PART II - DRAFT 03.10.2010
20th Century Climatic Change and the Instrumental Record
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Throughout the western U.S., complex biophysical conditions (e.g., topography) are important factors
influencing highly variable responses to broad-scale climatic conditions. As a consequence, climatic
conditions vary across elevational, latitudinal and longitudinal gradients and at different spatial and
temporal scales. While trends in natural climatic variability are evident within the different climatic
regions it is important to keep in mind that broad-scale patterns in temperature, precipitation,
snowpack and other conditions also vary greatly across smaller spatial scales. Additionally, as a direct
measurement of climate conditions for the past century, the instrumental record provides some of the
best data of past climates yet the data from these records represents only the locales where
instruments are located and are often geographically biased. Thus, instrumental records provide direct
measurements of past conditions where they are located and often better represent certain elevations
and locals better than others (e.g., mid and low elevations vs. alpine environments).
Temperature
For all climate regions in this study, 20th century climate change is characterized by high spatial and
temporal variability. At broad temporal and spatial scales, however, it is possible to summarize trends
in climatic conditions impacting large portions of the four climate regions. Since 1900, temperatures
have increased in most areas of the western U.S. from 0.5-2˚ C (Fig. 19, Pederson et al. 2009, Mote et al.
2003, Ray et al. 2008) although there are a few exceptions where cooling has occurred (e.g.,
southeastern Colorado, Ray et al. 2008, and central Idaho and northwestern Montana, CIG). The rate of
change varies by location and elevation but is typically 1˚ C from early 20th century to present (Hamlet et
al. 2007). For most of the northern portions of the study area, temperatures generally increased from
1900 to 1940, declined from 1940-1975 and have increased from 1975 to present (USGCRP 2005).
Similarly, in the southern U.S. Rockies, temperatures generally increased in the 1930s and 1950s with a
period of cooling in the 1960s-70s and a consistent increase to present (Ray et al. 2008). Compared to
temperature increased in the early 20th century, temperatures for much of the study area doubled from
the mid 20th century to present, largely associated with rapid warming occurring since 1975.
Figure 19. 1950–2007 trend in observed annual average North American surface temperature (°C, left) and the
time series of the annual values of surface temperature averaged over the whole of North America (right). Annual
anomalies are with respect to a 1971–2000 reference. The smoothed curve (black line) highlights low frequency
variations. A change of 1°C equals 1.8°F. (Data source: UK Hadley Center’s CRUv3 global monthly gridded
temperatures). Source: Ray et al. 2008
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Temperature increases are more pronounced during the cool season (Hamlet and Lettenmaier 2007)
and in the northern U.S. Rockies, are roughly triple that for the global average (Hall and Fagre 2003), a
pattern that is evident at northern latitudes and higher elevation sites throughout the western U.S.
(USGCRP 2005). Mean regional spring and summer temperatures for 1987 to 2003 were 0.87˚ C higher
than those for 1970 to 1986, and were the warmest since 1895 (Westerling et al. 2006). Additionally,
Bonfils et al. (2008) and Barnett et al. (2008) found a strong anthropogenic signal in the warming
observed warming trends in mountain areas across the west, suggesting that a portion of recent
observed warming is attributable to human influenced changes in greenhouse gas and aerosol
concentrations.
Precipitation
Trends in precipitation for the study area are much less clear. Instrumental data for much of the
northwestern U.S. show modest increases in precipitation during the past century (Fig. 20, Mote et al.
1999, 2003,2005) whereas records from parts of the southern Rockies do not show trends in
precipitation for the past century (Ray et al. 2008). Natural variability in precipitation is evident in the
instrumental record for all of the climate regions and long-term drought conditions in the past century
impacted large areas of the study area although 20th century droughts were not as severe as those
evident at other periods during the past millennium (Cook et al. 2007, 2004). For example, two
significant droughts in the 1930’s and 1950’s impacted much of the study area. The 1930’s drought was
more widespread and pronounced in the northern climate regions while the 1950’s drought was
centered more on the southcentral and southwestern U.S. (Cook et al. 2007, Gray et al. 2004, Fye et al.
2003) and research suggests that climatic conditions that influenced the nature and location of these
droughts is likely linked to low-frequency oscillations in ocean-atmospheric interactions (McCabe et al.
2004, Gray et al. 2007, Hidalgo 2004, Graumlich et al. 2003).
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Figure 20. Trends in 1 Apr SWE over the 1960–2002
period of record directly from snow course observations.
Positive trends are shown in blue and negative in red, by
the scale indicated in the legend. Source: Mote et al.
2005 Journal of Climate
Surface Hydrology
Since 1950 more precipitation has been falling as
rain than snow, snowmelt and peak runoff is
occurring earlier and river flows are decreasing
during summer months (Fig. 21, Pederson et al.
2010, 2009, Mote et al. 2006, Barnett et al. 2008).
Recent impacts on snowpack and surface hydrology
are strongly associated with more precipitation
falling as rain than snow and earlier snowmelt
(Knowles et al. 2006). Comparisons of paleodata
and instrumental records with changing climate
conditions suggest that cool season precipitation is
more strongly associated with natural multidecadal, decadal and inter-annual variability in
ocean-atmosphere conditions than rising
temperatures.
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Despite modest increases in precipitation in parts of the central and northern portions of the study area
(modest declines for parts of the southern U.S. Rockies), significant declines in snowpack are evident in
much of the northwestern U.S., especially in the northern U.S. Rockies and parts of the Upper Columbia
Basin ((Hamlet and Lettenmaier 2007, Gray et al. 2004, 2007, Pederson et al. 2004, 2010, Selkowitz et al.
2002, Mote et al. 2005, Mote et al. 2008). In Glacier National Park glaciers decreased nearly 30% by
1993 from their areal extent in 1850 (Hall and Fagre 2003) and of the 150 glaciers present in 1910, only
27 still exist (Fagre 2002). Much of the changes to surface hydrology of the western U.S. since 1950 can
be attributed to human-caused climate change related to increases in greenhouse gas emissions and
aerosols (Barnett et al. 2008, Bonfils et al. 2008). During the past century trends in drought conditions
have been increasing in central and southern parts of the study area and decreasing in northern areas of
the study area (Andreadis and Lettenmaier 2006) although drought conditions are expected to increase
for much of the study area in the future (Hoerling and Eischeid 2007).
Figure 21. Average winter (Dec-Feb; top),
spring (Mar-May; middle), and annual
(bottom) minimum temperatures from
SNOTEL (water year Oct-Sep) and valley
MET (calendar year Jan-Dec) stations.
