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Climate and Ecosystem Change in the U.S. Rocky Mountains and Upper Columbia
Basin: Historic and Future Perspectives for Natural Resource Management
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David B. McWethy*, 2Stephen T. Gray, 3Philip E. Higuera, 4Jeremy S. Littell, 5Gregory T. Pederson, 1Cathy
Whitlock*
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Dept. Earth Sciences, Montana State University, Bozeman, MT 59717
Water Resource Data System, University of Wyoming, Laramie, WY 82071
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Department of Forest Resources, University of Idaho, Moscow, ID 83844
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Climate Impacts Group, University of Washington, Seattle, WA 98195
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U.S. Geological Survey, Northern Rocky Mountain Science Center, Bozeman, MT 59715
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Introduction
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At large spatial and temporal scales, climatic conditions act as primary controls shaping the structure
and distribution of ecosystems and the species they support. Changes in climate have dramatically
altered ecosystem dynamics by shifting plant communities, creating opportunities for recruitment of
new species, and restructuring land-surface processes and nutrient cycles (IPCC 2007, Frank et al. 2009).
The controls of climate vary on different time-scales (i.e., millennial, centennial, multi-decadal, decadal,
annual and interannual), and the ecological response to climate change varies accordingly.
Paleoecological data show that ecosystems in the western U.S. have undergone significant and
sometimes rapid changes since the last glacial maximum (ca. 20,000 years ago) and the biotic
assemblages that we observe today are relatively recent phenomena (Thompson et al., 1993; Whitlock
and Brunelle, 2007, Jackson et al. 2009). The future is also likely to be characterized by rapid biotic
adjustment, including the possibility of novel assemblages as species respond individualistically to
projected climate change (Williams et al. 2007). Understanding the drivers and rates of climate change
in the past and the sensitivity of ecosystems to such changes provides critical insight for assessing how
ecological communities and individual species will respond to future climate change (Frank et al. 2009,
MacDonald et al. 2009, Shafer et al. 2005, Whitlock et al. 2003, Overpeck et al. 2003).
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The purpose of this report is to provide land and natural resource managers with a foundation of both
climate and ecosystem response information that underpins management-relevant biophysical
relationships likely to play an increasingly important role over the coming decades of rapid climate
change. To accomplish this goal we begin by synthesizing the climate and vegetation history over the
last 20,000 years following the retreat of late-Pleistocene glaciers. This time span provides examples of
ecosystem responses to long-term (e.g. millennial length) climate warming as well as several well-known
periods of rapid climate change (i.e., substantial decadal- to centennial-scale climate perturbations). To
contextualize past climate and ecosystem changes, and to provide a best estimate of future climate
conditions, we also report on the most current statistically and dynamically downscaled Global Climate
Model (GCM) projections of future changes in key climate variables (e.g. precipitation, temperature,
evapotranspiration, streamflow; [CIG 2010]). Overall, our objective is to use the past to highlight a
range of climate driven biophysical responses to illustrate potential system trajectories (and associated
uncertainties) under future climate conditions. To meet the requested needs of the U.S. National Park
Service (NPS) and the U.S. Fish and Wildlife Service (FWS) the geographic scope of this report
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encompasses the core regions of Great Northern Landscape Conservation Cooperatives (LCCs), and the
NPS high-elevation Parks; defined here as the U.S. northern and central Rockies (NR), the Greater
Yellowstone Area (CR-GYA), and the southern U.S. Rockies (SR) (Fig. 1).
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The synthesis is organized into the following four sections: (1) biophysical responses and drivers of
climate changes occurring on multi-centennial to millennial times-scales during the last 20,000 years; (2)
biophysical responses and drivers of climate changes on annual- to centennial-scales over the last 2,000
years; (3) the past century of climate and ecosystem change as observed by high-resolution instrumental
records; and (4) the next century of likely future climate and ecosystem changes under a range of
greenhouse gas emission scenarios. For each section, we present the large-scale and regional drivers of
climate change, as inferred from GCMs and paleoenvironmental data. We then detail the associated
biophysical and ecological responses documented in both modern and paleoecological proxy data with a
focus on the implications for managers tasked with maintaining key resources in the face of changing
conditions. The synthesis concludes with a discussion of challenges in planning for future conditions where there is high uncertainty in both climatic change and ecological response - by providing a
planning approach designed to deal with a wide-range of potential future conditions. The purpose of
this review is to highlight important climate and ecosystem linkages using records relevant to the
regions of great conservation and natural resource management value shown in Figure 1. Accordingly,
this document does not necessarily provide a comprehensive review of the suite of available climate and
ecosystem related research available for the entire study region.
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Figure 1. Location of study area climate region boundaries. Climate region boundaries modified from Littell et al.
(2009), Kittell et al. 2002, and, originally Bailey’s (1995) ecoprovince boundaries. Climate regions represent coarse
aggregations of biophysical constraints on modern ecological assemblages and the interaction between climate,
substrate, elevation and other conditions.
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Climate controls and variability at different spatiotemporal scales
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The influence of variation in climate on ecosystems changes at different temporal and spatial scales, and
understanding the potential influence of climate on biophysical processes often requires local- to
synoptic-scale information on the type and magnitude of forcings (e.g. local albedo and relative
humidity related changes versus changes in solar radiative output, long-lived greenhouse gasses, and
ocean-atmosphere interactions) along with an understanding of the sensitivity of the ecosystem
property that is being measured (Bartlein 1988, Overpeck et al. 2003) (Table 1). It’s also important to
recognize the hierarchical nature of climate variation and change. More specifically, understanding that
short-term (i.e., daily to inter-annual) events are superimposed on longer ones (i.e., decadal to
millennial), effectively amplifying or dampening the magnitude of the underlying physical controls that
influence ecosystem dynamics.
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Table 1. Climatic variation at different time scales and biotic response (modified from Overpeck et al. 2003).
Frequency and
Scale of
Variation (yrs.)
Kind of Variation
Driver
Biotic Response
Millennial
Deglacial and
postglacial
variations
Ice sheet size, insolation,
trace gases, regional
ocean-atmospheric-ice
interactions
Species migration, range expansion
and contraction; community
reorganization and establishment,
species extirpation/extinction
Decadal and
centennial
anomalies-climate
variations
Internal variations in the
climate system, solar
variability, volcanism
Shifts in relative abundance and
composition of different taxa through
recruitment, mortality and succession
Storms, droughts,
ENSO events
Internal variations in the
climate system, solar
variability, volcanism
Adjustments in physiology, life history
strategy, and succession following
disturbance
1000-10,000
InterdecadalCentennial
10-100
AnnualInterannual
< 10
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At the longer temporal range of major changes in earth’s climate system, climate variations on time
scales of 10,000-100,000 years are attributed to slowly varying changes in the earth’s orbit known as
Milankovitch Cycles (i.e. changes in earth’s precession, tilt, obliquity [Hays et al. 1976, Berger 1978,
Berger and Loutre 1991, Loutre et al. 2001], Fig. 2). The net result of slowly varying changes in earth’s
orbit included multiple glacial and inter-glacial cycles over the past several million years that are driven
by changes in global average temperatures. Changes in global average temperatures are initiated by
changes in the seasonal and annual cycle of solar insolation over the high-latitudes of the Northern
Hemisphere that result in substantial positive feedbacks arising from changing concentrations of
atmospheric greenhouse gasses (Vettoretti and Peltier 2004, Bond et al. 2001, Kutzbach and Guetter,
Kutzbach et al. 1998, 1993).
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Figure 2. Climate change at millennial, centennial and decadal/interannual time scales for particular records. (A)
Temperature reconstruction from the GISP2 ice core record, and the forcing mechanisms thought to influence
variation during the glacial period (ca. 49-12 ka BP) and the Laurentide ice sheet in North America began to recede
and climate warming commenced (ca. 12 ka BP – present); Gray bands in (A) indicate Younger Dryas cold period
(ca. 12,800-11,600 BP), and the 8.2 ka cool event; (B) Multiproxy temperature anomalies for the Northern
Hemisphere based on multiple proxy data (e.g., ice core, ice borehole, lake sediment, pollen, diatom, stalagmite,
foraminifera, and tree ring records) from Moberg 2005 (black line), Mann and Jones (2003, red line) and
instrumental record (blue) for the past 2000 yrs (Viau 2006). It has been suggested that Mann and Jones (2003)
indicates less variability because of regression techniques (see Moberg et al. 2005). Continental patterns of
drought, and inter-annual and decadal climate variability are indicated during the Medieval Climate Anomaly (ca.
