biol
55
(
9
):
1845
1860
.
Yarie
J
,
Billings
S
.
2002
.
Carbon balance of the taiga forest within Alaska: present and future
.
Can J Forest Res
32
(
5
):
757
767
.
Yarie
J
.
1981
.
Forest fire cycles and life tables: a case study from interior Alaska
.
Can J Forest Res
11
(
3
):
554
562
.
Yoshikawa
K
,
Hinzman
LD
.
2003
.
Shrinking thermokarst ponds and groundwater dynamics in discontinuous permafrost near Council, Alaska
.
Permafrost Periglac
14
(
2
):
151
160
.
Zhang
T
,
Frauenfeld
OW
,
Serreze
MC
,
Etringer
A
,
Oelke
C
, et al
2005b
.
Spatial and temporal variability in active layer thickness over the Russian Arctic drainage basin
.
J Geophys Res-Atmos
110
(
D16
).
Zhang
T
.
2005a
.
Influence of the seasonal snow cover on the ground thermal regime: An overview
.
Rev Geophys
43
(
4
).
Zhuang
Q
,
McGuire
AD
,
Melillo
J
,
Clein
J
,
Dargaville
R
, et al
2003
.
Carbon cycling in extratropical terrestrial ecosystems of the Northern Hemisphere during the 20th century: A modeling analysis of the influences of soil thermal dynamics
.
Tellus B
55
(
3
):
751
776
.
Zimov
S
,
Davydov
S
,
Zimova
G
,
Davydova
A
,
Schuur
E
, et al
2006a
.
Permafrost carbon: Stock and decomposability of a globally significant carbon pool
.
Geophys Res Lett
33
(
20
).
Zimov
SA
,
Schuur
EA
,
Chapin
III FS
.
2006b
.
Permafrost and the global carbon budget
.
Science
312
(
5780
):
1612
1613
.
Zoltai
S
,
Morrissey
L
,
Livingston
G
,
Groot
WD
.
1998
.
Effects of fires on carbon cycling in North American boreal peatlands
.
Env Rev
6
(
1
):
13
24
.

This project was supported by funding from the U.S. Army Corps of Engineers Responses to Climate Change and Actions for Change programs and the Department of Defense’s Strategic Environmental Research and Development Program’s Resource Conservation and Climate Change Program (project RC-2110).

Competing Interests

We are not aware of any competing interests for any of the authors.

This is an open-access article distributed under the terms of the Creative Commons Attribution License, which permits unrestricted use, distribution, and reproduction in any medium, provided the original author and source are credited.

Supplementary data