Source: Mote et al. 2006
Ocean-Atmosphere Interactions
Ocean-atmosphere interactions (e.g.,
Pacific Decadal Oscillation, Atlantic
Multidecadal Oscillation, and El Niño
Southern Oscillation) are important
drivers of multi-decadal, decadal and
inter-annual variability in temperature
and precipitation (Pederson et al. 2010,
2009, Gray et al. 2007, 2004, 2003,
McCabe et al. 2004, Hidalgo 2004) but
their impacts vary greatly across latitudinal, elevational and longitudinal gradients. The Pacific decadal
oscillation (PDO) exhibits a cool and warm phase that typically last for 20-30 year intervals (Mantua et
al. 1997). During the 20th century several switches occurred between warm and cool phases and the
magnitude of PDO phases increased in the latter half of the past century (Fig. 22, McCabe et al. 2004,
Mantua et al. 2002, 1997).
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Figure 22. Time series of the annual PDO and AMO.
Shaded areas indicate combinations of positive (+) and
negative (-) PDO and AMO periods. Source: McCabe et
al. 2004 PNAS
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Ocean-atmosphere interactions - PDO and ENSO.
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The Pacific Decadal Oscillation (PDO) and El Niño-Southern Oscillation (ENSO) are the predominant source of interdecadal (PDO) and inter-annual (ENSO) climate variability for much of the study area and the potential for
temperature and precipitation extremes increases when ENSO and PDO are in the same phases (CIG, other cits.).
Natural variations in PDO and ENSO are characterized by changes in sea surface temperature, sea level pressure,
and wind patterns (Mantua 1997, Wolter and Timlin 1993).
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The PDO is described as being in one of two phases: a warm phase (positive index value) and a cool phase
(negative index value). Figure 1 shows the sea surface temperature (SST) anomalies that are associated with the
warm phase of PDO. The spatial patterns are very similar: both favor anomalously warm sea surface temperatures
near the equator and along the coast of North America, and anomalously cool sea surface temperatures in the
central North Pacific. The cool phases for PDO and ENSO, which are not shown, have the opposite patterns of SST
anomalies: cool along the equator and the coast of North America and warm in the central north Pacific. During
the 20th century, each PDO phase typically lasted for 20-30 years (Figure X previous page). Studies indicate that
the PDO was in a cool phase from approximately 1890 to 1925 and 1945 to 1977. Warm phase PDO regimes
existed from 1925-1946 and from 1977 to (at least) 1998. Pacific climate changes in the late 1990's have, in many
respects, suggested another reversal in the PDO (from "warm" to "cool" phase and possibly back to “warm”).
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Figure 1. Warm Phase PDO and ENSO. The
spatial pattern of anomalies in sea surface
temperature (shading, degrees Celsius) and sea
level pressure (contours) associated with the
warm phase of PDO for the period 1900-1992.
Note that the main center of action for the
PDO (left) is in the north Pacific, while the main
center of action for ENSO is in the equatorial
Pacific (right). Contour interval is 1 millibar,
with additional contours drawn for +0.25 and
0.5 mb. Positive (negative) contours are
dashed (solid). Source: Climate Impacts Group,
University of Washington.
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ENSO variations are commonly referred to as El Niño (the warm phase of ENSO) or La Niña (the cool phase of
ENSO). An El Niño is characterized by stronger than average sea surface temperatures in the central and eastern
equatorial Pacific Ocean, reduced strength of the easterly trade winds in the Tropical Pacific, and an eastward shift
in the region of intense tropical rainfall (Fig. 2). A La Niña is characterized by the opposite – cooler than average
sea surface temperatures, stronger than normal easterly trade winds, and a westward shift in the region of intense
tropical rainfall.
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Natural variation in the strength of PDO and ENSO events impact climate regions in different ways. In the
northwestern U.S. and parts of the central U.S. Rockies, warm-phase PDO and El Niño winters tend to be warmer
and drier than average with below normal snowpack and streamflow whereas La Niña winters tend to be cooler
and wetter than average with above normal snowpack and streamflow (Graumlich et al. 2003, Cayan et al. 1999).
The southern U.S. Rockies and the southwestern U.S. respond differently, here warm-phase PDO and El Niño
winters tend to be wetter than average with above normal snowpack and streamflow and La Niña winters tend to
be drier than average with below normal snowpack and streamflow (Gray et al. 2004).
Figure 2. Multivariate ENSO index, 1950-2009.
Positive (red) index values indicate an El Niño event.
Negative (blue) values indicate a La Niña event
(Wolter and Timlin 1998, 1993).
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As with other indices of ocean-atmosphere conditions, the PDO index influences precipitation differently
across the western U.S., with cool-season precipitation negatively correlated with a positive (warm)
phase PDO in the Upper Columbia Basin and northern U.S. Rockies and parts of the central U.S. Rockies
and GYA and positively correlated with the warm PDO in parts of the central and southern U.S. Rockies
(Fig. 23, Mote et al. 2005). Thus, the northwestern U.S. generally receives more winter precipitation
during when PDO index is low or negative and the southwestern U.S. receives more precipitation when
the PDO index is high or positive.
Fig 23. Relationships between two climate
indices, NPI and PDO, and 1 Apr SWE, over the
1960–2002 period of record. (a), (b) Correlations
are shown as red for negative and blue for
positive; circles indicate statistically significant
trends, and/or indicates insignificant trends. (c),
(d) The trend explained by regression with the
index, SWEX [see Eq. (3)], in units of cm as in Fig.
5. Source: Mote et al. 2005 Journal of Climate
Significant droughts appear to be linked to
complex interactions between PDO, AMO
and, to a lesser extent, variations in ENSO
although longer droughts likely result from
low-frequency oscillations in PDO and AMO.