950-1250 AD) and fewer fires during the Little Ice Age (ca. 1400-1700 AD). Time span indicated by Medieval
Climate Anomaly and Little Ice Age (give dates) from Mann et al. (2009) but varies by region (Cook et al. 2004,
MacDonald et al. 2008, Bradley et al. 2003, Carrara 1989). (C) Recent temperature anomalies for where (black line,
based on 1900-2000 mean) from HadCRUT3v global instrumental reconstruction (Brohan et al. 2006) and the
magenta line shows the Multivariate ENSO Index (MEI), an ENSO index based on six observed ocean-atmosphere
variables (Wolter and Timlin 1993, 1998). (Source NOAA Earth System Research Laboratory (ESRL)
URL:http://www.esrl.noaa.gov/psd/enso/enso.mei_index.html accessed 03/01/2010)
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On multi-millennial time scales (here specific to the last 20,000 years), the presence or absence of the
large North American ice-sheets, particularly the Laurentide ice-sheet, result in important ocean-iceatmosphere interactions that drive changes in atmospheric circulation patterns (i.e. position of
westerlies and preferential positioning of stormtracks [Fig. 2a]) that result in major changes in the
distribution and structure of ecosystems. For example, as Northern Hemisphere summer insolation
increased and the ice-sheets and glaciers began to retreat, preferential positioning of seasonal
stormtracks shifted and paleoecological records show widespread reorganization of plant communities
throughout the western U.S. (e.g., MacDonald et al. 2008, Thompson et al., 1993, Bartlein et al. 1998,
Whitlock and Brunelle 2006, Jackson et al. 2005, Betancourt et al. 1990). These large-scale and longterm changes are important features of the earth’s climate dynamics because they influence the
persistence and strength of storm tracks, subtropical high-pressure systems, ocean-land temperature
gradients, and consequently important inter-annual to decadal-scale drivers of climate variability such
as the El Niño-Southern Oscillation (ENSO). For example, higher-than-present summer insolation in the
Northern Hemisphere during the early Holocene (11,000 to 6,000 cal yr BP) led directly to increased
summer temperatures and indirectly to a strengthening of the northeastern Pacific subtropical highpressure system off the northwestern U.S. - effectively intensifying summer drought in the region
(Bartlein et al. 1998). Records of early-Holocene glacier dynamics, lake levels, aeolian activity (i.e.
blowing dust), vegetation, and fire show the effects of this large-scale control (i.e., increased summer
insolation) on ecosystem responses at local- to subcontinental-scales (e.g. Whitlock and Brunelle 2007,
Whitlock et al. 2008, Jackson et al. 2009a, Luckman and Kearney 1986, Osborne and Gerloff 1997,
Rochefort et al. 1994, Graumlich et al. 2005, Fall 1997, Booth et al. 2005, Dean et al. 1996, Dean 1997).
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On shorter and perhaps more management-relevant time scales, climate variations at inter-decadal to
centennial timescales are related to changes in solar activity, volcanism, sea-surface temperature and
pressure anomalies in both the Atlantic and Pacific Oceans - and more recently important contributions
arise from rapidly increasing atmospheric greenhouse gas concentrations (Barnett et al. 2008, Fig. 2b-c).
The Medieval Climate Anomaly could be considered an example of important centennial-scale climate
variation due to the relatively warm and dry conditions that prevailed across the Western U.S. from
approximately 900-1300 AD. Over this interval, the Western U.S. experienced substantially reduced
streamflows (Meko et al. 2007) shifts in upper treeline (Rochefort 1994, Fall 1997), and increased fire
activity (Cook et el. 2004). Thought the exact causes of the Medieval Climate Anomaly are still debated,
current prevailing evidence suggests the climatic conditions were driven by changes in solar activity,
volcanism, and perhaps sustained La Niña like conditions in the tropical pacific (Mann et al. 2009). At
decadal to inter-decadal time scales, sustained sea surface temperature anomalies in the North Pacific
and Atlantic Oceans appear to be important drivers of climate variability across western North America
(e.g. McCabe et al. 2004, Einfeld et al. 2001). The major indices that capture these important modes of
inter-decadal variability include Pacific Decadal Oscillation (PDO; see Mantua et al. 1997), and the
Atlantic Mulidecadal Oscillation (AMO; see Enfield et al. 2001, Regonda et al. 2005). Decadal-scale shifts
in climate associated with changes in the PDO and AMO are well expressed in 20th century records of
drought and winter precipitation (e.g. Cayan et al. 1998, McCabe et al. 2004) as well as proxy-based
reconstructions of precipitation and streamflow (Fig. 2d; e.g Gray et al. 2003). These events are often
regional to subcontinental in scale, initiate and terminated rapidly (within years), but often have
widespread physical and ecological effects (e.g. Allen and Breshears 1998, Bitz and Battisti 1999,
Pederson et al. 2004, Watson and Luckman 2004). Examples include widespread bark beetle outbreaks,
increased forest fire activity and stress related tree mortality, and rapid changes in glacier mass balance,
snowpack and streamflow.
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The El Niño Southern Oscillation (ENSO) is a major global control of both temperature and moisture
patterns that operates on inter-annual (i.e. year to year) time scales (Fig. 2c, 3). ENSO events are
defined by changes in atmospheric pressure gradients across the tropical Pacific that are related to
patterns of warming (El Niño) and cooling (La Niña) in the central and eastern equatorial Pacific with a
typical duration of 6 to 18months, and typically reoccur every 2 to 7 yrs (Philander 1990, Chang and
Battisti 1998). The strength of these ocean-atmosphere teleconnections varies from event to event, but
typically exerts a substantial influence on regional temperature and precipitation patterns (McCabe et
al. 2004, Einfeld et al. 2001). For example, El Niño events typically result in warm and dry conditions
across the northwestern U.S. and a southerly displacement of the winter stormtrack (i.e. the jet stream)
resulting in cool and wet conditions across the southwestern U.S. (McCabe et al. 2004, Einfeld et al.
2001). The inverse is typically true for La Niña-like conditions, but in all cases this mode of climate
variability appears to exert its strongest influence across the southwestern U.S. (e.g. Swetnam and
Betancourt 1998) with and important, but somewhat attenuated spatial signature across the
northwestern U.S. (e.g. Dettinger et al. 1998). In summary, climate-ecosystem linkages are evident
across many time scales, from individual records as well as regional and global compilations. Shifts in
species distributions and abundance are a response to climate variations occurring over from seasons to
millennia. Knowledge of climate drivers on all time scales is necessary to better identify temporal and
spatial dimensions of future changes and possible ecological responses to these changes.
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How can understanding millennial to interannual scale climatic variability inform management?
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Knowledge of natural variations in climate at different scales provides a context for understanding how
communities and individual species might respond to current rates of change and rates predicted for the
future (Table 2). For example, past changes in fire regimes have been largely driven by large-scale
change in climate on millennial, centennial, decadal and annual time scales (Whitlock and Brunelle 2007,
Kitzberger et al. 2007, Littell et al. 2009, Westerling et al. 2006). In ecosystems where fire regimes are
expected to change with future climate conditions, management efforts should focus on the ecological
response to rapidly changing and new conditions as opposed to maintaining current or past conditions
(Whitlock et al. 2009). Additionally, paleoenvironmental records showing evidence of rapid changes in
climate and attendant ecological responses suggest that even small changes in climate can have large
consequences and provide an important context for anticipating ecosystem response to future climate
change (Whitlock and Brunelle 2007, Gray et al. 2004, MacDonald et al. 2009, Gray et al. 2006, Lyford et
al. 2003).
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Table 2. Examples of paleoenvironmental proxy data, spatiotemporal resolution and range and type of
reconstruction.
Proxy
Source
Tree
growth
Charcoal
Tree-cores
Pollen
Oxygen
Isotopes
Lake-peat
sediments
Lake-sediments
Corals, tree,
lake, ocean or
ice cores
Temporal
Resolution
High (annual)
Spatial
Resolution
High
Temporal range
High (annualdecadal)
Moderate
(multidecadalcentennial)
High (annual) to
moderate (decadal)
High-moderate
(1 to several km2)
Moderate
(several km2)
High (many
millennia)
High (many
millennia)
Moderate to low
High to very high
(100,000 +)
100-1000’s yrs
Type of
reconstruction
Temperature,
moisture
Fire
Vegetation
Temperature,
ice volume
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How do we know about past conditions? Reconstructing past environments using proxy data.
Reconstructing past climatic conditions and the associated ecological response involves a number of direct and
indirect measurements of past change. Direct measurements include ground temperature variations, gas
content in ice core air bubbles, ocean sediment pore-water change and glacier extent changes. Indirect
measurements or paleoclimate proxy typically come from organisms (e.g., trees, corals, plankton, diatoms,
chironomids and other organisms) that are sensitive changes in climate through changes in growth rates,
abundance, or distribution. Past changes are thus recorded in past growth of living or fossil specimens or
assemblages of organisms. Each proxy represent past change at different scales and resolutions, representing
millennial to centennial, decadal and inter-annual variation in conditions.
