In addition to the familiar daily, seasonal, and yearly fluctuations in weather, there are longer term natural variations in the Earth’s climate. Climate can be defined as the “average weather,” or more specifically, as “the statistical description in terms of the mean and variability of relevant quantities over a period of time ranging from months to thousands or millions of years” (IPCC 2007). Past variation in the Earth’s climate has been cyclical, as opposed to being random or following linear trends (fig. 1). It is important to understand this natural cyclical variability in climate when considering and evaluating human-caused climate change.

Figure 1—Variation in temperature and CO2 over the past 400,000 years. Source: Petit et al. 1999.

Figure 1—Variation in temperature and CO₂ over the past 400,000 years. Source: Petit et al. 1999.

Cycles in the Earth’s climate are nested and on multiple time scales, from year to year (interannual) to decades, centuries, and millennia. Various cycles are caused by independent physical mechanisms. Thus, for example, there are major glacial (cold) and interglacial (warm) periods on multimillennial time scales, caused by changes in the Earth’s orbit around the Sun. Other cycles in the Sun’s activity drive climate variations at the century scale. Cyclical patterns in circulation of the oceans and atmosphere lead to decadal (30 to 40 year) patterns, such as the Pacific Decadal Oscillation (PDO), which affects the west coast of North America. Cycles in the ocean-atmosphere system also lead to interannual variations in climate, such as the El-Niño/La Niña cycle (called ENSO, for El-Niño Southern Oscillation). Climate at any one time is an expression of all of these nested mechanisms and cycles operating together.

Multimillennial climate cycles—

Long-term climatic change is driven primarily by changes in solar radiation and atmospheric composition of gases such as CO₂. Variations in the Earth’s orbit influence the amount of solar radiation received at the surface. Several parameters of the Earth’s orbit change over time, including 1) eccentricity, or how elliptical (versus circular) the Earth’s orbit is around the sun; (2) tilt, or the angle of the Earth’s tilt on its axis; and (3) precession, a “wobbling” in Earth’s axis of rotation, resulting in variation in the time of year when the Earth is closest to the sun. The eccentricity, tilt, and precession of the Earth’s orbit vary on time scales of approximately 100,000, 41,000, and 23,000 years, respectively (Chapin et al. 2002) (fig. 2). Together, these variations in the Earth’s orbit produce the Milankovitch cycles of solar input. These cycles are strongly associated with the glacial and interglacial cycles over the last 800,000 years, which are determined from analysis of ocean sediments and ice cores.

The patterns of historical temperature changes associated with the glacial-interglacial cycles are also correlated with changes in atmospheric concentrations of carbon dioxide and methane, two greenhouse gases. Concentrations of carbon dioxide were relatively higher during warm interglacial periods in the past but decreased during cold glacial periods (fig. 1). The strong relationship between temperature and greenhouse gases suggests a mechanistic relationship. It is estimated that about half of the glacial–interglacial temperature change is due to greenhouse gas feedbacks (Petit et al. 1999). The potential CO2 increase through the 21st century may be sufficient (at the upper end of the uncertainty bounds) to produce a temperature increase on the magnitude of a full glacial – interglacial cycle (IPCC 2001).

Figure 2. Variation in parameters of the Earth's orbit over the last 250,000 years. From Chapin et al. 2001.

Figure 2—Variation in parameters of the Earth’s orbit over the last 250,000 years. Source: Chapin et al. 2001.

Century- to millennial-scale climate cycles—

In addition to multimillennial glacial and interglacial cycles, there are shorter cold-warm cycles that last from 100 to 1,000 years. These cycles, referred to as “Bond cycles,” have been documented for at least the last 130,000 years. The average length of a Bond cycle is 1300 to 1500 years, with each warm and cold phase of the cycle lasting for about 700 years. The Little Ice Age, a global cold period from 1450 to 1920, is an event that is thought to exemplify a cold phase of a Bond cycle (Grove 1988, Overpeck et al. 1997). Like the Milankovitch cycles, the Bond cycles are currently thought to be driven by changes in the sun (Bond et al. 2001).