Research indicates a link between the warmphase (positive) of AMO and past drought
with positive and negative phases of PDO
moderating the geographic center of
droughts (centered more in northwestern
U.S. when PDO is positive and southwestern
U.S. when PDO is negative (Fye 2003, Hidalgo
2004). For example, the dust bowl drought
was associated with a positive AMO and a
positive PDO and was centered more over
the southwestern U.S. whereas the 1950’s drought (positive AMO and PDO) was centered more over
Wyoming, Montana and the Canadian Rockies (Gray et al. 2004, Fye et al. 2003). Drought conditions in
the interior western U.S. are strongly associated with low-frequency variations in AMO and to a lesser
extent PDO (McCabe et al. 2004, Hidalgo 2004, Graumlich et al. 2003, Enfield et al. 2001) and these
variations appear more pronounced in the northern and central U.S. Rockies than parts of the
southwestern U.S. (Hidalgo 2004). In the coastal Pacific Northwest and southwestern U.S., variations in
precipitation and warm-season water availability appear more sensitive to low-frequency ENSO
variations than PDO and AMO although different combinations of these phases tend to amplify or
dampen ENSO signals in climatic and hydrologic records (Gray et al. 2007, 2004, McCabe et al. 2004,
Hidalgo 2004). While ocean-atmospheric interactions including ENSO and PDO are partially responsible
for variations in climatic conditions across the climate regions that are the focus of this synthesis,
research suggests that since the late 20th century, human influences, via increased greenhouse gas and
aerosol concentrations, are amplifying, dampening and, in some cases, overriding the influence of
natural variability of these phenomena (Barnett et al. 2008, Bonfils et al. 2008, McCabe et al. 2008, Gray
et al. 2006, 2003).
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Changes in storm track and circulation patterns
Simulations of 21st century climate suggest a northward movement of the storm-track influencing
precipitation for much of the western U.S. (Yin 2005, Lorenz and DeWeaver 2007) which has the
potential to reduce precipitation for large parts of the study area. McAfee and Russell (2008) show that
a strengthening of the Northern Annual Mode, an index of sea level pressure poleward of 20˚ N, which
results in a poleward displacement of the Pacific Northwest stormtrack, increased zonal (west to east)
flow and reduced spring precipitation west of the Rocky Mountains and increased spring precipitation
east of the Rocky Mountains (McAfee and Russell 2008). This shift in the storm track is expected to
persist well into the future and may reduce the length of the cool-season, when circulation patterns
provide the bulk of precipitation for large areas of the central and northern parts of the study area
(McAfee and Russell 2008). If this becomes a more permanent shift in the storm-track position, this
phenomena could increase warm season (i.e., warm and dry) conditions for the Upper Columbia Basin,
northern U.S. Rockies and parts of the central U.S. Rockies whereas the central and southern U.S.
Rockies could receive increased spring season precipitation east of the Rockies. Changes in storm-track
position and circulation patterns will be superimposed on current trends, amplifying or dampening
changing conditions (e.g, reduced snowpack, earlier snowmelt and peak flows, diminished summer
flows, increased evapotranspiration) depending on location.
Ecological Impacts
Recent changes in climate conditions such as warming temperatures and associated declines in
snowpack and surface-hydrology are already influencing ecosystem dynamics. Several examples
illustrate the types of ecological impacts that have been observed in the past century including: earlier
spring bloom and leaf out of plant species, forest infill at and near treeline, and increased impact of
disturbances such as wildfire and insect outbreaks, all of which are likely to continue with increased
warming. Throughout much of the western U.S. spring bloom of a number of plant species has occurred
earlier, in some cases as much as several weeks (Cayan et al. 2001, Schwartz and Reiter 2000). In the
northern U.S. Rockies, increased density of trees at or near treeline has been observed at a number of
sites (Fagre et al. 2004). This “infill” phenomenon is not uncommon in the western U.S. and is predicted
to continue where minimum temperatures rise, snowpack in high-snowfall areas decreases and
moisture is not limiting (Graumlich et al. 2005, Lloyd and Graumlich 1997, Fagre et al. 2004, Rochefort et
al. 1994). While evidence for “infill” is widespread, upslope movement in treeline position is much more
variable and research suggests that upslope movement will be characterized by a high degree of spatial
heterogeneity in relation to variables that control treeline (Graumlich et al. 2005, Lloyd and Graumlich
1997).
Changing climate conditions are also influencing important disturbance processes that regulate
ecosystem dynamics. Warming temperatures, earlier snowmelt and increased evapotranspiration are
increasing moisture stress on forest species making them more susceptible to insect invasions. An
increase in the extent, intensity and synchronicity of mountain pine beetle attacks in the western U.S.
and Canada have been linked to forests stressed by dry intervals and are less able to resist beetle
infestations (Bentz et al. 2009, Hicke et al. 2006, Romme et al. 2006, Logan et al. 2003, Carroll et al.
2004, Breshears et al. 2005). Warming temperatures have also influenced bark beetle population
dynamics though reduced winter kill and have helped facilitate the reproduction and spread of these
insects (Black et al. 2010, Carroll et al. 2004). In some cases, past forest management (e.g., factors
related to structural characteristics of host stands) might also facilitate beetle infestation (Black et al.
2010, Bentz et al. 2009). Another example involves the area of the western U.S. burned by forest fires
annually. The extent of the western U.S. burned by fires each year is strongly linked to changing climate
conditions (Littell et al. 2009, 2008, Higuera et al. 2009, Morgan et al. 2008) and changes in surface
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hydrology associated with reduced snowpack, earlier spring runoff and peak flows, and diminished
summer flows have been linked to increased frequency and duration of large fires and a lengthened fire
season in the western U.S., impacts where are most evident at mid-elevation forests in the northern U.S.
Rockies (Westerling et al. 2006). These are only a few examples of ecological impacts linked to recent
changes in climate, and all of these impacts are expected to become more pronounced in many parts of
the study area with future changes.
Northern U.S. Rockies
Temperature
Over the course of the 20th century, the instrumental record in the northern Rockies (NR) shows a
significant increase in average seasonal, annual, minimum and maximum temperatures (Figs. 24-25,
Loehman and Anderson 2009, Pederson et al. 2010, 2009). Regional average annual temperatures in
the Northern Rockies increased between 1-2C (Pederson et al. 2009, Pederson et al. submitted,
Loehman and Anderson 2009). Within this framework of increasing regional temperatures, seasonal
and annual minimum temperatures are generally increasing at a rate greater than that of the maximum
temperatures (Pederson et al. 2009, Pederson et al. submitted). In particular summer and winter
seasonal average minimum temperatures are increasing at a rate significantly greater than that of the
respective season’s average maximum temperatures, causing a pronounced reduction in the seasonal
daily temperature range (DTR) index (Pederson et al. 2009). The magnitude of minimum temperature
increases also appears seasonally variable: in area SNOTEL sites, Pederson et al. (submitted) estimated
minimum temperature increases since 1983 of +3.8 ± 1.72˚ C in winter (DJF), of +2.5 ± 1.23˚ C in spring
(MAM), and of +3.5 ± 0.73˚ C annually (Fig. 25). The magnitude of changes varies locally but there are
few exceptions to this general trend in warming.