Lake-sediment, tree-ring cores and packrat middens are three of the primary proxy used for
reconstructing past conditions for the western U.S. (Whitlock et al. 2006, Whitlock and Larsen 2001, Fritts and
Swetnam 1991, Betancourt 1990). Lake-sediment cores often provide some of the longest records of vegetation
and fire through analysis of pollen, plant macrofossils, and charcoal particles at regular intervals throughout the
core. Most lakes in the northwestern U.S. were formed during the course of deglaciation, and thus they provide
a sedimentary record spanning the last 15,000 years or longer, depending on the time of ice retreat. Sediment
cores are retrieved from modern lakes (and wetlands) using anchored platforms in summer or from the ice
surface in winter. Samples for pollen and charcoal analyses are removed from the cores at regular intervals that
depend on the detail and temporal resolution required. The pollen extracted from the sediment is chemically
treated, identified under the microscope, and tallied for each sediment level sampled. Pollen counts are
converted to percentages of terrestrial pollen and these are plotted as a pollen percentage diagram. The
reconstruction of past vegetation (and climate) from pollen percentage rests on the relationship between
modern pollen rain and present-day vegetation and climate. Modern pollen samples have been collected at
lakes throughout North America, and this information is calibrated to the modern vegetation and climate. Past
fire activity is inferred from the analysis of particulate charcoal, which is extracted and tallied from the sediment
cores (Whitlock and Bartlein 2004). High-resolution charcoal analysis involves extraction of continuous samples
from the sediment core such that each sample spans a decade or less of sediment accumulation. These samples
are washed through sieves and the charcoal residue is tallied under a microscope. Examining these relatively
large particles enables a local fire reconstruction, because large particles do not travel far from a fire. Charcoal
counts are converted to charcoal concentration, which is then divided by the deposition time of each sample
(yr/cm) to yield charcoal accumulation rates (particle/cm2/yr). Detection of fire events involves identification of
charcoal accumulation rates above background levels.
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How do we know about past conditions? Reconstructing past environments using proxy data –
Continued…
Tree-rings provide records of past change at centennial and millennial time scales and have several features that
make them well suited for climatic reconstruction, such as ease of replication, wide geographic availability,
annual to seasonal resolution, and accurate, internally consistent dating. Networks of tree-ring width and
density chronologies are used to infer past temperature and moisture changes based on calibration with recent
instrumental data, recording past change for century to millennia. Tree growth is highly sensitive to
environmental changes and therefore tree-ring records are powerful tools for the investigation of annual to
centennial variations in climate. Tree-ring chronologies are used to reconstruct past climates such as growing
season temperature, and precipitation. The most sensitive trees are those growing in extreme environments
where subtle variations in moisture or temperature can have a large impact on growth. For example,
precipitation and/or drought reconstructions are often derived from extremely dry sites, or sites at
forest/grassland boundaries, where moisture is the strongest limiting factor of growth. Similarly, sites at
altitudinal and latitudinal treelines with ample moisture are often targeted for temperature-sensitive
chronologies. The year-to-year variability in individual tree-ring-width series (or other tree-ring parameters
such as density) from long-lived stands of trees are combined to produce site histories (or chronologies) that
span centuries or millennia. These chronologies contain considerable replication (e.g. two cores per tree,
minimally 10-15 trees per site) and dating accuracy is rigorously verified by comparing ring-width patterns
among trees, also known as “crossdating”. Cross-dating also allows tree-ring series from ancient dead wood
(found in historic dwellings, lakes, sediments and on the surface in cold/dry environments) to be combined with
overlapping records from living trees, thereby extending records further back through time. Statistical
relationships established between annual tree ring-width chronologies and instrumental climate records are
used to hindcast estimates of precipitation and/or temperature back through time.
Packrat middens left by woodrats of the genus Neotoma also provide long-term records of past change
and are often found in arid environments where other approaches for reconstructing past environments are less
viable. When packrats build nests, plant and animal remains often become crystallized and mummified in
packrat urine, preserving rich deposits of macrofossils that can be used to reconstruct past vegetation and
climate. Middens located in caves or under rock ledges shelters that provide protection from water are
especially well preserved. Once packrat middens are excavated from a site, plant and animal parts are
dissected, identified and aged using radiocarbon dating techniques. A single midden typically represents a
relatively discrete time interval when material was accumulated (one to several decades, [Finley 1990]) but a
network of middens within one site can be stacked chronologically to provide a record of vegetation and climate
change over a longer time period. A reconstruction of vegetation surrounding the site typically represents the
vegetation within 30-100 meters surrounding a site (USGS).
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Past 20,000 Years of Paleoenvironmental Change in the Western U.S.
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Drivers of millennial-scale climate variation
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The climate variations of the last 20,000 years occurring on millennial time-scales are best understood
by paleoclimate model-based simulations that look at the regional response to large-scale changes in
the climate system, and by paleoenvironmental data that measure specific components of climate
change. Broad-scale climate variations were described by Bartlein et al. (1998) using NCAR CCM1 (4.4˚
latitude by 7.5˚ longitude spatial resolution, mixed-layer ocean, crude depiction of western Cordillera
topography). The model simulation provided estimates of climatic conditions over six discrete time
periods (21,000 cal yr BP=full glacial period with full-sized ice sheets; 16,000 and 14,000 cal yr BP =lateglacial period with shrinking ice sheets; 11,000 cal yr BP = early Holocene insolation maximum, 6000 cal
yr BP= middle Holocene transition, and present). Other recent regional-scale modeling studies have
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provided better temporal and spatial resolution with more realistic topography (Hostetler 2009,
Hostetler et al. 2003). Results from these efforts are summarized below.
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Figure 3. Area of the Laurentide ice sheet (LIS) over time, modified from Shuman et al. 2002 (area of LIS estimated
from Dyke and Prest (1987) and Barber et al. (1999) and oxygen isotope record (bottom panel black line)
associated with variations in Northern Hemisphere temperature from GISP2 (Stuiver et al. 1995). The Younger
Dryas Chronozone (YDC) was an abrupt cooling ca. 12,900-11,600 which represented a temporary reversal in
warming during the Pleistocene/Holocene transition (Alley et al. 1993)(modified from Shuman et al. 2002).
At the time of the Last Glacial Maximum (ca. 21,000 cal yr BP), the large Laurentide and Cordilleran icesheets strongly influenced climatic conditions in the western U.S. (Bartlein et al. 1998, Fig. 3).
Specifically, the presence of the ice sheets depressed temperatures (by approximately 10˚C for areas to
the south of the ice sheets) and steepened the latitudinal temperature gradient. Additionally, the
presence of the large ice sheets displaced the jet stream south of its present position greatly reducing
winter precipitation in the northwestern U.S. and Canada while increasing precipitation across the
southwestern U.S. Another element of the full-glacial climate was the stronger-than-present surface
easterlies related to a strong high-pressure system that persisted over the ice sheets. This steepened
latitudinal temperature gradient and weakened the westerly stormtracks, resulting in colder and
effectively drier conditions across the northwestern U.S. and cold wet conditions in the southwestern
U.S. (Hostetler et al. 2000, Bartlein et al. 1998, Thompson et al. 1993).
The glacial/Holocene transition, spanned 16,000 to 11,000 yrs BP, and was a period of increasing
summer insolation in the Northern Hemisphere. Solar insolation over the high-latitude Northern
Hemisphere landmasses increased during this period and reached a maximum ca. 11,000 year BP when
summer insolation was 8.5% higher than present and winter insolation was 10% lower than present at
45˚N latitude. One consequence was a northward shift of winter storm tracks from its full-glacial
position, which brought wetter winter conditions to the northwestern U.S. whereas the southwestern
U.S. became increasingly dry (Bartlein et al. 1998). Increasing summer insolation resulted in warmer
growing season temperatures causing alpine glaciers and ice sheets to melt (or ablate) rapidly (see Fig. 4
for example of modern air mass circulation).
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Figure 4. Primary air masses that influence study
area today. Ice-sheet dynamics and oceanatmosphere interactions have altered large-scale air
mass circulation patterns throughout the lateglacial and Holocene, influencing important climatic
variables including temperature and precipitation.
The U.S. Rocky Mountains lie in an important
transition zone where many of these air masses
interact to strongly influence local and regional
climate. Changes in the position and strength of
these air masses, driven by large-scale changes in
the climate system, are responsible for climatic
variations on different time scales and act as an
important control on vegetation as discussed
further in the text (modified from Ahrens 2008 and
Aguado and Burt 2008).
Greater-than-present summer insolation (8% above present) and lower-than-present winter (8% lower
than present) insolation in the early Holocene (11,000 to 7,000 cal yr BP) profoundly affected the
climate and ecosystems of the western U.S. Directly, increased summer insolation led to warmer
temperatures throughout the region during the growing seasons, while winters were likely colder than
at present. Indirectly, model simulations show that increased insolation led to a strengthening of the
eastern Pacific subtropical high-pressure system which suppressed summer precipitation in much of the
northwestern U.S. At the same time, it strengthened inflow of moisture from the Gulf of California to
the southwestern U.S. and southern and central Rocky Mountain region (Whitlock and Bartlein 1993),
resulting in greater than present summer precipitation. East of the Rocky Mountains, increased summer
precipiaton was likely offset by increased temperatures and increased rates of evapotranspiraton,
meaning conditions were effectively drier than present which is consistent with low lake levels and dune
activation (Shuman et al. 2009, Stokes and Gaylord 1993). Today the western U.S. is characterized by
summer-dry areas under the influence of the subtropical high (i.e. the Northwestern U.S.) and summerwet areas where summer precipitation reflects monsoon activity (i.e. the Southwestern U.S.). These
two precipitation regimes are defined by topography (e.g., Yellowstone Plateau) and the boundary
between them is relatively sharp (Whitlock and Bartlein 1993, Gray et al. 2004). The indirect effects of
greater-than-present summer insolation strengthened both of these precipitation regimes, making
summer-dry regions drier in the early Holocene and summer-wet regions wetter than at present
(Thompson et al. 1993, Whitlock and Bartlein 1993, Bartlein et al. 1998). During the middle Holocene
(ca. 7000 to 4000 cal yr BP) and the late Holocene (4000 cal yr BP to present), summer insolation
decreased (and winter insolation increased) gradually to present levels. Summer-dry regions became
cooler and wetter and summer-wet regions became cooler and drier than before.