Interannual- to decadal-scale climate cycles—

The well-known El-Niño Southern Oscillation (ENSO) is a large-scale cyclical change in the atmosphere-ocean system that occurs on interannual to decadal time scales. ENSO events are part of a large-scale air-sea interaction that couples atmospheric pressure changes (the southern oscillation) with changes in ocean temperature (El-Niño) over the equatorial Pacific Ocean (Chapin et al. 2002). Every few years, hemispheric trade winds that usually blow warm tropical ocean water in a westerly direction across the Pacific Ocean stall, resulting in warm water accumulating in the eastern Pacific Ocean. This leads to changes in water temperatures off the shore of North and South America. Each year there is some degree of El Niño, or its opposite effect, La Niña. On average, ENSO events occur every 4 to 7 years. El Niño events bring different conditions to different parts of the world. For example, El Niño events result in dry weather in the Pacific Northwest but wet weather in the Southwest U.S. (fig. 3). The reverse occurs during La Niña events.

Figure 3. Typical winter conditions in North America during El Niño and La Niña years. Source: NOAA Climate Prediction Center .

Figure 3—Typical winter conditions in North America during El Niño and La Niña years. Source: NOAA Climate Prediction Center (http://www.cpc.noaa.gov/products/analysis_monitoring/ensocycle/nawinter.shtml)

Recently, climate cycles on multidecadal timescales have also been described. The Pacific Decadal Oscillation (PDO), which affects western North America, seems to be regulated by decadal changes in ocean circulation patterns in the high-latitude Pacific Ocean. The effects of the PDO are similar to ENSO (Mantua et al. 1997), with warm (positive) phases and cool (negative) phases that last from 10 to 25 years (fig. 4). There are other decadal-scale ocean-mediated patterns that affect other parts of the world, such as the North Atlantic Oscillation (NAO).

Figure 4—Top: Typical wintertime sea surface temperature (colors), sea level pressure (contours), and surface wind stress (arrows) anomaly patterns during positive and negative phases of the Pacific Decadal Oscillation (PDO). Temperature anomalies (colors) are in degrees Celsius. Bottom: Monthly values for the PDO index, 1900-2004. Source: S. Hare and N. Mantua, Climate Impacts Group, Center for Science in the Earth System, Joint Institute for the Study of the Atmosphere and Ocean, University of Washington, Seattle.

Climate Mechanisms

Earth’s energy budget—

The Earth’s energy budget is the balance between incoming and outgoing radiation, which determines the amount of energy available to drive the Earth’s climate system (Chapin et al. 2002). About 30 percent of solar radiation that reaches Earth is reflected back into space by clouds, air molecules, dust, haze, and the Earth’s surface. Another 20 percent of incoming solar radiation is absorbed by the atmosphere. The remaining solar radiation reaches the Earth’s surface and is absorbed. The Earth’s surface radiates this energy back to the atmosphere in the form of infrared radiation. Most (90 percent) of this infrared radiation is trapped in the atmosphere by greenhouse gases, such as carbon dioxide (CO₂), methane (CH₄), and nitrous oxide (N₂O). The energy absorbed by the greenhouse gases is reradiated in all directions. The energy that is directed back towards the Earth’s surface contributes to the warming of the planet. This phenomenon is called the greenhouse effect (fig. 5).

Figure 5. The greenhouse effect. Source: Climate Change 1995, The Science of Climate Change, contribution of working group 1 to the second assessment report of the Intergovernmental Panel on Climate Change.

Figure 5—The greenhouse effect. Source: Climate Change 1995, The Science of Climate Change, contribution of working group 1 to the second assessment report of the Intergovernmental Panel on Climate Change.

Without the energy-absorbing greenhouse gases in the Earth’s atmosphere, the mean temperature at Earth’s surface would be about 33 ºC cooler than it is today and would probably not support life (Chapin et al. 2002). However, long-term records of the concentration of greenhouse gases in the atmosphere (from atmospheric measurements and ice core analysis) show steep increases in greenhouse gas concentrations since the beginning of the Industrial Revolution 150 years ago (fig. 6). These unprecedented increases in greenhouse gases are largely due to human activities, such as the burning of fossil fuels and industrial activities. As concentrations of greenhouse gases increase, more radiation emitted by the Earth is trapped by the atmosphere, thus enhancing the greenhouse effect and leading to increased temperatures at the Earth’s surface.