Fig. 24 Comparison of variability and trends in western Montana (blue-green) and Northern Hemisphere (dark blue
line) annual average temperatures. A 5-year moving average (red line) highlights the similarity in trends and
decadal variability between records. Source: Pederson et al. 2010 Climatic Change
Temperature trends within the NR generally track Northern Hemisphere trends across temporal scales
(Fig. 24). This similarity between regional and continental scale trends suggests that large-scale climate
forcings such as greenhouse gases, patterns in sea surface temperatures (SSTs), volcanic activity, and
solar variability that drive decadal scale temperature global and continentally also drive regional
temperatures in the NR (Pederson et al. 2009).
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Figure 25. SNOTEL station Tmin records have
been fit using a non-linear quadratic equation
due to characteristics of these time series. All
trends shown are significant (p ≤ 0.05) and
note the y-axis temperature scale changes for
each panel. Source: Pederson et al. submitted
Precipitation
When compared to the distinctive,
statistically significant trends present in
the 20th century temperature records for
the NR, no long-term (centennial scale)
trends are evident in the precipitation
time series data. Throughout the west,
high inter-annual, annual, and decadal
precipitation variability hinders trend
detection within the time series to derive
a consistent, centennial-scale trends in
precipitation that are statistically
significant (Ray et al. 2008). General
patterns throughout the latter part of the
20th century indicate that areas within the
NR experienced noticeable, albeit modest,
decreases in annual precipitation (Mote et al. 2005, Knowles et al. 2006). Although few statistically
significant long-term trends can be derived from regional 20th century precipitation time series data,
rising temperatures throughout the west have led to an increasing proportion of precipitation falling as
rain rather than snow (Knowles et al. 2006). However, due to regional winter temperatures averaging
significantly less than 0C, areas in the NR are generally less sensitive to shifts in precipitation type
spurred by rising temperature when compared to other regions in West (Knowles et al. 2006).
Surface Hydrology
In the NR, like most of the western United States, snow water equivalent (SWE) within a winter
snowpack largely controls surface runoff and hydrology for the water year and consequently has
significant impacts on water resources throughout the year (e.g. Pierce et al. 2009, Barnett et al. 2008,
Stewart et al. 2005, Pederson et al. submitted). Over the second half of the twentieth century, studies
have demonstrated a statistically significant decrease in winter snowpack SWE across the region
(Barnett et al. 2008, Pederson et al. submitted). Additionally, earlier onsets of springtime snowmelt and
streamflow have been documented (Stewart et al. 2005).
Ocean-atmosphere interactions
In the NR during the 20th century, the warm-phase (positive) of PDO is associated with reduced
streamflow and snowpack (Fagre et al. 2003) and the cool-phase of PDO is associated with increased
streamflow and snowpack (Fig. 26-27). Ecological responses are also evident in changes in the
distribution of mountain hemlock (Tsuga mertensiana). At high elevations growth of mountain hemlock
is limited by snowpack free days where a warm-phase PDO often results in decreased snowpack and
increased mountain hemlock growth (Fagre et al. 2003). The response is opposite at low elevation sites
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where moisture is limiting and a warm-phase PDO commonly leads to less moisture which constrains
mountain hemlock growth and establishment (Fagre et al. 2003).
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Figure 26. First EOF amplitudes representing spatially
broad patterns of (top) variability in winter wet-day
minimum temperature and (bottom) fraction of winter
precipitation falling as snow, with PDO phases indicated.
The thick, wavy lines are low-passfiltered versions of
each amplitude. The portion of the low-pass curve that
best describes a given period is used for that period.
Source:
Figure 27. Comparison of GNP summer drought, winter snowpack (May 1
SWE) and the Pacific Decadal Oscillation. All time series have been smoothed
using a 5 yr moving average. (a) Relationship between the average annual
instrumental PDO index (blue line, inverted for ease of comparison) [Mantua
and Hare, 2002] and May 1 SWE anomalies (red line) for GNP. (inset)
Correlations between winter (October–March) precipitation and the PDO index
for all U.S. climate divisions spanning 1949–2003
(http://www.cdc.noaa.gov/USclimate/Correlation/). (b) Reconstructed PDO
[D’Arrigo et al., 2001] used as a proxy for snowpack anomalies back to 1700.
Positive (negative) values of the reconstructed PDO correspond with low (high)
winter snowpack as indicated by the strong relationship with instrumental May
1 SWE anomalies (r = _0.688) for the common period of overlap (1922–1979).
(c) Summer drought (MSD) reconstruction for GNP. Source: Pederson et al.
2004 Geophysical Research Letters
Pederson et al. (submitted) summarize how variation in Pacific SSTs,
atmospheric circulation and surface feedbacks influence climate
conditions, snowpack and streamflow for the northern Rocky
mountains. Winters with high snowpack in the NRMs tend to be associated with negative PDO
conditions, a weakened Aleutian Low, and low pressure centered poleward of 45°N across western
North America (Fig. 28). During years of high snowpack, for example, the tendency is for mid-latitude
cyclones to track from the Gulf of Alaska southeast through the Pacific Northwest and into the NRMs.
The relatively persistent low-pressure anomaly centered over western North America is also conducive
to more frequent Arctic-air outbreaks resulting in colder winter temperatures. ENSO is also an
important driver of snowpack and streamflow at inter-annual time-scales, and the influence of related
tropical Pacific atmospheric circulation anomalies persists well into the spring.
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Changes in spring (MAM) temperatures and precipitation are associated with changes in regional
atmospheric circulation, and are also shown to strongly influence the timing of NRM streamflow (Fig.
28). Springtime geopotential heights over western North America influence the amount, but more
importantly the timing, of snowmelt and streamflow across the northern U.S. Rockies. Specifically, high
pressure anomalies centered over western North America correspond with increased spring
temperatures and consequently the increasing number of snow-free days, early arrival of snow meltout, and streamflow CT. Atmospheric circulation changes occurring in March and April can, in turn,
initiate surface feedbacks that contribute to surface temperature and hydrograph anomalies (Fig. 28).
Hence, in the northern U.S. Rockies warming temperatures influence earlier snowmelt and runoff and
associated decreasing snowpack and streamflow but can also be partially attributed to seasonallydependant ocean-atmosphere teleconnections and atmospheric circulation patterns as well as surfacealbedo feedbacks that interact with broad-scale controls on snowpack and runoff (Pederson et al.
submitted).