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Mid-Holocene Transition Period (ca. 7,000-4,000 cal yr BP)
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In the Pacific Northwest, including the Columbia Basin, the middle Holocene is a period of transition
between the warm dry early Holocene and the cooler wetter late Holocene. In British Columbia, Hebda
and Mathewes (1984) term this period the mesothermal and trace the expansion of hemlock and
Douglas fir during this period. This same signal is also evident in the northern Rockies and perhaps as far
south as the Great Yellowstone region. In the southern and central Rockies, both summer-wet regions,
paleoclimate data suggest several anomalously dry/wet periods during the Holocene (Shuman et al.
2009, Stone and Fritz 2006). Lake-level data for example indicate that numerous sites throughout the
Rocky Mountains experienced drier conditions than today during the middle Holocene (ca. 6,000 cal yr
BP) (Shuman et al. 2009), which is similar to climatic conditions across the Great Plains. Anomalously
dry conditions for the central and southern U.S. Rockies also contrast with wetter-than-present
conditions in the American Southwest (Betancourt et al. 1990, Davis and Shafer 1992, Thompson et al.
1993, Fall 1997, Mock and Brunelle-Daines 1999, Harrison et al. 2003). Regional climate simulations
from the Colorado Rockies shed some light on the drivers of mid-Holocene aridity (Diffenbaugh et al.
2006, Shuman et al. 2009). They note that variations in the seasonal cycle of insolation imposed localscale surface feedbacks (e.g., reduced snowpack and soil moisture) that were important drivers of
submillennial-scale changes in precipitation and moisture (Shuman et al. 2009, Diffenbaugh et al. 2006).
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At the end of the late-Pleistocene, much of the Northern Hemisphere experienced an abrupt cooling
known as the Younger Dryas cool period (YDC; ca. 12,900-11,600 BP) (Alley et al. 1993). This event is
clearly registered in the North Atlantic region and across Europe, and is related to changes in ocean
circulation during the melting of the Laurentide ice sheet. Evidence of the YD cooling event is less
obvious in the western U.S., although in some areas, a readvance of mountain glaciers (Osborn and
Gerloff 1998, Reasoner and Huber 1999, Friele and Clague 2002, Menounos and Reasoner 1997) and
changes in vegetation indicate cooling (Reasoner and Jodry 2000). During the YDC, temperatures
decreased as much as 5-10 ˚ C in Greenland but pollen data for the northwestern U.S. indicate much
more moderate cooling 0.4-0.9 C (Reasoner and Jodry 2000). The YDC ended rapidly with warming of ~7
˚C in Greenland occurring within one to several decades (Alley 2000). Consequently, the period is
closely scrutinized as an analogue for abrupt climate change (Alley et al. 2003). The evidence of YDC is
indicated by a lowering in treeline in central Colorado (Reasoner and Jodry 2000), and changes in the
isotopic signature in speleothem data. Most paleoenvironmental data from the western U.S. show little
or no response to the YDC, in terms of glacial activity (Licciardi et al. 2004, Heine 1998), vegetation
change (Grigg and Whitlock 2002, Briles et al. 2005, Brunelle et al. 2005, Huerta et al. 2009), or shifts in
fire activity (Marlon et al., 2009).
A similar abrupt cooling occurred at ca. 8.2 ka BP (Figs. 2, 4), where a large influx of freshwater disrupted
circulation in the North Atlantic, causing cooling lasting several centuries in the North Atlantic region (Le
Grande et al. 2006, Alley and Agustsdotti 2005). The 8.2 ka event again is not registered at many sites in
the western U.S., either because the signal is weak or the sampling resolution is inadequate to detect it.
Geochemical proxies from lake sediments suggest that it has been associated with drier conditions at
Bear Lake, UT (Dean et al. 2006). The YDC and the 8.2 ka event illustrate how rapid climate changes due
to ocean-atmosphere-ice interactions can occur. With respect to this evaluation, it is also important to
recognize that they also show a variable signal in regions distal to the North Atlantic origin, so that some
sites show a response but others do not in the western U.S.
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Southern Canadian and Northern U.S. Rockies
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Prior to 14,000 cal yr BP, much of the Northern Canadian and U.S. Rockies were glaciated (Fig. 5). The
first plant communities to colonize deglaciated regions were alpine in character, composed of shrubs
and herbs dominated by sagebrush, grasses and alder (Artemisia, Gramineae and Alnus) (Whitlock et al.
2002, MacDonald 1989, Reasoner and Hickman 1989). By the early-Holocene, warmer-than-present
conditions allowed shrub and herb communities to spread upslope to higher elevations than at present,
and forests of pine and alder (Pinus albicaulis/flexilis and Alnus) developed at mid elevations. The lesser
abundances of spruce and pine (Picea and Pinus contorta) suggest warmer growing season conditions
than present. Low-elevation forests were characterized by spruce and lodgepole pine (Pinus contorta)
and valley floors supported open grassland with limber pine. Upper treeline was at least 90 m above
modern elevation during the early- and middle-Holocene from ca. 8500 to ca. 3000 yr BP. Low/midelevations experienced the greatest fire activity during the early Holocene ca. 9-8 ka BP (MacDonald
1989) and forest composition resembled modern subalpine spruce-fir (Picea-Abies) forest. The lateHolocene was marked by cooler summer (growing) season conditions than over the early- and midHolocene, increased glacial activity, and a substantial lowering of upper treeline (Reasoner and Hickman
1989). Cooler and moister conditions led to tundra at high-elevations, spruce and fir forests at midelevations, and spruce, lodgepole pine and aspen at low elevations.
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Figure 5. Ecological response to changing climatic conditions following the retreat of glaciers in the southern
Canadian Rockies (derived from MacDonald 1989 and Reasoner and Hickman 1989)(Whitlock, unpublished).
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Ecosystem response to centennial-scale climatic variations is evident from a 3,800 year history of
climatic, vegetation, and ecosystem change inferred from pollen and charcoal concentrations in the
lake-sediment record from Foy Lake in northwestern Montana. Formed over 13,000 years ago as ice
retreated from the Flathead Valley, the lake is situated at the eastern edge of the Salish Mountains 3 km
southwest of Kalispell, Montana (Stevens et al. 2006). Several studies from the site provide historical
reconstructions of climate and hydrologic variability and ecosystem response to climate change over the
past several millennia (Stevens et al. 2006, Power et al. 2006, Shuman et al. 2009). Paleolimnologic and
pollen data indicate that ca. 2700 BP, an abrupt rise in lake levels coincides with a transition from steppe
and pine forest to pine forest/woodland to mixed conifer forest (Power et al. 2006), a transition linked
to an increase in effective moisture (winter precipitation) shown in lake level records (Stevens et al.
2006, Shuman et al. 2009). Following the establishment of mixed conifer forests, lake levels decreased
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from 2200 to 1200 cal yr BP and increases in grass, pine and sagebrush and declines in Douglas-fir/larch
lead to the development of a steppe/parkland/forest mosaic ca. 700 yr BP (Power et al. 2006, Stevens et
al. 2006). Recent increases in grass and sagebrush (late 19th and early 20th centuries) coincide with
human activities. Notable climatic events during this period include a long, intense drought ca. 1140 AD
following a wetter period from 1050 to 1100 AD (Stevens et al. 2006).
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Figure 6. Pollen percentage of selected pollen taxa. AP/NAP ratio includes all upland tree pollen divided by all nonaquatic shrubs and herbs. Gray silhouettes are 5_ exaggeration of pollen percentage data. Fire frequency is
recorded as number of episodes/1000 year. Source: Power et al. 2006, copyright reprint permission needed
These shifts in vegetation were accompanied by pronounced changes in patterns of fire as evidenced by
the charcoal record (Fig. 6). Intervals dominated by forests coincide with high-magnitude and frequent
fires (e.g., stand replacing fires), periods dominated by steppe/parkland vegetation are associated with
smaller and less frequent fires (surface fires) and a decline in charcoal deposition in the last century
likely reflects the impact of fire suppression in the area (Fig. 6). The Foy Lake record demonstrates the
impacts of centennial scale climatic variations and their associated ecosystem response during the past
two millennia. While relatively modest changes in vegetation cover occurred after the conifer forests
were established (ca. 2700 yr BP) multi- decadal shifts in climate are evident in the fire reconstruction
for the past several millennia.
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Central U.S. Rockies and the Greater Yellowstone Area (GYA)
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Pollen records from Yellowstone and Grand Teton National Parks show the nature of biotic change that
occurred in conjunction with broad climatic changes at different elevations throughout the Parks (Fig. 7,
Whitlock 1993). Deglaciation (ca. 17-14 ka BP) was followed by colonization of tundra vegetation.