Figure 6—Variations in carbon dioxide concentrations in the Earth’s atmosphere over the last 400,000 years. Source: Robert A. Rohde and Global Warming Art.

Figure 6—Variations in carbon dioxide concentrations in the Earth’s atmosphere over the last 400,000 years. Source: Robert A. Rohde and Global Warming Art (http://www.globalwarmingart.com/wiki/Image:Carbon_Dioxide_400kyr_Rev_png)

Human influence on climate—

The 2007 United Nations Intergovernmental Panel on Climate Change (IPCC) Fourth Assessment Report (AR4) provides the most substantive and authoritative support for human-caused global climate change to date. This report, based mostly on peer-reviewed and published scientific literature, states that, "Warming of the climate system is unequivocal, as is now evident from observations of increases in global average air and ocean temperatures, widespread melting of snow and ice and rising global average sea level." The report further states that, "Most of the observed increase in globally averaged temperatures since the mid-20th century is very likely due to the observed increase in anthropogenic greenhouse gas concentrations."

According to the IPCC AR4 report, the global linear warming trend of 0.13 °C per decade over the 50 years 1956-2005 is nearly twice that for the 100 years 1906-2005 (fig. 7). In addition, 11 of the last 12 years (1995-2006) rank among the 12 warmest years in the instrumental record of global surface temperature (since 1850).

Observed rise in sea levels is consistent with global warming (fig. 7). According to the IPCC AR4, global average sea level has risen at an average rate of 1.8 mm/yr since 1961 and at 3.1 mm/yr since 1993, partly due to melting glaciers and ice. The observed decreases in snow and ice extent are also consistent with global warming (fig. 7). Satellite data show that annual average Arctic sea ice extent has shrunk by 2.7 percent per decade since 1978, with larger decreases of 7.4 percent per decade in the summer (IPCC 2007).

Long-term changes in other aspects of climate have also been observed. From 1900 to 2005, precipitation increased significantly in eastern parts of North and South America, northern Europe and northern and central Asia, whereas precipitation declined in the Sahel, the Mediterranean, southern Africa and parts of southern Asia (IPCC 2007). At a global scale, the area affected by drought has likely increased since the 1970s (IPCC 2007).

In addition to observed temperature and precipitation changes, analysis using General Circulation Models (GCM) of the Earth’s atmosphere is further evidence of human-caused global climate change. General Circulation Models (GCM) of the atmosphere are now being coupled with those of the oceans (AOGCM), ice, and the Earth’s terrestrial biosphere. These models have been under development for many decades. They spontaneously exhibit interannual and interdecadal oscillations like those observed in the real Earth system. They are run under different starting conditions and using different amounts of solar, volcanic and greenhouse gas "forcing" of the atmospheric dynamics. Using this "ensemble" approach, various AOGCMs have successfully simulated the Earth’s climate over the past 1,000 years. However, they cannot capture the rapid increase in global temperature of the past half century without including greenhouse gas forcing (IPCC 2007). Similarly, the models are able to simulate the warming of the upper 700 meters of all the major oceans of the world over the past 40 years, but only if they include the greenhouse gas emissions of the industrial age (IPCC 2007).

Past and future greenhouse gas emissions are expected to lead to further warming and changes in the global climate system in the future. Globally, average temperatures are expected to increase from 1.1 to 6.4 ºC by 2099 (IPCC 2007) (fig. 8). Further rise in sea level, decreasing extent of glaciers and ice, and increases in both drought and flooding are also expected globally owing to climate change. Effects of climate change will differ by geographic location. Predictions and potential implications of climate change specific to the Western United States are discussed in the next section.