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Figure 28. Idealized relationship
between NRM snowpack and
streamflow anomalies with associated
Pacific SSTs, atmospheric circulation,
and surface feedbacks. Source:
Pederson et al. submitted
Central U.S. Rockies and GYA
Temperature
Temperatures for the central U.S. Rockies and the GYA (CR-GYA) have increased 1-2C during the past
century, with the greatest increases occurring in the latter half of the 20th century. Trends in
temperatures during the past century are slightly higher than the southwestern U.S. and slightly lower
than the northern U.S. Rockies, following a pattern of more pronounced temperature increases for
higher latitudes in the latter half of the 20th century (Cayan 2001, Ray et al. 2008). Increasing winter and
spring temperatures have resulted in reduced snowpack, earlier spring snowmelt and peak flows, and, in
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some cases, lower summer flows for major basins in the climate region (Watson et al. 2009, Gray et al.
2007, 2004, 2003, Graumlich et al. 2003)
Precipitation
Records of precipitation for the CR-GYA show highly variable patterns across gradients in elevation,
latitude and longitude. No long-term trends over the past century are evident although patterns in
inter-annual, decadal and multi-decadal variation are evident in reconstructions of past hydrology from
tree-ring records (Fig. 29, Watson et al. 2009, Gray et al. 2007, 2004, Graumlich et al. 2003). A greater
proportion of precipitation is falling as rain versus snow in this climatic region but the impacts are less
pronounced than other parts of the western U.S. (Knowles et al. 2006). In many parts of the CR-GYA,
the 1930s and 1950s were significantly drier and the 1940s were wetter than average although subregional variation is high, likely associated with the location of this climate region in a transition
between Pacific Northwest and southwestern U.S. atmospheric circulation patterns discussed further
below (Watson et al. 2009, Gray et al. 2007, 2004, 2003, Graumlich 2003).
Figure 29. (a) Observed annual (previous July through current year June) precipitation for Wyoming Climate
Division 1 (gray line) vs. estimates of precipitation based on the stepwise regression (black line) model. (b) The full
stepwise version of reconstructed annual precipitation (black line) for AD 1173 to 1998. The horizontal line (solid
gray) near 400 mm represents the series mean, and the vertical line (dotted gray) at AD 1258 divides the wellreplicated portion of the record from reconstructed values in the earlier, less-replicated years (see Methods).
Source: Gray et al. 2007 Quaternary Research
Ocean-atmosphere interactions
The influence of ocean-atmospheric interactions on decadal, multi-decadal and inter-annual variation in
climatic conditions in the CR-GYA is more spatially variable than the other climate regions because the
region falls in a transition area between northwestern U.S. and southwestern U.S. circulation patterns
that strongly influence climatic conditions differently for these two areas of the western U.S. (Gray et al.
2007, 2004, Graumlich et al. 2003). Because the CR-GYA encompasses this transition between major
circulation types, a complex interaction between natural variations in ocean-atmospheric interactions
such as El Niño/Southern Oscillation (ENSO) and Pacific Decadal Oscillation (PDO) and topography,
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latitude and longitude often result in opposite trends in climatic conditions at sites within the same
region (Gray et al. 2007).
Variation in precipitation and water availability in the CR-GYA is complex, largely due to the fact that CRGYA is influenced by both Southwest and Pacific Northwest modes of ocean-atmospheric variability and
is split between the winter-wet and summer-wet divisions outlined by Whitlock and Bartlein (1993).
High elevation snow basins within the central Rockies and western portions of the GYA typically
responds similar to the Pacific Northwest where the cool-phase (negative) PDO results in cool, wetter
than average winters and warmer and drier than average winter precipitation coincides with the warmphase (positive) PDO (Gray et al. 2007, 2004, Graumlich et al. 2003, Dettinger et al. 1998). Similar to the
Pacific Northwest, these portions of the climate region experience increased precipitation during La
Nina ENSO events and decreased precipitation during El Niño events, and ENSO seems to be linked to
the magnitude of PDO anomalies, especially so during winter months (Gray et al. 2007). Alternatively,
lower elevation sites and eastern portions of the GYA respond more like areas of the southwestern U.S.
or show little response to ENSO events (Gray et al. 2004). Here, years with strong El Nino SST’s result in
increased winter precipitation and drier conditions during La Nina events. This decoupling of high and
low and low elevation precipitation regimes is common throughout the central U.S. Rockies,
complicating predictions of future precipitation for the region (Woodhouse 2001).
Southern U.S. Rockies
Temperature
In the last 30 years temperatures have increased between 1-2 ˚F throughout the Southern U.S. Rockies.
The north central mountains of Colorado warmed the most, ~2.5 ˚F and high elevations have warmed
more quickly than lower elevations (Diaz and Eischeid 2007). Warming is evident at almost all locations
and temperatures have increased the most in the north central mountains and the least in the San Juan
Mountains of southwestern Colorado (Ray et al. 2008). Only the Arkansas River Valley in southeastern
Colorado show slight cooling trend during the 20th century and no trend is evident in this area for the
latter half of the century (Ray et al. 2008).
Precipitation
Records of precipitation for the Southern U.S. Rockies (SR) for the past century indicate highly variable
annual amounts and no long term trends are evident for the region (Ray et al. 2008, Dettinger et al.
2005). Like elsewhere in the interior west, a greater proportion of precipitation is falling as rain versus
snow than in the past but these changes are less pronounced than in the Northern U.S. Rockies
(Knowles et al. 2006). Decadal variability is evident in records of precipitation and surface flows and is
linked to variability in ocean-atmosphere and land-surface interactions (Fig. 30, Dettinger et al. 2005).
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Fig 30. Observed time series (18952007) of annually averaged
precipitation departures areaaveraged over the Upper Colorado
drainage basin (top panel) and annual
Colorado River natural flow
departures at Lees Ferry in million
acre-feet (bottom panel). The
precipitation data are based on 4kmgridded PRISM data. Colorado River
natural flow data from the Bureau of
Reclamation. Source: Ray et al. 2008
Surface Hydrology
Paralleling trends evident throughout the interior west, more precipitation is falling as rain than snow,
spring snowpack is decreasing, especially at lower and mid-elevations below 2500 m and peak
streamflows are occurring earlier because of warmer spring temperatures (Knowles et al. 2006, Bales et
al. 2006, Stewart et al. 2005, Hamlet et al. 2005, Clow 2007, Mote 2006, 2003). Additionally, summer
flows are typically lower although annual flows show high variability but no significant trends in most
locations (Ray et al. 2008).
Ocean-atmosphere interactions
The Colorado River Basin spans an important transition area where the influence of Pacific Northwest
and southwestern U.S. circulation patterns show opposite trends (Gray et al. 2007, Clark et al. 2001).