During the late-Glacial and into the mid-Holocene, tundra vegetation was replaced by subalpine
communities of spruce, fir, and pine in many regions, first as an open parkland and then as a closed
forest. With the warmer growing season conditions of the eary-Holocene period, pine, juniper and birch
trees were present at low-elevations; whereas lodgepole pine (Pinus contorta), Douglas-fir (Pseudotsuga
menziesii) and aspen (Populus) trees established and characterized mid-elevation forests. Subalpine
forests (Picea and Abies) expanded their ranges upslope to higher elevations, and upper treeline was
located at an elevation similar to the present-day upper treeline elevation. Decreased summer
insolation in the late-Holocene (ca. 3-4 ka BP) led to cooler wetter conditions. Sagebrush-steppe
became present at low-elevations, and forests of limber pine (Pinus flexilus), Douglas fir and lodgepole
pine dominated mid-elevations (Whitlock et al. 1993). Subalpine communities were comprised of mixed
spruce, fir and pine forests, and the increasingly cooler conditions resulted in a lowering of upper
treeline to an elevation close to its present-day position.
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Figure 7. Ecological response to changing climatic conditions following the retreat of glaciers in the summer-dry
region of Greater Yellowstone Area (derived from Whitlock 1993).
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Figure 8. Ecological response to changing
climatic conditions following the retreat of
glaciers in Yellowstone National Park
(Millspaugh et al. 2000). Pollen and charcoal
diagram from Cygnet Lake, central
Yellowstone. Modified from Millspaugh et al.
2000
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Cygnet Lake
An examination of charcoal, pollen and
climate conditions from central
Yellowstone provides an example of the
tight linkage between fire and climate,
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even in the absence of vegetation change (Figure 8; from Millspaugh et al. 2000, Whitlock 1993). Pollen
from Cygnet Lake in the Central Plateau of Yellowstone indicates the area was dominated by tundra
community of sagebrush (Artemisia) and grass (Poaceae) prior to 12,000 cal yr B.P. when cool and wet
conditions prevailed. After 12,000 cal yr BP, the vegetation was dominated by lodgepole pine (Pinus
contorta), which is favored in areas of infertile rhyolite soils. The last 12,000 years at Cygnet Lake
reveals little change in vegetation in contrast to other sites on non-rhyolite substrates that show strong
responses to early-, middle-, and late-Holocene climate forcings (Whitlock, 1993). Despite the
complacency of the vegetation, the fire record at Cygnet Lake shows the effect of early-Holocene
drought on fire regimes with charcoal data showing rising fire activity in response to increasing summer
insolation. Fire return intervals were 75-100 years between 11,000 and 7,000 cal yr BP, and after 7,000
cal yr BP cooler wetter conditions coincide with decreasing fire frequency. The present fire return
interval of 200-400 years was reached in the late-Holocene.
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Southern U.S. Rockies
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A network of vegetation reconstructions from pollen and macrofossil data provides a history of climate
and vegetation change in southwestern Colorado during the late-glacial and Holocene (Fall 1997). Prior
to 14,000 cal yr BP, cooler (2-5˚ C cooler than present) and wetter (7-16 cm wetter than present)
conditions supported tundra vegetation at high-elevations (~300-700 m below present treeline), and
spruce parkland at low-elevations (Fig. 9, Fall 1997). During the late-glacial fir (Abies) increased in
abundance in and subalpine forest was established at low- and mid-elevations (Fall 1997). Summer
insolation increased during the early Holocene, and warmer temperatures allowed subalpine forests to
expand upslope above its present position. Though the central and northern U.S. Rocky Mountains
generally experienced warm and dry conditions during the early Holocene, the southern U.S. Rockies
(SR) experienced warmer and wetter growing season driven by a more intense North American
Monsoon (Thompson et al. 1993) as evidenced by Markgraf and Scott (1981) who record a retreat of
subalpine forest upslope (due to warmer conditions) and an expansion of pine forests at both lower and
upper treeline facilitated by warmer but still wet conditions. Similarly, Fall (1997) found that from 9-4
ka BP, warm summer temperatures (1.9˚ C mean summer temperature above present) facilitated the
upslope expansion of subalpine forests of spruce and fir (Picea engelmannii and Abies lasiocarpa), at
least 300 m greater than today to and elevation of almost 4,000 meters (Fall 1997). These forests also
expanded to lower elevations below 3000 m, along with montane conifers (Pinus spp.). These
assemblages persisted through the mid-Holocene where summer precipitation is estimated to have
increased by 8 -11 cm (Fall 1997).
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Figure 9. Ecological response to changing climatic conditions following the retreat of glaciers in the Southern U.S.
Rockies (derived from Fall 1997).
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In the late Holocene, lower elevation montane forests mixed with steppe vegetation, and low-elevation
subalpine forests were defined by an increasingly open stand structure. Summer temperatures declined
to pre-industrial era (ca 1850 AD) values, and spruce and fir (Abies lasiocarpa) dominated subalpine
forests and krummholz vegetation. Montane taxa retreated upslope, sagebrush steppe expanded at
lower elevations and alpine tundra dominated larger range of high-elevations - suggesting drier
conditions increased for parts of the SR (Markgraf and Scott 1981, Fall 1997). Conditions similar to
present were established approximately two millennia ago with modest treeline elevation fluctuations
during the Medieval Climate Anomaly.
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Upper Columbia Basin
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Paleoenvironmental data are available from the Upper Columbia Basin (UCB) (Barnosky 1985, Whitlock
et al. 2000, Mack et al. 1978, 1976, Mehringer 1996, Blinnikov et al. 2002), the Snake River Plain (Davis
1986, Beiswenger 1991, Bright and Davis 1982) and the mountains of southern Idaho and Montana
(Doerner and Carrara 2001, Whitlock et al. in review, Mumma 2010).
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Figure 10. Ecological response to changing climatic conditions in the western Columbia Basin following
deglaciation (derived from Carp Lake, Washington [Whitlock et al. 2000]).
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During glacial times, tundra-steppe communities dominated by Artemisia and Poaceae were widespread
in the basins, reflecting cold, dry conditions (Fig. 10-11). As the climate warmed in the late-glacial
period, pine (Pinus), spruce (Picea) and fir (Abies) parkland developed. The early-Holocene period in the
western Columbia Basin and Snake River Plain featured steppe vegetation, and records from adjacent
mountains record an expansion of juniper (Juniperus), sagebrush (Artemisia), and Chenopodiaceae.
Summer drought was more pronounced throughout much of the region during the early Holocene as
the amplification of the seasonal cycle of insolation resulted in warmer and effectively drier conditions
at mid and lower elevations (Whitlock et al. 2000). Cool and dry conditions in the Okanogan Highlands
continued to support grasses and sagebrush and the forest-steppe ecotone was located north of its
present location by as much as 100 km (Mack et al. 1978). Pine, spruce and fir were present in areas
with greater precipitation and (e.g., Waits Lake, eastern Washington, Whitlock and Brunelle 2006, Fig.
11). In the middle Holocene, increased effective moisture allowed for the establishment and expansion
of Pinus woodland at middle and higher elevations, and like other regions, upper treeline lay above its
present elevation. In the western Columbia Basin, the expansion of pine woodland (primarily Pinus
ponderosa) was followed by the invasion of mixed forest in the late Holocene (Douglas-fir, larch, fir,
western hemlock, and oak, Whitlock and Brunelle 2006). Some sites at higher elevations signal an
interval of cooling during the late Holocene (ca. 1.7-3.5 ka BP) when the abundance of spruce (Picea)
and fir (Abies) pollen increase (Whitlock and Brunelle 2006) and led to an expansion of mixed forests
throughout the Okanogan highlands. Modern assemblages of Douglas-fir, fir, western hemlock and
spruce were established during this time period (Whitlock and Brunelle 2006).
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Figure 11. Late-glacial and early-Holocene vegetation history along the southern margin of the Cordilleran ice
sheet based on a transect of pollen records from western Washington to western Montana. Abies, fir; Alnus, alder;
Artemisia, sagebrush; cal yr BP, calendar years before present; ID, Idaho; MT, Montana; Picea, spruce; Pinus
contorta, lodgepole pine; Poaceae, grass; Pseudotsuga, Douglas-fir; Shepherdia, buffaloberry; Tsuga heterophylla,
western hemlock; WA, Washington. Source: Whitlock and Brunelle 2006, copyright reprint permission needed
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What can we learn from long-term paleoenvironmental records?
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Representative of broad-scale changes that occurred for much of the study area during the Holocene,
these case studies paint a general picture of dynamic vegetation response to warming and subsequent
deglaciation during the early-Holocene, warmer-drier conditions during the mid-Holocene, and the
establishment of the cool/moist conditions during the late-Holocene that characterize most of the U.S.
Rockies and Upper Columbia Basin today. These longer term records of past change in the U.S. Rockies
and Upper Columbia Basin indicate:
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
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Latitudinal and seasonal distribution of incoming solar radiation and the presence of the
Laurentide ice sheet were important drivers of atmospheric circulation (e.g., southward
displacement of jet stream) and the general climate of the Rocky Mountains and Upper
Columbia Basin at the end of the last glacial.
Deglaciation was followed by dramatic biotic reassembly and millennial scale variation in
insolation and ocean-atmosphere and land-surface interactions became the primary driver of
vegetation change.
At smaller temporal and spatial scales, local factors (e.g., substrate, disturbance) interact with
climatic variation to influence the distribution of vegetation.
