Climate
Climate of Interior Alaska
Interior Alaska is wedged between three mountain ranges: the Alaska Range to the south, the Brooks Range to the north, and the McKenzie Mountains (an extension of the Brooks Range/Rocky Mountain Range) to the east. These mountain ranges provide effective barriers to coastal air masses, producing a strongly continental climate with cold winters and warm relatively dry summers. These forests are characterized by drastic seasonal fluctuation in day length (more than 21 hours on June 21 and less than 4 hours on December 21), and temperature ranges from -50°C in January to +33°C in July, with a short growing season (135 days or less from early May to mid-September). Montane rain shadows and distance from the Bering Sea minimize precipitation from coastal storms, and the average annual precipitation is only 287 mm in Fairbanks, 35% of which falls as snow (Slaughter and Viereck 1986). Snow covers the ground from mid-October until mid- to late April, and maximum accumulation averages 75-100 cm (Viereck et al. 1993). Soil temperatures are consistently low. In interior Alaska, permafrost distribution and active layer thickness (portion of soil profile above permafrost that thaws and re-freezes annually) are closely related to the topographic conditions of slope, aspect, drainage, soil moisture content, thermal properties of the parent material, and vegetation.
Temperature and Precipitation
The continental climate of interior Alaska has a wide range of air temperature between summer and winter and large fluctuations around the seasonal means. Mean annual temperatures in the Tanana Valley average -3.1oC at the Fairbanks International Airport, with the warmest month, July, averaging 16.3oC, and the coldest (January) averaging -23.5°C (1917 to 2000 averages). However, these averages do not present a good picture of either the summer or winter air temperatures. For example, in the Tanana Valley, periods of extreme cold ranging in the vicinity of -40oC to -45oC are not uncommon at any time from late November through February. In contrast, daily maximum temperatures occasionally reach 35oC to 37oC in June and July, often with only modest night cooling because of persistent daylight.
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 Annual precipitation in interior Alaska is low and decreases from west to east, with a 50-year average for Fairbanks of 287 mm and a range from 142 mm in 1957 to 478 mm in 1990. Most summer and winter precipitation is generated from major frontal systems that cross the State, but convective storms add significantly to the summer precipitation. Precipitation events in early summer (May, June, and early July) are typically light and showery, with high spatial variability. The relatively dry summer conditions are replaced by the fall rain events, which can be heavy and sustained. On average, precipitation increases through the summer. There is considerable variability in annual precipitation in Alaska with low precipitation years, such as 1957, generating frequent wildfires, while high-precipitation years, such as 1967, often result in flooding. Although precipitation during the growing season may be low, evaporation rates are also low because of the relative short growing season and cool temperatures. Nonetheless, as much as 76 to 100 percent of the summer precipitation may be lost as evapotranspiration (Dingman 1966). Thus much of the summer precipitation probably derives from recycling of water that evaporated from land (Serreze and Etringer 2003).
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Snow is an important climate and ecological factor in the boreal forest. The ground is usually snow-covered from mid-October until early to mid-April, although in some years the snow period can be considerably longer. In 1992, for example, snow remained on the ground until the last week of May and returned permanently on 10 September giving only a 3-month snow-free period. Although snow accumulates during the entire winter period, maximum snowpack is generally shallow in interior Alaska with an average maximum snowpack depth in mid-March to early April of 75 cm depth with a water equivalent of 110 mm. There are occasional heavy snow years such as 1971, 1991 and 1993 when maximum snow accumulations may reach 2 meters. There are also low snow years, such as 1969-70, 1979-80, 1980-81, 1985-86, 1986-87, and 1998-99, when less than 40 cm may accumulate throughout the entire winter. Based on precipitation data from the Fairbanks Airport (1948-2000), snowfall accounts for about 35% of the yearly total precipitation (range 13-77%, standard deviation = 13%). Little winter snowmelt occurs due to typically below-freezing temperature throughout the winter. During the snowmelt period, (generally late April) snow is released as stream-flow over a relatively short period, making snowmelt the major hydrological event of the year. Incoming solar radiation (supplied by long daylight hours) is the major factor governing snowmelt as the albedo of the snow decreases from approximately 0.75-0.90 (fresh snow) to 0.40-0.70 (old snow, Campbell and Norman 1998). Melting of snow from the ground surface is negligible as ground temperatures are <0°C. During snowmelt, south-facing slopes typically have less snow than north-facing slopes (more interception by vegetation and greater sublimation losses), and the snow pack melts one to two weeks earlier. Spring snowmelt is the primary source of groundwater recharge in interior Alaska with infiltration occurring through frozen ground, primarily in areas absent of ice-rich permafrost (Hinzman et al. 2006).