Averages in SWE and annual runoff during El Nino years reflect this transition as northern parts of this
climate region experience drier-than-average conditions, whereas the southern portions experience
wetter-than-average conditions and the opposite conditions occur during La Nina years (i.e., wetter than
average in the north and drier than average in the south and west) and anomalies tend to be more
pronounced in spring in southern portions of the climate region. Long-term droughts are linked more
closely to low-frequency oscillations in PDO and AMO and are most commonly associated with the
interaction between a cool-phase (negative) PDO and warm-phase (positive) AMO (McCabe et al. 2004).
Upper Columbia Basin
Temperature
For most of the Upper Columbia Basin (UCB), average annual 20th century temperatures (data from
1920-2003) have increased by 0.7-0.8 °C and the warmest decade was the 1990s (Fig. 31, CIG 2010).
Average temperatures have increased as much as 2° C in parts of the climate region and increases have
been more pronounced at higher elevations (cits.). Winter and spring average temperatures and daily
minimum temperatures have increased more than other seasons and more than maximum
temperatures during the mid-20th century. During the latter half of the 20th century minimum and
maximum temperatures have increased at similar rates (Watson et al. 2009, Karl et al. 1993, Dettinger
et al. 1995, Lettenmaier et al. 1994, CIG 2010).
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Figure 31. 20th century trends in (a) average annual PNW
temperature (1920-2000). This figure shows widespread
increases in average annual temperature for the period
1920 to 2000. The size of the dot corresponds to the
magnitude of the change. Pluses and minuses indicate
increases or decreases, respectively, that are less than the
given scale. Source: Climate Impacts Group, University of
Washington.
Precipitation
Trends in precipitation for the Upper Columbia Basin are less clear than trends in temperatures and
observations indicate high decadal variability. Precipitation has generally increased in the northwestern
U.S., by 14% for the entire region (1930-1995) and range between 13%-38% for different parts of the
region (Fig. 32, Mote et al. 2003) although trends are often not statistically significant depending on the
area and time interval measured. Paralleling trends observed for much of the interior western U.S.,
variability in winter precipitation has increased since 1973 (Hamlet and Lettenmaier 2007).
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Figure 32. 20th century trends in average annual
precipitation (1920-2000). Increases (decreases) are
indicated with blue (red) dots. The size of the dot
corresponds to the magnitude of change. Source: CIG,
University of Washington.
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Surface hydrology
Spring snowpack and snow water equivalent (SWE) declined throughout the UCB in the latter half of the
20th century and was most pronounced at low and mid-elevations (Figure 33, Mote et al. 2003, Hamlet
et al. 2005, Mote 2006). Declines in snowpack and SWE are associated with increased temperatures and
declines in precipitation over the same time period and declines of up to 40% or more are recorded for
some parts of the climate region (Mote et al, 2003, 2005). In addition to declines in snowpack, the
timing of peak runoff has shifted 2-3 weeks earlier for much for much of the region during the latter half
of the 20th century (Stewart et al. 2004) and the greatest shifts have occurred in the mountain plateaus
of Washington, Oregon and western Idaho (Hamlet et al. 2007). Because ecosystems in the Upper
Columbia Basin rely on the release of moisture from snowpack these shifts are significantly impacting
plant species which are blooming and leafing out earlier in the spring (Mote et al. 2005, Cayan et al.
2001, Schwartz and Reiter 2000).
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Figure 33. 20th century trends in (b) April 1 snow
water equivalent (1950-2000). This figure shows
widespread increases in April 1 snow water equivalent
(an important indicator for forecasting summer water
supplies) for the period 1950 to 2000. The size of the
dot corresponds to the magnitude of the change.
Pluses and minuses indicate increases or decreases,
respectively, that are less than the given scale. Source:
CIG, University of Washington.
Ocean-atmosphere interactions
Variations in climatic conditions in the UCB are related to ocean-atmosphere and land-surface
interactions, namely El Niño/Southern Oscillation (ENSO) and Pacific Decadal Oscillation (PDO)
phenomena. In their warm phases (i.e., El Niño conditions for ENSO), both ENSO and PDO increase the
chance for a warmer winter and spring in the UCB and decrease the chance that winter precipitation will
meet historical averages. The opposite tendencies are true for cool phase ENSO (La Niña) and PDO: they
increase the odds that UCB winters will be cooler and wetter than average (Clark et al. 2001). While
typically warmer than average, SWE anomalies during strong El Niño are often less pronounced and
winter precipitations are commonly close to historical averages (Clark et al. 2001, CIG). Clark et al.
(2001) suggest that El Nino circulation anomalies are centered more in the interior west than for La Nina
circulation anomalies and are most evident in the middle of winter.
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Lessons from the recent past: what can we learn from 20th century observations?
Small changes can have large impacts
Modest increases in temperature are already having dramatic impacts on surface hydrology through
earlier and diminished spring snowpack (Mote 2006, 2003, Stewart et al. 2005, 2004, Pederson et al.
submitted), increases in the proportion of winter precipitation as rain versus snow (Knowles et al. 2006,
Bales et al. 2006), decreased snow season length at most elevations (Bales et al. 2006), earlier blooming
dates for plants (Cayan et al. 2001, Schwartz and Reiter 2000) and lower summer flows (Meko 2007).
Changes in the distribution of minimum temperatures and frost-free days illustrate how small changes
in temperature may have large changes to surface hydrology (Barnett et al. 2004, Stewart et al. 2004).
Evidence from a number of studies suggests that even small increases in temperatures are expected to
have dramatic impacts on water availability for much of the western U.S. Along with changes in
snowpack and earlier spring runoff, modest temperature increases predicted for the future suggest that
drought conditions will become much more common (Hoerling and Eischeid 2007, Seager et al. 2007,
Barnett and Pierce 2009). Increased winter precipitation that is predicted for central and northern
areas of the study area will likely not be adequate to offset increased rates of evapotranspiration,
leading to increased number of dry days and increased frequency (and possibly intensity) of droughts for
much of the interior west (Gray et al. 2009, Hoerling and Eischeid 2007, Seager et al. 2006).
Shifting distributions and new norms
The western U.S. is vulnerable to small changes in temperature and drier summers because much of the
western U.S. consists of arid to semi-arid ecosystems that depend on limited water resources regulated
by dynamics of mountain snowpack (Gray et al. 2009). While many ecosystems are adapted to natural
variations in water availability, a shift in the distribution of drought like conditions where they become
the norm and the frequency and magnitude of dry conditions exceed historical patterns could result in a
tipping point where major redistribution of vegetative communities ensues (Fig. 34, Jackson et al. 2009).