Large-magnitude and abrupt climatic changes are evident in paleoclimatic records (e.g., the
Younger Dryas, 8.2 ka event, MWP, LIA) and vegetation response was rapid, occurring on lags of
only years to decades following some of these events.
Several regionally synchronous events occurred during the Holocene (e.g., Medieval Climate
Anomaly and the Little Ice Age) but were not uniformly manifested across the study area. For
example, the warm and dry conditions during the early Holocene and MCA led to increased fires
in the summer-dry areas of Yellowstone whereas the most severe drought conditions in the
summer-wet areas of Yellowstone occurred in the past 7,000 yrs. (Whitlock and et al. 2003).
Centennial, multi-decadal and decadal scale droughts have occurred throughout the Holocene
and prolonged droughts have occurred in the last millennia that rival 20th century droughts.
Additionally, these case studies provide important lessons for understanding climate-vegetation
linkages. First, the four climate regions are characterized by heterogeneous biophysical and
environmental conditions and are influenced by contrasting patterns of atmospheric circulation that are
evident during the Holocene (e.g., summer-wet and summer-dry regions in close proximity in
Yellowstone). For example, warm dry conditions during the early-Holocene increased fire activity in the
Central Plateau of Yellowstone but other sites in summer-wet settings of northern Yellowstone
experienced wetter conditions and lower fire activity (Whitlock and Bartlein 1993, Huerta et al. 2008,
Millspaugh et al. 2004). The contrast between summer-wet and summer-dry regions was greatest in the
early-Holocene and weakened in the late-Holocene and the boundaries between summer-wet and
summer-dry conditions defined by complex topography are persistent but may be defined by different
climate responses in the future. Second, while climate exerts strong controls on the distribution of
vegetation at large spatio-temporal scales, other factors also act as important controls on vegetation.
For example, the persistence of lodgepole pine forests on rhyolitic soils in the Central Plateau of
Yellowstone throughout much of the Holocene illustrates how substrate can act as an important longterm control on vegetation in some settings (Whitlock et al. 1993, Millspaugh et al. 2000) (see also Briles
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et al. in review). Third, climate change can influence vegetation via direct climate constraints or by
influencing key ecosystem processes such as fire. Feedbacks related to changes in vegetation can also
influence ecosystem processes such as fire by increasing fuel availability. Lastly, conditions such as
vegetation structure, composition and patterns of fire at the time of European settlement are often
used as a baseline for restoration efforts despite the fact that these states are relatively recent
phenomena and suggest that:
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Natural variability in climate, change and accompanying ecological responses are not adequately
represented in the instrumental and historical record of the past century and thus do not provide
adequate scale for considering baseline natural variability for restoration efforts.
Many terrestrial ecosystems like those represented in these case studies are relatively young,
dynamic and have only recently established in response to the interaction of a number of climate
drivers operating at different scales.
As evidenced by changes in fire regimes of the later 19th and early 20th centuries, human activities
have left a strong imprint on the landscape and can rapidly influence ecosystems and key ecosystem
processes such as fire.
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Past 2,000 Years of Paleoenvironmental Change
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Primary drivers of change
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The primary external climate forcings during last 2,000 years include ocean-atmosphere interactions,
volcanic eruptions, changes in incoming solar radiation, and increases in atmospheric greenhouse gases
and aerosols due to human activities (Fig. 2). Greenhouse gases and tropospheric aerosols varied little
from AD 1 to around 1700 AD when human activities began to significantly impact levels of greenhouse
gases and aerosol concentrations (IPCC 2007, Keeling 1976). Prior to the rise of anthropogenic GHGs,
volcanic eruptions and solar fluctuations were likely the most strongly varying external forcings during
this period. Climate model simulations indicate that solar and volcanic forcings combined with oceanatmosphere interaction likely resulted in periods of relative warmth and cold during the preindustrial
portion of the last 2,000 years (Amman et al. 2007, Mann et al. 2003, Mann 2007, Jones et al. 2009,
Esper et al. 2002, 2010). The recent rapid rise in 20th century global temperatures; however, are best
explained by the combination of natural plus anthropogenic GHG forcings, with GHGs playing an
increasingly dominant role over recent decades.
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When compared to the conditions driving continental deglaciation and the Pleistocene/Holocene
transition, orbital and radiative forcings over the past two millennia have remained relatively constant
and can be considered more analogous to modern conditions. However, major centennial-scale climate
variations are evident during this time period and can be linked to changes in solar output, volcanic
forcing, and ocean atmosphere interactions (Amman et al. 2007, Mann et al. 2009, Mann 2007, Jones et
al. 2009, Esper et al. 2002, 2010). Because the rates of climatic change during the past two millennia
were much smaller in magnitude than those associated with the late glacial and early Holocene, the
ecological response to climatic variation during the past two millennia was generally less dramatic. For
much of the study area considered in this synthesis, climatic variation during this period led to shifts in
the extent and abundance of modern vegetation assemblages and rarely led to widespread changes in
Biophysical Conditions
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dominant vegetation. Since the late-Holocene, vegetation types throughout this period are considered
similar to present-day conditions and, overall, the magnitude and duration of the changes are not
comparable to those of the Pleistocene/Holocene transition. Rather, the past 2000 years of change
reflect smaller magnitude, and shorter duration (centennial, decadal, inter-annual) climatic controls on
ecosystems - that never the less result in societally- and ecologically-relevant changes in both
ecosystems and natural resources. At these shorter time scales ocean-atmosphere interactions such as
the Pacific Decadal Oscillation (PDO), the North Atlantic Oscillation (NAO), the Atlantic Multidecadal
Oscillation (AMO), and the El Nino Southern Oscillation (ENSO) interact to influence temperature,
precipitation, and atmospheric circulation and help explain droughts, wet periods at inter-annual to
inter-decadal time scales (e.g. some citations). Past records of temperature illustrate centennial,
decadal and interannual variation which and provide an important context for understanding ecosystem
changes that occurred in different regions (Fig. 12).
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Figure 12. Comparison
of regional and global
temperature
reconstructions.
Derived from Luckman
and Wilson 2005,
Salzer and Kipfmueller
2005 (both based on
tree-ring records) and
Mann et al. 2008
(based on a multi-proxy
reconstruction).
Tree-ring records from throughout the western U.S. show natural variation in temperature and
precipitation/available moisture during the past two millennia, some of which are synchronous across
large areas of the four climate regions in this study, while others are more representative of local
phenomena (Figs. 12-13, Pederson et al. 2006, Cook et al. 2004). These records show that decadal and
multi-decadal fluctuations in precipitation are a defining characteristic of climate during the past
millennia, and exert important controls on ecosystem processes and species distributions (e.g. Pederson
et al. 2006, Cook et al. 2004, 2007). Regionally synchronous wet and dry intervals have been linked to
low-frequency variations and state changes in sea surface temperature and pressure anomalies in both
the Atlantic and Pacific Oceans which are discussed in more detail later (McCabe et al. 2004, Gray et al.
2003, Cayan et al. 1998).
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Figure 13. Long-term aridity changes in
the West as measured by the percent area
affected by drought (PDSIb−1) each year
(B) (redrawn from Cook et al., 2004). The
four most significant (p<0.05) dry and wet
epochs since AD 800 are indicated by
arrows. The 20th century, up through
2003, is highlighted by the yellow box. The
average drought area during that time,
and that for the AD 900–1300 interval, are
indicated by the thick blue and red lines,
respectively. The difference between
these two means is highly significant
(p<0.001). Source: Cook et al. 2004, 2007
copyright reprint permission needed
Two major, well-documented examples of centennial-scale climate change over the past two millennia
are the Medieval Climate Anomaly (MCA; 950-1250 AD) and the Little Ice Age (LIA; 1400-1700 AD). As
indicated by Figures 13-14 (Cook et al. 2007, Mann et al. 2009) a number of anomalous warm (dry) and
cool (wet) periods exist during this time frame (see Biondi et al. 1999, Meko et al. 2007, Salzer and
Kipfmuller 2008, and Cook et al. 2007), and result in extensive hydrological and ecological impacts.
During the MCA, warm temperature anomalies persisted across western North America (Fig. 13). While
the general climatic conditions during this time were defined by warmer, drier conditions, the local
effects were highly spatially variable. Elevated aridity and “mega-droughts” (i.e. droughts of ~50 years
duration or greater) were commonplace across the western U.S., with more land area synchronously
experiencing sustained drought conditions as compared to the LIA and majority of the 20th century (Fig.
13, Cook et al. 2007).
Data from a number of sites suggest regionally synchronous drought events occurring regularly during
the Medieval Climate Anomaly, with durations and extents unmatched in the late Holocene (Cook et al.
2007, 2004). Glacial retreat in mountainous locations occurred in Colorado, Wyoming, Montana, and
the Cascades, with substantial reductions in streamflow (e.g. Meko et al. 2007, Gray et al. 2003) and lake
levels throughout the study area (Millspaugh et al. 2000, Brunelle and Whitlock 2003). However, spatial
and temporal variations in the generally warm dry conditions were widespread, as the CR-GYA and parts
of the SR experienced increased summer moisture (Davis 1994, Whitlock and Bartlein 1993). Upper
treelines in some areas advanced upwards in elevation and areas now covered by krummholz were
occupied by arborescent trees (Graumlich et al. 2005, Whitlock et al. 2002, Fall 1997, Rochefort et al.