Figure 7—Observed changes in (a) global average surface temperature; (b) global average sea level from tide gauge (blue) and satellite (red) data and (c) Northern Hemisphere snow cover for March-April. All differences are relative to corresponding averages for the period 1961-1990. Smoothed curves represent decadal averaged values and circles show yearly values. The shaded areas are the uncertainty intervals estimated from a comprehensive analysis of known uncertainties (a and b) and from the time series (c). Source: IPCC 2007.

Figure 8—Projected surface temperature changes for the late 21st century (2090-2099). The map shows the multi-AOGCM average projection for the A1B SRES scenario. Temperatures are relative to the period 1980-1999. Source: IPCC Climate Change 2007.

Figure 8—Projected surface temperature changes for the late 21^st century (2090-2099). The map shows the multi-AOGCM average projection for the A1B SRES scenario. Temperatures are relative to the period 1980-1999. Source: IPCC Climate Change 2007.

Effects and Implications of Climate Change in the Western United States

Temperature and Precipitation

Over almost the entire Western United States, there have been increases in both cool season and warm season temperatures between 1916 and 2003 (Mote et al. 2005, Hamlet et al. 2007) (fig. 9). Although the rate of change varies with location and the time period examined, the warming has been on the order of 1ºC per century over the 1916 to 2003 time period (Hamlet et al. 2007). The rate of increase from 1947 to 2003 is roughly double that of the longer period from 1916 to 2003, and much of the observed warming has occurred from about 1975 to present.

Temperature increases in the west over the next century are expected to range from 2 to 3 ºC at the low end of the uncertainty range to 5 to 6 ºC at the upper end of the uncertainty range (IPCC 2007, Miles et al. 2007). Beyond mid-century, future warming is dependent on greenhouse gas emission levels in the next few decades, which are dependent on human activities.

There have been increases in winter (November-March) precipitation since 1930 over much of the Western United States, although patterns are variable in different regions (Mote et al. 2005) (fig. 9). Precipitation changes in the West over the next century are complex and more uncertain than temperature changes. Expected changes in precipitation patterns differ by region. Summer rains in the Southwest may intensify and shift to the North. Winter rains might decrease in the Southwest but increase in the northern half of the West (Salathé 2006). Interannual and interdecadal variability via El Niño-La Niña cycles may also intensify (Timmermann et al. 1999), producing extreme winter events in both the Southwest and the Northwest.

Figure 9—Linear trends in November-March (a), (b) temperature and (c), (d) total precipitation of the period indicated for the Western United States and Canada. For temperature, negative trends are indicated by blue circles, and positive trends are indicated by red circles; values are given in degrees Celsius per century. For precipitation, trends are given as a percentage of the starting value (1930 or 1950), and positive trends are shown as blue circles. Source: Mote et al. 2005.

The Hydrologic Cycle

In the Western United States, increased temperatures have led to more precipitation falling as rain rather than snow, earlier snowmelt and snowmelt-driven streamflow (Stewart et al. 2005, Hamlet et al. 2007), and reduced spring snowpack (Mote 2003, Mote et al. 2005, Barnett et al. 2008) (fig. 10). For the mountainous regions of the Western United States, snowmelt provides approximately 70 percent of annual streamflow (Mote et al. 2008). Both increased winter rain (as opposed to snow) and shifts to earlier spring snowmelt result in greater winter and spring streamflows and reduced summer streamflows in snowmelt-dominated and transient (rain/snow) watersheds (fig. 11). This reduction in summer streamflow could have major implications for fisheries, wildlife, water supply, and agriculture, particularly in drier regions. The current and expected future trends in hydrology suggest a coming crisis in water supply for the Western United States (Barnett et al. 2008).

Increased temperatures may also result in decreased soil moisture in arid regions of the Western United States (Miles et al. 2007). Changes in soil moisture are expected to differ by region. In the Pacific Northwest, it is expected that mountainous regions will have 80 percent or less of historical average soil moisture, while arid regions will have 90 to 95 percent of historical soil moisture (Miles et al. 2007).