Rapid changes in the storage and distribution of surface flows related to snowpack could have significant
impacts on ecosystems adapted to past snowpack dynamics that may change to the point that they
exceed recent natural ranges of variation and ecosystem thresholds where adaptation and/or
fundamental transitions occur (Fig. 34).
Figure 34. Idealised version of a
coping range showing the
relationship between climate
change and threshold
exceedance, and how adaptation
can establish a new critical
threshold, reducing vulnerability
to climate change (modified from
Jones and Mearns 2005). Source:
IPCC 2007.
Small increases in temperatures are already resulting in greater evaporative losses from lakes, streams,
wetlands and from terrestrial ecosystems and modest increases in winter precipitation that some
models predict for the northern parts of the study area are not expected to keep pace with and offset
increased evapotranspiration (Gray et al. 2009, Arnell 1999). For example, models from Gray and
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McCabe (in review) estimate a 15-25% decrease in water in the Yellowstone River in the future and,
even with increased precipitation it would take the wettest years of last 800 years to offset increased
evapotranspiration associated with warmer temperatures (Gray and McCabe in review). Additionally,
the frequency of extreme episodes of precipitation and temperature has occurred in the past century
(Karl et al. 2009, IPCC 2007, Groisman et al. 2005, Kunkel et al. 2003, Madsen and Figdor 2007) and
shifts in the distribution of climate conditions (e.g., average summer temperatures) will likely result in
increased occurrence of extreme conditions when compared with 20th century conditions (Fig. 35).
Research indicates that delivery of precipitation in many parts of the western U.S. will occur as extreme
events and will result in more dry days as there will be longer intervals without precipitation (Groisman
et al. 2005, Kunkel et al. 2003, Madsen and Figdor 2007).
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Figure 35. Schematic showing the effect
on extreme temperatures when the mean
temperature increases, for a normal
temperature distribution. Source: IPCC
2007
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What can we expect in the future?
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Many of the trends in climate evident in the past century are expected to continue in the future.
Temperatures are expected to increase across the landscape, although short-term variation is still
expected. An example of short-term variation that some point to as a reversal of warming temperatures
relates to a modest cooling that much of the Northern Hemisphere experienced in the last decade, a
phenomenon which has now been largely attributed to a temporary decrease in atmospheric
concentrations of water vapor in the lower stratosphere (Solomon et al. 2010). This short-term cooling
illustrates of how short-term variations will be superimposed on longer-term trends. Changes in the
amount and spatial distribution of precipitation are related to a complex interaction of global circulation
patterns, ocean- and land surface -atmosphere interactions that are still poorly understood.
Additionally, the influence of human activities on natural variations appears to be increasing (Barnett et
al. 2008, Bonfils et al. 2008). While models generally predict increased precipitation for parts of the
northern and central U.S. Rockies and the Upper Columbia Basin and decreased amounts for the
southwestern U.S., both uncertainty and variability is expected to be high and is reflected in model
predictions. Levels of uncertainty are a critical component of any prediction of future conditions and
careful consideration of how levels vary should inform how managers anticipate future conditions.
Scenario planning can be an effective approach for considering future conditions that area highly
uncertain, an approach discussed in more detail later.
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Model projections of future conditions
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Analysis of future projections for temperature, precipitation, snowpack, stream flow, drought, growing
season. Jeremy Littell will provide figures representing wall-to-wall downscaled models (VIC model
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forecasts) for key variables (temperature, precipitation, soil moisture, snowpack, growing season and
possibly others e.g. extreme events?) for most of our study area. He will also provide some text to
accompany these figures and, if possible, text discussing key sources of uncertainty associated with
future projections.
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Figures of future projections for a few key variables (temperature, precipitation, soil moisture, snowpack
or streamflow, growing season) for each of the climate regions and text on uncertainty.
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1. Northern U.S. Rockies
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2. Central U.S. Rockies and the GYA
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3. Southern U.S. Rockies
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4. Upper Columbia Basin
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Ecological Response: Area burned example
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Phil Higuera is interested in providing predictions for future area burned under different climate
scenarios. This is new modeling that he has been working on and would provide an example of
projected ecological (process) response to biophysical changes identified in the VIC and other models.
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Trends that will likely continue to impact large parts of the study area
Climate conditions
 2-6 + C degrees in temperature and higher at higher latitudes
 Increased but highly variable precipitation for parts of the Upper Columbia Basin, northern and
central U.S. Rockies.
 Decreased but highly variable precipitation for parts of the central and southern U.S. Rockies.
 Increased evapotranspiration for most of the western U.S. which will likely not be offset by
increased precipitation (Fig. 36, Hoerling and Eischeid 2007, Seager et al. 2006).
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Fig. 36. Modeled changes
in annual mean
precipitation minus
evaporation over the
American Southwest
(125°W to 95°W and 25°N
to 40°N, land areas only),
averaged over ensemble
members for each of the
19 models. The historical
period used known and
estimated climate
forcings, and the
projections used the SResA1B emissions scenario. The median (red line) and 25th and 75th percentiles (pink
shading) of the P − E distribution among the 19 models are shown, as are the ensemble medians of P (blue
line) and E (green line) for the period common to all models (1900–2098). Anomalies (Anom) for each model
are relative to that model's climatology from 1950–2000. Results have been 6-year low-pass Butterworthfiltered to emphasize low-frequency variability that is of most consequence for water resources. The model
ensemble mean P − E in this region is around 0.3 mm/day. Source: Seager et al. 2007 Science
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Surface Hydrology
 Greater proportion of winter precipitation falling as rain than snow (Knowles et al. 2006, Bales
et al. 2006, Dettinger et al. 2004)
 Decreased snow season length at most elevations Bales et al. 2006
 Less spring snowpack (Pederson et al. submitted, Mote 2006, 2005, 2003)
 Earlier snowmelt runoff and peak streamflows (Stewart et al. 2005, 2004, Hamlet et al. 2005,
Clow 2007).
 Increased frequency of droughts and low summer flows (Gray et al. 2009, Meko et al. 2007).
Increased evapotranspiration that will likely not be offset even where precipitation increases
amplifying dry conditions (Hoerling and Eischeid 2007, Seager et al. 2006).
 Underground recharge may decline with declines in snowpack (Winnograd et al. 1998)
 With poles getting wetter and subtropics getting drier variability in mid-latitude precipitation is
expected to increase (Dettinger et al. 2004)
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Extreme conditions: Drought, Floods, Heat Waves
 More episodes of extreme temperatures (US Global Change report, Karl et al. 2009)
 Increased frequency of extreme precipitation (storm) events, rain on snow and consequent
winter/spring floods in mountains (Madsen and Figdor 2007, Groisman et al. 2005, Kunkel et al.