1994, Winter 1984, Kearney and Luckman 1983). Additionally, alpine larch expanded 90 kilometers
north of its current range (Reasoner and Huber 1999, Reasoner and Hickman 1989) ca. 950 1100 BP.
The mechanisms and drivers leading to the MCA are still debated but there is increasing evidence that
low-frequency variation in ocean-atmosphere interactions were an important factor contributing to
MWP anomalies (Fig. 14, Mann et. al 2009).
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Fig. 14. Decadal surface temperature
reconstructions. Surface temperature
reconstructions have been averaged over (A) the
entire Northern Hemisphere (NH), (B) North
Atlantic AMO region [sea surface temperature
(SST) averaged over the North Atlantic ocean as
defined by Kerr (1984)], (C) North Pacific PDO
(Pacific Decadal Oscillation) region (SST averaged
over the central North Pacific region 22.5°N–
57.5°N, 152.5°E–132.5°W as defined by (Mantua et
al. 1997) and (D) Niño3 region (2.5°S–2.5°N,
92.5°W–147.5°W). Shading indicates 95%
confidence intervals, based on uncertainty
estimates discussed in the text. The intervals best
defining the MCA and LIA based on the NH
hemispheric mean series are shown by red and
blue boxes, respectively. For comparison, results
are also shown for parallel ("screened")
reconstructions that are based on a subset of the
proxy data that pass screening for a local
temperature signal [see Mann et al. 2008 for
details]. The Northern Hemisphere mean Errors in
Variables (EIV) reconstruction is also shown for
comparison. Source: Mann et al. 2009 Science
In contrast to the MCA, the Little Ice Age (LIA) was a period of anomalous Northern Hemisphere cooling,
where mountain glaciers throughout western U.S. expanded, many reaching their Holocene maximums
(e.g. Pederson et al. 2004, Luckman 2000, Watson and Luckman 2006). While temperatures across the
study area were persistently cooler than long-term average (> -1˚ C) some data suggests that the
magnitude of the cooling decreased with latitude (Whitlock et al. 2002). In the Northern Rockies where
the most pronounced cooling occurred, conditions in the late LIA may have approached those of the late
Pleistocene (Pederson et al. 2007, Luckman 2000, Clark and Gillespie 1997). Both tree-ring and glacier
data indicate sustained cool summer conditions and increased winter precipitation across the Northern
Rockies resulted in the major LIA glacier advance documented across the NR (Watson and Luckman
2004, Pederson et al. 2004). The Southern and Central U.S. Rockies did not see advances of this
magnitude, demonstrating the spatial variability in LIA climatic anomalies (Clark and Gillespie 1997). As
with the MCA, the drivers and mechanisms that influenced cooler conditions during the LIA are not well
understood. Decreased sunspot activity during a period called the Maunder Minimum led to decreased
incoming solar radiation from ca. 1645-1715 AD and is considered one of the main variables explaining
cooler conditions for this interval of the LIA (Eddy 1976, Luckman and Wilson 2005).
Case studies from the U.S. Rockies and Upper Columbia Basin
The following case studies from each of the four climate regions highlight examples of biophysical and
biotic response to climatic changes over the past two millennia and provide clues to the timing and
extent of future biotic changes.
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Drought variability and ecosystem dynamics in Glacier National Park
An examination of records showing past hydroclimatic changes, glacier dynamics, and fire activity in
Glacier and Waterton National Parks provide an example of how interdecadal and long-term changes in
climate interact to alter important ecosystem processes in the Northern Rockies (Figure 15).
Specifically, long-term changes in regional temperature (e.g., the relatively cool conditions of the LIA vs.
the warm temperatures in the 1350s-1450s and especially over the latter half of the past century Fig.
12) combined with persistent shifts in summer and winter moisture regimes over decadal to
multidecadal timescales have a particularly strong impact on fire regimes in the GNP region (Figure
15a,b,c). Regional records spanning the past three centuries show that periods from the 1780s to the
1840s and the 1940s to the 1980s, for example, had generally cool and wet summers coupled with high
winter snowpack resulting in extended (> 20 yr) burn regimes characterized by small, infrequent fires
with relatively little area burned. Conversely, decadal and longer couplings of low snowpack and warm
dry summers resulted in burn regimes characterized by frequent fires and large total area burned (e.g.
1910 to 1940, 1980s to present).
Figure 15. Relationship between Glacier
National Park historic summer drought,
inferred winter snowpack (May 1st SWE), and
ecosystem processes (i.e. fire area burned and
glacial recession) back to A.D. 1700. (a)
Instrumental and reconstructed summer
drought (MSD) for Glacier National Park
normalized by converting to units of standard
deviation and smoothed using a 5yr running
mean. (b) Measured Spring snowpack (May 1st
SWE) anomalies (1922-present) and average
annual instrumental and reconstructed PDO
anomalies (1700-2000). Each time series was
normalized and smoothed using a 5 yr running
mean to highlight decadal variability. For ease
of comparison the instrumental and
reconstructed PDO index was inverted due to
the strong negative relationship between PDO
anomalies and May 1st snowpack. (c) Fire area
burned timeline for the Glacier National Park
region spanning A.D. 1700-present. Timeline is
presented with maps exemplifying fire activity
using decadal snapshots in time over periods
corresponding to interesting winter and
summer precipitation regimes. (d) Maps
showing the decrease in area of the Sperry
Glacier at critical points in time from 1850 to
1993. The retreat patterns of the Sperry
Glacier are representative of other regional
patterns of recession for glaciers sensitive to
regional climate variability. Source: Pederson
et al. 2006, copyright reprint permission
requested
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Similarly, long-duration summer and winter temperature and moisture anomalies drive glacial dynamics
in the Northern Rockies (Watson and Luckman 2004, Pederson et al. 2004). Prior to the height of the LIA
(~1850), four centuries of generally cool summer conditions prevailed (from ~1450-1850, Fig. 12).
Ultimately, what drove the glaciers to reach their greatest extent since the Last Glacial Maximum was
approximately 70 years of cool wet summers coupled with generally high snowpack conditions between
1770 and 1840 AD (Watson and Luckman 2004, Pederson et al. 2004). Over subsequent periods when
summer drought and snowpack were generally in opposing phases (e.g. 1850-1910) and summer
temperatures remained relatively cool (Fig 15), glaciers experienced moderate retreat rates (1-7 m/yr).
During the period from 1917 to 1941, however, sustained low-snowpack, extreme summer drought
conditions, and high summer temperatures drove rapid glacial retreat. The Sperry Glacier, for example,
retreated at 15-22 m/yr and lost approximately 68% of its area (Figure 15d). Other glaciers such as the
Jackson and Agassiz Glaciers at times retreated at rates  100 m/yr (Carrara and McGimsey 1981, Key et
al. 2002, Pederson et al. 2004). Climatic conditions over the middle of the 20th century (1945-1977)
became generally favorable (i.e. level to cooling summer temperatures, high snowpack w/variable
summer drought conditions) for stabilization and even accumulation of mass at some glaciers (Figure
15a,b). Since the late 1970s, however, exceptionally high summer temperatures (Pederson 2009)
combined with low winter snowpack and multiple periods of severe and sustained summer drought
conditions have resulted in a continuation of rapid glacier retreat. These records representing diverse
settings throughout the U.S. Rockies provide further support that biophysical and ecosystem processes
are strongly influenced by moisture and temperature variability at decadal and longer timescales.
As demonstrated by these records for Glacier National Park and surrounding regions, biophysical and
ecosystem processes of the Northern Rockies are strongly influenced by moisture and temperature
variability at decadal and longer timescales. This relationship between biophysical and ecosystem
processes with non-stationary hydroclimatic changes that also exhibit regime-like behavior at decadal
scales pose challenges to management and sustainability in three key ways. First, long-duration proxy
reconstructions call into question the conventional strategy of defining reference conditions or
management targets based solely on short duration (< 100 yr) observational records. For example, the
use of 30 yr climatology for allocation of natural resources, and development of resource management
goals is flawed as it is probable that any 30 yr climatic mean may only capture a single mode of climate
variability (i.e. an extended regime of wet or dry conditions). Second, step-like changes from one
climate regime to the next can onset rapidly, and have prolonged impacts on ecosystem processes.
These persistent and frequent shifts in the availability of natural resources and services may be
misinterpreted as resulting from management activities. Likewise such climate-related shifts may
amplify or dampen the effects of management activities. Lastly, decadal and longer persistence of
either deficits or abundances of climate-related biophysical processes can lead to management policies
and economic strategies that, while appropriate during the current regime, may not be robust under
subsequent climates. Overall, greater awareness of regime-like behavior must shape our expectations
for ecosystem services and, in the face of global change, the institutions that manage these resources.
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Upper Columbia Basin
Millennial-scale climate variation and fire-related sedimentation in central Idaho and northern
Yellowstone.