The Warming and Drying of Interior Alaska
Warming in Alaska results from both natural climate variability and climate forcing caused by increases in greenhouse gas concentrations produced by human activities concentrated in temperate and tropical latitudes (Karl et al. 2009).
Since the early 20th century, interior Alaska has experienced an increase in mean annual temperature:
Temperature records from the Fairbanks International Airport record (mid 1948 to present) and the UAF Experiment Station (1904 to mid 1948) (Glenn Juday).
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Over the past 50 years, the rate of winter warming (2 °C decade-1) has exceeded the rate of summer warming (0.5 °C decade-1).
Climate records combined with dendrochronological climate reconstructions based on d13C of tree rings (Barber et al. 2004) indicate that the rate and length of time summers have been warming in Fairbanks is unprecedented over the past several centuries:
Annual precipitation has declined in Fairbanks over the past century:
Precipitation records from the Fairbanks International Airport
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This slight annual decline is the net result of gradual declines in late summer rainfall (left) but gradual increases in wintertime snowfall (right):
Overall, growing seasons have lengthened by 2.5 days decade-1 (Keyser et al. 2000), with most change occurring in spring (Euskirchen et al. 2009a):
Global Climate Change Impacts in the United States. 2009. Thomas R. Karl, Jerry M. Melillo, and Thomas C. Peterson, (eds.). Cambridge University Press.
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Effects of the Pacific Decadal Oscillation (PDO) on Alaskan Climate
Climate change in interior Alaska involves secular shifts in both mean annual temperatures and the seasonality of precipitation patterns and frequency of precipitation events, but also modes of long-term oscillations in global atmospheric pressure and sea surface temperatures that have strong effects on regional climate at decadal to multi-decadal scales.
Strengthening of the Aleutian Low during positive phases of the PDO increases advection of warm, moist air into interior Alaska, particularly during fall and winter (Hartmann and Wendler 2005). PDO expresses a bidecadal mode of variability that correlates with lightening strikes and fire distributions (Duffy et al. 2005, Fauria and Johnson 2006, 2008) as well as beetle outbreaks (Fauria and Johnson 2009). Tree-ring reconstructions indicate that climate-affected reversals in the PDO oscillations have occurred for at least the past 400 years (Biondi et al. 2001), and patterns of warming and cooling in interior Alaska over the past century correlate with major reversals in PDO:
Spring-summer PDO index obtained from the Joint Institute for the Study of the Atmosphere and Ocean ( JISAO;
http://tao.atmos.washington.edu/). Note the regime shifts in 1924, 1947 and 1976/1977 (arrows) (from Fauria and Johnson 2009).
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Landscape Consequences of a Warmer and Drier Climate
Longer summers and higher temperatures are causing drier conditions, even in the absence of strong trends in precipitation.
For example, white spruce forests in Alaska’s interior are experiencing declining growth due to drought stress and continued warming could lead to widespread death of trees (Barber et al. 2000):
Tree-ring properties in relation to Fairbanks climate during the 20th century. Smoothed (5-year running mean); all climate values are normalized with zero mean and scaled as standard deviation units. (b) Ringwidth and climate index, (d) d13C discrimination versus May through August temperature, (f) maximum late-wood density and density climate index. (from Barber et al. 2000).
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Another consequence of warmer, drier summers is the reduction in surface water, and associated shrinkage/disappearance of closed-basin ponds throughout the Alaskan boreal region as a result of increased evaporation and permafrost thaw (Riordan et al. 2006):
Original data from Riordan et al. (2006); graph on the right is from: Climate Change Impacts in the United States. 2009. Thomas R. Karl, Jerry M. Melillo, and Thomas C. Peterson, (eds.). Cambridge University Press.
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Other disturbance regimes affected by recent warming and drying include the recent increase in the size and severity of fire (Kasischke and Turetsky 2006), and the outbreak behavior of insects and pathogens (USDA 2008, Ruess et al. 2009), resulting in apparent threshold shifts in biogeochemical cycling, vegetation dynamics and successional pathways (Hollingsworth et al. submitted), and ecosystem and landscape function (Johnstone et al. submitted, Jorgenson et al. submitted).