Warmer temperatures and higher rates of evapotranspiration with climate change in some areas, such as the Southwest United States, will likely lead to increased drought frequency and severity. Overall, drought-affected areas are projected to increase in extent (IPCC 2007). Although increased temperatures will likely lead to decreased runoff in some areas, increased frequency of heavy precipitation events will likely lead to increased flood risk in many regions (IPCC 2007). Earlier snowmelt and runoff owing to increased temperatures could also lead to increased winter and spring flooding.

Figure 10—Changes in April 1 snow water equivalent in the Western United States. Linear trends in April 1 snow water equivalent (SWE) relative to 1950 at 798 snow course locations in the Western United States and Canada for the period 1950-1997. Negative trends are shown by red circles and positive by blue circles. SWE is a common measurement for the amount of water contained in snowpack if it were melted instantaneously. Sources: Mote et al. 2005, Casola et al. 2005.

Figure 11—Winter precipitation sensitivity and projected changes in monthly streamflow for the Yakima River basin in Washington State. Source: Casola et al. 2005.

Ecosystem Function and Process

Climate controls ecosystem structure and processes such as species distribution and abundance, regeneration, vegetation productivity and growth, and disturbance, including insects, and fire. Increasing temperatures and changes in precipitation with climate change will impact both ecosystem structure and ecosystem processes. This section highlights some of potential effects of climate change on vegetation, wildlife and ecosystem disturbance processes.

Vegetation—

Abundance and distribution of plant species shift individualistically in response to climate fluctuations. Plant species respond according to both thermal constraints and water constraints. For example, regeneration of tree species increases with changes in limiting factors, such as snowpack, length of growing season, and summer soil moisture levels. Thus, with increasing temperature, regeneration of species in high-snow environments will likely increase, while regeneration of species in drier, lower elevation environments will likely decrease.

Tree growth and productivity will also change with increasing temperatures. Lower snowpacks, and longer growing seasons may result in increasing growth and productivity in subalpine forests. However, forest productivity may decrease in lower elevation forests owing to water limitations.

With increased temperatures in the Western United States, the highest and coldest alpine (tundra) zones will likely contract significantly. The boreal and temperate forest zones (primarily conifer dominated) will likely shift up in elevation helping to squeeze the high-elevation zones into smaller domains. The frost-sensitive vegetation of the subtropical zone, including oaks and other woody and ephemeral species, will also likely expand up in elevation and north. This expansion of southern species could result in a contraction of the Great Basin shrublands.

Water constraints will have complex effects on vegetation distribution in the Western United States. Although precipitation may increase in some areas, increased evapotranspiration with increased temperatures may lead to increased water stress. However, higher concentrations of carbon dioxide in the atmosphere may reduce water stress. Changes in frequency and severity of fire may also influence vegetation distribution. Model simulations of western vegetation response to future climate change show a shifting of the water-limited boundaries, such as between closed forest and open tree-savanna, further down in elevation in the northern half of the West (north of the Oregon-California border) (Bachelet et al. 2001) (fig. 12). Other water-limited vegetation in these same regions, such as pine and juniper woodlands, is expected to expand (Bachelet et al. 2001) (fig. 12). In the Southwest, winter precipitation may decrease, but summer precipitation might increase. With the benefit of increased water use efficiency from elevated CO₂ concentrations, lower ecotones might shift down the mountains. At the lower elevations, the reduction in winter precipitation may limit woody vegetation. Increased summer precipitation would benefit the summer C4 grasslands.

Figure 12—Potential vegetation distribution simulated by the Mapped Atmosphere-Plant-Soil System (MAPSS) model and a dynamic model (MC1) for current conditions (historical for MAPSS and 1990 for MC1) and for future conditions (2070-99 for MAPSS and 2095 for MC1) under two future scenarios: HADCM2SUL and CGCM1.

Wildlife—

Viability of a species is dependent on the availability of suitable habitat. Animal species respond to climate variability in the short term through shifts in geographic range (migration) when suitable habitat is not available in the former range. Mortality and population extirpation in parts of a species’ former range often occur. Over time, extirpation and colonization events cumulatively result in shifts of the species’ distribution range (Davis and Shaw 2001, Delcourt and Delcourt 1991).