2003).
 Droughts are expected to become more frequent as a result of increased temperatures,
evapotranspiration and changes to surface hydrology that impact warm season conditions (Gray
et al. 2009, Meko 2007).
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Circulation patterns
 A long-term trend in the northward shift in the winter storm track and jet stream will likely
result in more winter/autumn precipitation for the northwestern U.S. and less for southern
Rockies and Southwest U.S. although variability is expected to by high (McAfee and Russell
2009).
 Natural variation in ocean-atmosphere interactions (e.g., PDO and ENSO) will continue to
influence climate in the western U.S. but human drivers of change (i.e., greenhouse gases and
aerosol concentrations) are now interacting with these natural variations and are dampening,
amplifying and, in some cases, overriding these natural drivers of change (Meehl et al. 2009,
Barnett et al. 2008, Bonfils et al. 2008).
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Productivity – Phenology
 Earlier blooming dates for many plant species (Cayan et al. 2001, Schwartz and Reiter 2000)
 Longer growing season (Bates et al. 2006)
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Disturbance
 Increased impact of disturbances linked to drought stress. Examples include, increased large
fires resulting from increased temperatures, changes in surface hydrology and snowpack and
conditions during the warm season (Westerling et al. 2006)
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
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Higher frequency of large-fires, longer fire season and increased area of western U.S. burned by
fire (Westerling et al. 2006, Littell et al. 2009, 2008, Higuera et al. 2009, Morgan et al. 2008)
Greater drought stress will likely result in more insect infestations and disease affecting forests
(Black et al. 2010, Bentz et al. 2009, Hicke et al. 2006, Romme et al. 2006, Logan et al. 2003,
Carroll et al. 2004, Breshears et al. 2005).
Planning for the future
I have made an initial attempt to adapt a part of Jackson et al. 2009 for this section but it needs further modification (Perhaps
Steve could look more closely at this).
Planning for future conditions that are highly uncertain presents a significant challenge for land
managers. Scenario planning provides an approach for preparing for future climate conditions that are
highly uncertain by anticipating a range of future conditions. Scenario planning uses a combination of
scientific input, expert opinion and forecast data to develop alternative scenarios for the future
(Schwartz 1991, van der Heijden 1996), contrasting with more traditional attempts at developing
precise, quantitative assessments of future conditions, which are often useless because of compounded
uncertainties. In scenario planning, alternative scenarios can be used as a starting point for exploring
species or ecosystem vulnerabilities under a range of future conditions, and as a means for examining
how management strategies might play out in the face of multiple drivers of change.
Jackson et al. (2009) developed an example to illustrate this process. In this example, alternative futures
can be arrayed along two axes comprising integrators of potential climate-change (drought frequency)
and potential changes in disturbance regimes (fire size). In concert with monitoring and modeling,
studies of past climates can define the range of drought frequency we might reasonably expect, and
past studies of fire can place bounds on potential fire size. This exercise yields four quadrants, each
comprising a distinct combination of climatic and fire-regime change (Fig. 37). These quadrants each
provide a contrasting scenario that can be used to explore potential impacts on species or ecosystems
and for examining the relative costs and benefits of various mitigation and adaptation measures.
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Figure 37. Example of a scenario planning matrix.
Each axis represents a critical driver of system
change or a significant trend in the environment.
In common practice, the variables chosen for
analysis are likely to have the strongest influence
on the system or they are associated with a high
degree of uncertainty (Shoemaker 1995). In the
case presented here, the axes represent a
continuum between conditions that are similar to
those observed in the historical record and
conditions that are significantly altered from those
seen today. Combining these two drivers produces
four alternative scenarios for the future conditions
(e.g., frequent drought and large fires in the upper
right) that can then be further developed into
“storylines” that provide details about how each
scenario might unfold. Depending on the
application and available data, axes and the
resulting storylines may be defined quantitatively,
or they may be based on qualitative assessments alone. Source: Jackson et al. 2009 Paleontological Society Papers
20
908
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911
912
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916
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921
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930
931
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933
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937
938
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940
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951
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959
At one extreme, major climate change and altered disturbance regimes interact to drive emergence of
novel ecosystems. Given limited experience with ecosystem turnover in many of the climate regions,
consideration of long-term paleoenvironmental records serves as a primary means for adding texture
and substance to the scenario. It also points out the risk of finding ourselves in any one of the four
quadrants. For example, transition to “novel ecosystems” is analogous to the transition observed
11,000 yr BP at Yellowstone while the transitions to “inevitable surprises” and “patches and fragments”
are analogous to the late Holocene transitions observed in (Fig. 38).
Figure 38. Multiple time scales of vegetation change. At the 15 year scale, a decline in the area of living lodgepole
pine forest in Colorado and Wyoming has arisen from a mountain pine beetle infestation; based on USFS data. At
the 150 years scale, the percent of a 129,600-ha subalpine study area in central Yellowstone occupied by three
different types of forest stages, recently burned (dashed line), even-aged (gray line), and all-age mixed stands
(black line), show successional and disturbance influences on forest structure (Romme and Despain, 1989). The
same data are shown at the 1500-year scale (extending back >250 yrs before AD 1950), and are compared with the
percent pine (Pinus) pollen in pollen records from Yellowstone region, which show limited plant assemblages
changes at that scale. More dramatic changes in plant assemblages are evident from the percent of pine and
sagebrush (Artemisia) pollen over 15,000 years from the same sites. Gray arrows indicate different types of
ecosystem changes. Pollen records shown here include the data used for climate reconstruction in Fig. 4 from
Blacktail Pond, black line (Gennett and Baker, 1986); data from location of charcoal record in Fig. 4 from Slough
Creek Pond, thick gray line (Millspaugh, 1997); data from the Romme and Despain (1989) study area from
Buckbean Fen, short dashed line (Baker, 1976); Emerald Lake, long dashed line (Whitlock, 1993); and other
datasets from Cygnet Lake Fen, gray line (Whitlock, 1993); Hedrick Pond, gray line (Whitlock 1993). Source: Jackson
et al. 2009
21
960
961
962
963
964
965
966
The greatest value in scenario planning comes from uncovering vulnerabilities and potential responses,
particularly those common to multiple story lines…
Could spend more time discussing novel climates and novel communities (Jackson et al. 2009 PNAS)
Final summary…
22
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