Collectively, increased insect/disease outbreaks, reduced photosynthesis, growth and a change in root/leaf carbon allocation in response to warmer and drier growing season climate have led to a “browning of interior Alaska” as determined by satellite-detected trends in NDVI over the past several decades (Verbyla 2008). This is in contrast to the “greening of the Arctic” (Jia et al. 2003), which has likely resulted from the continued expansion of shrubs (Tape et al. 2006) and an increase in the NPP of wet meadows in response to a lengthened growing season.
Linear trends in annual (1982-2003) maximum normalized difference vegetation index (NDVI) values for each 64-km2 pixel in the study region. Only pixels with significant (P < 0.05) linear regression slopes are displayed. Albers equal area map projection (standard parallels 55° N, 65° N) (from Verbyla 2008).
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Feedbacks to Regional and Global Climate
Warming of the boreal forest is leading to significant climate feedbacks resulting from landform changes and associated atmospheric C, water and energy exchanges (McGuire and Chapin 2006, McGuire et al. 2006, Randerson et al. 2006, Riordan et al. 2006, McGuire et al. 2007, Euskirchen et al. 2009a, Euskirchen et al. 2009b, McGuire et al. 2009, Euskirchen et al. in press, submitted). Research generally suggests that the net effect of a warming climate is a positive regional feedback to warming (Euskirchen et al. submitted).
Currently, the primary positive climate feedback is likely related to changes in surface albedo due to decreases in snow cover resulting from a longer snow-free period (Figure A, B).
Figure A. Spatial distribution of changes in snow cover (days yr-1) for changes in snow melt (a), snow return (b), and changes in the total length of the snow cover duration (c) in the western Arctic between 1970 - 2000. From Euskirchen et al. (in press).
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Figure B. Changes in energy (W m-2 decade-1) due to changes in the timing of snow melt (a) changes in energy due changes in the timing of snow return (b), and changes in energy due to changes in the length of the snow cover duration in the western Arctic between 1970 - 2000. From Euskirchen et al. (in press).
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While negative feedbacks to climate have been quantified, including (1) those associated with increases in surface albedo due to a greater proportion of younger forests on the landscape under a heightened fire regime (Figure C,D), and (2) increased carbon uptake by the vegetation due to CO2 fertilization and longer growing seasons (Table A), these negative feedbacks may not be large enough to counterbalance the large positive feedbacks (Figure E).
Figure C. Changes in the distribution of the age of the ecosystems post-fire based on simulations with the ALFRESCO model and four different future climate scenarios. A trend toward younger aged forests indicates a trend towards a greater proportion of deciduous stands compared to coniferous stands across the landscape. A least square regression line is fit to the means of the data, with p < 0.0001 for all regressions in panels (a-c) . [a] = slope, [b] = intercept. From Euskirchen et al (2009b).
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Figure D. Changes in albedo due to the changes in the post-fire distributions of boreal forest ecosystems in the Western Arctic for the years 2003 - 2100. Shown are four different future climate scenarios, and the mean of the scenarios. For the summer albedo in (a), these changes are based on a synthesis of field collected data (Amiro et al. 2006)and for the summer albedo in (b), these changes are based on remotely sensed data (Lyons et al. 2008). For the winter albedo in (b), only the field measured data were used (Euskirchen et al, 2009b).
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Figure E. Summary of changes in atmospheric heating due to changes in land surface albedo and carbon and methane uptake/emissions in boreal Alaska, from available estimates. A negative sign represents a negative feedback for a sink term and a positive sign represents a positive feedback for a source term. Values are converted between Pg C to W m-2 based on published methodologies (Zhuang et al. 2006). 'Climate and CO2' refers to model simulations incorporating transient data pertaining to climate and atmospheric concentrations of carbon dioxide. Currently, the change in albedo due to the fire regime between 1970 - 2000 has not been estimated, as represented with the '?'. After Euskirchen et al., in press.
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Table A. Mean annual changes in carbon storage for Alaska from 2003-2100a driven by SRES A2 and B2 scenarios output by CGCM2 (Balshi et al. 2009). |
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Effects* |
Scenario |
Region |
CO2 |
Climate |
Fire |
Total |
With CO2 Fertilization |
A2 |
Alaska |
26.7 |
21.5 |
-12.0 |
36.2 |
B2 |
Alaska |
18.4 |
14.9 |
-9.4 |
23.9 |
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Without CO2 Fertilization |
A2 |
Alaska |
0.0 |
16.9 |
-11.0 |
5.9 |
B2 |
Alaska |
0.0 |
12.6 |
-8.3 |
4.3 |
*Units are in Tg C yr-1. Positive values indicate carbon sequestration by terrestrial ecosystems. Negative values indicate a release of carbon from land to atmosphere. |
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