Species distributions have already changed in response to climate change in the Western United States. For example, the northern boundary of the sachem skipper butterfly (Atalopedes campestris) has moved from California to Washington State (420 miles) over a 35-year period (Crozier 2003, 2004) (fig. 13). Studies show that winter cold extremes determine the northern range limit (Crozier 2003, 2004).

Figure 13—Overwintering range of the sachem skipper butterfly (shaded) in Washington, Oregon, California, and Nevada from Opler (1999), modified to include the western range expansion (lighter shading). Colonization by the sachem skipper butterfly in four cities in Oregon and Washington show the chronology of range expansion. Contour lines represent the January average minimum -4 ºC isotherm 1950-1999 (solid) and 1990-1998 (dotted) (NCDC 2000). Source: Crozier 2003.

Changes in phenology, or timing of life history events, of both plant and animal species with climate change could influence wildlife. For example, in response to a 1.4 ºC rise in local temperatures at the Rocky Mountain Biological Laboratory in Colorado between 1975 and 1999, yellow-bellied marmots (Marmota flaviventris) emerged from hibernation 23 days earlier. However, the flowering plant phenology did not shift in that time period. Thus, the change in marmot behavior shifted the relative phenology of marmots and their food plants (Inouye et al. 2000). Shifts in prey behavior could similarly influence predator species.

Population extinctions have occurred in the Western United States in response to increasing temperatures over the last few decades. In the Great Basin, since being recorded in the 1930’s, 7 out of 25 recensused populations of the pika (Ochotona princeps) were extinct (Beever et al. 2003). There is little human disturbance in the high-elevation pika habitat. It was observed that extinct populations were those that had been at significantly lower elevations than populations still present (Parmesan and Galbraith 2004). Experiments show that adult pikas are sensitive to high temperatures (Smith 1974).

Land-use changes, urban development, and introduction of invasive species often impede the ability of species to respond to climate change adaptively. For instance, many land-use changes impose barriers to species’ migration to favorable new environments; small population sizes and isolation of populations as a result of landuse impede gene flow, and landscape fragmentation reduces corridors for movement (Joyce et al., in press).

Fire—

Widespread fire years and fire extent are associated with warmer and drier spring and summer conditions in the Western United States (McKenzie et al. 2004, Westerling et al. 2006, Heyerdahl et al. 2008, Taylor et al. 2008). Warmer spring and summer conditions lead to relatively early snowmelt, and lower summer soil and fuel moisture, and thus longer fire seasons (Westerling et al. 2006). Increased temperatures and drought occurrence in some locations owing to global warming will likely lead to increased fire frequency and extent. Intensity of fires may also increase in some areas if higher temperatures interact with fuel characteristics to increase fire intensity.

Insects—

Insect outbreaks may become more frequent and widespread because warmer temperatures may accelerate insect life cycles. Winter minimum and nighttime temperatures are forecasted to increase faster than maximum temperatures through the 21st century. This release of winter constraints is already occurring, increasing survival rates for insect larva and accelerating adult reproduction rates, thus leading to increased insect outbreaks. For example, the mountain pine beetle (Dendroctonus ponderosae) has invaded higher elevations and latitudes and significantly expanded its range in British Columbia owing to release of cold constraints (fig. 14). Thus, many forests that have historically never experienced these infestations are now being severely threatened and will continue to be threatened in the future.

In addition to effects of increased temperatures on insect life cycles, increased temperatures will also increase drought stress of some forest tree species, thus making some forests more susceptible to insect infestation. In addition, insect infestations can interact with fire. Recently burned forests may be more susceptible to insect damage. In turn, dead and weakened trees that have been infested with insects increase fire risk.

Figure 14—Mountain pine beetle damage in British Columbia. Photo taken by Lorraine Maclauchlan, Ministry of Forests, Southern Interior Forest Region

Figure 14—Mountain pine beetle damage in British Columbia. Photo taken by Lorraine Maclauchlan, Ministry of Forests, Southern Interior Forest Region (http://www.for.gov.bc.ca/hfp/mountain_pine_beetle/bbphotos.htm)

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