An examination of fire-related sediment deposits in central Idaho illustrate millennial-to-centennial scale
climate variation and its control on fire regimes and attendant sedimentation. In central Idaho, Pierce et
al. (2004) dated sediment deposits in alluvial fans to reconstruct erosion events related to fires in dry
forests dominated by Ponderosa pine (Pinus ponderosa) and frequent, low-severity fires. They found
that small sedimentation events occurred more frequently during the Holocene, especially during the
Little Ice Age (LIA) and were associated with frequent fires of low to moderate severity. Large
sedimentation events were associated with prolonged periods of drought and severe fires that were
most pronounced during the Medieval Climate Anomaly (MCA) (Pierce et al. 2004, Fig. 16).
Figure 16. Probability distributions were smoothed using a 100-yr running mean to reduce the influence of shortperiod variations in atmospheric 14C, but retain major probability peaks representing the most probable age
ranges. Calibrated years before present (cal. yr bp) are equal to the number of calendar years before ad 1950. The
general trend of decreasing probability over time, and overall lower probability values before 4,000 cal. yr bp
reflect lesser exposure and preservation of older deposits. a, SFP 'small events' (blue line) in Idaho are thin
deposits probably related to low- or moderate-severity burns; note correspondence of peaks with minima in YNP
fire-related sedimentation, and major peak during the 'Little Ice Age' (LIA), 650–50 cal. yr bp. (Fewer nearsurface deposits in the 400 cal. yr bp–present range were selected for dating because of bioturbation and large
uncertainties in 14C calibration.) b, SFP 'large events' are major debris-flow deposits probably related to severe
fires; note correlation with the YNP record, and prominent peak in large-event probability during the 'Medieval
Climatic Anomaly' (MCA), 1,050–650 cal. yr bp. Source: Pierce et al. 2004 copyright permission needed
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These results were then compared with a similar record from northern Yellowstone National Park where
mixed conifer forests are associated with infrequent, stand-replacing fires (Meyer et al. 1995). Peaks in
sedimentation events there occurred during the Little Ice Age (LIA) when cooler and wetter conditions
(although highly variable) led to infrequent high severity fires in northern Yellowstone. In central Idaho,
in contrast, cooler wetter conditions during the LIA are inferred to have maintained high-canopy
moisture that inhibited canopy fires and facilitated understory grass growth and fuels linked to lowseverity fires (Pierce et al. 2004). Patterns of fire were synchronous between central Idaho and
northern Yellowstone during the warmer and drier MCA. In central Idaho, longer intervals of warm and
dry conditions allowed for sufficient drying of large fuels to increase the frequency of large, standreplacing canopy fires and large fire-related erosion events. Hence, strong climatic controls changed a
fuel-limited, infrequent, low-severity fire regime to a fuel-rich (large fuels) high-severity stand-replacing
fire regime. Large, stand replacing fires also increased in mixed conifer forests of northern Yellowstone
during the MCA (Fig. 16). In northern Yellowstone, large pulses of sedimentation associated with fire
were found to coincide with the Medieval Climate Anomaly (ca. AD 1150) and Meyer et al. infer that this
activity resulted from increased intensity and interannual variability of summer precipitation during this
period. Meyer et al. (1995) also found a correlation between millennial and centennial-scale climatic
variability and periodic and alternating fire-related debris flow and sedimentation events (300-450 yr
period) and terrace development (1300 yr period).
These results were then compared with a similar record from northern Yellowstone where lodgepole
pine (Pinus contorta) forests are associated with infrequent, stand-replacing fires (Meyer et al. 1995).
Peaks in sedimentation events contrasted at the two sites during the Little Ice Age (LIA) when cooler and
wetter conditions (although highly variable) appears to have reduced the frequency of high severity fires
in northern Yellowstone and to have maintained conditions conducive to infrequent, low-severity fires
in central Idaho (Fig. 16). In central Idaho, cooler and wetter conditions during the LIA are inferred to
have maintained high-canopy moisture that inhibited canopy fires and facilitated understory grass
growth and fuels linked to low-severity fires (Pierce et al. 2004). Patterns of fire were more
synchronous in central Idaho and northern Yellowstone during warmer and drier conditions during the
MCA. In central Idaho, longer intervals of warm and dry conditions appear to have allowed for sufficient
drying of large fuels to increase the frequency of large, stand-replacing canopy fires and attendant
increase in hillside erosion and large debris-flow events in Idaho, overriding what is typically a fuellimited system to the extent that a switch from infrequent, low-severity fires to high-severity standreplacing fires occurred. Large, stand replacing fires also increased in lodgepole pine forests of northern
Yellowstone during the MCA (Fig. 16). In northern Yellowstone, large pulses of sedimentation
associated with fire were found to coincide with the Medieval Climate Anomaly (ca. AD 1150) and
Meyer et al. infer that this activity resulted from increased intensity and interannual variability of
summer precipitation during this period. Meyer et al. (1995) also found a correlation between millennial
and centennial-scale climatic variability and periodic and alternating fire-related debris flow and
sedimentation events (300-450 yr period) and terrace development (1300 yr period).
Paleoenvironmental investigations like these provide further evidence that millennial and centennial
scale variations in temperature and precipitation have influenced both biophysical conditions as well as
important ecosystem processes, such as fire across the UCB and CR-GYA and illustrate the complex
linkages between climatic variability, ecosystem processes and biophysical conditions. The results of
these studies and similar research (Whitlock et al. 2003) are important because they suggest that,
ultimately, natural climatic variability acts as a primary controls on ecosystem processes such as
disturbance and have important implications for management. Efforts to manage fuels in different
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stand types to restore specific fire regimes may be trumped by future climate variations and actively
managing for stand conditions that supported what are considered historical (20th century) structural
and disturbance characteristics may have limited impacts under anomalous climatic conditions. The
examples from central Idaho and northern Yellowstone illustrate how climatic variability, such as the
long-term droughts occurring during the MCA, may override efforts to manage fuels and maintain or
restore disturbance processes that support a desired stand condition.
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A long-term reconstruction of changing distributions of Utah Juniper (Juniperus osteosperma) in the CRGYA provides an example of how climate variation, complex topography and spatial distribution of
suitable habitat and biotic factors interact to govern plant invasions and provides evidence contrasting
the popular idea that plant invasions are characterized by a steady and continuous march across
landscapes in response to changing climatic conditions. The distribution of Utah juniper in the
mountains of Wyoming, southern Montana, and Utah during the late Holocene has been tracked from
radiocarbon-dating fossilized woodrat middens ( Lyford et al. 2003) During a dry period in the midHolocene (ca. 7500 – 5400 BP), Utah juniper migrated into the Central Rockies from the south via a
series of long distance dispersal events. Further range expansion and backfilling of suitable habitat was
stalled during a wet period from 5400 – 2800 BP (Lyford et al. 2003). In response to warmer and drier
conditions that developed after 2800 years BP, Utah juniper populations rapidly expanded within the
Bighorn Basin, Utah juniper establishment was particularly high from 2800 to 1000 years BP. The
notable absence of significant Utah juniper establishment and expansion during the MCA suggests that
long-term climate variations determine distributions of species with centennial scale life expectancies.
In the case of the Utah juniper, Lyford et al. (2003) note that establishment rates are significantly more
affected by adverse climatic conditions than individual or population survival where they are already
established. This could, in part, explain the tendency for Utah juniper populations to remain static
instead of contracting during the Holocene.
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In general, Utah juniper’s range expanded during periods characterized by warmer, drier conditions with
expansion and establishment cessation during cool, wet periods (Lyford et al. 2003). Thus, the migration
and establishment of Utah juniper into the Central Rockies was at least in part controlled by millennial
scale climatic variations within the Holocene. Although Utah juniper’s distribution is severely limited by
cool temperatures and high precipitation in higher elevations of the Central Rockies, Utah juniper
inhabits only a fraction of the suitable climate space in the region (Fig. 17). Within its suitable climate
space, Utah juniper distribution is limited by available suitable substrate as present distributions cover
over ninety percent of the substrate deemed highly suitable for Utah juniper survival in the region.
Central U.S. Rockies and GYA
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Figure 17. Suitable habitats for Utah juniper in Wyoming
and adjacent Montana. Areas shown in black indicate
extremely suitable habitats, while gray areas indicate
moderately to highly suitable habitats. (a) Climate (ratio
of growing-season precipitation to growing-season
temperature. (b) Substrate (including soil, bedrock, and
surficial-material type. (c) Climate and substrate
combined. (d) Modern distribution of Utah juniper
(Knight et al. 1987, Driese et al. 1997; available online
[see footnote 3]). Note that climate (a) overpredicts the
distribution (d), which is strongly constrained by
substrate variables (b), and that favorable habitat is
patchily distributed (c). Source: Lyford et al. 2003,
copyright reprint permission needed
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Tree-ring records from Wyoming’s Green River basin provide a reconstruction of drought conditions for
the last millennia and reveals how natural variations in dry and wet periods are a defining characteristic
of the CR-GYA. Tree rings were used to develop a 1,100-year record of the Palmer Drought Severity
Index, a measure of drought incorporating both precipitation and temperature trends, southwestern
Wyoming (Cook et al. 2004). This record is typical of many areas of the CR-GYA showing above average
precipitation in the early 20th century (Woodhouse et al. 2006, Gray et al. 2004, 2007, Meko et al. 2007)
and also shows the potential for severe, sustained droughts far outside the range of 20th century
observations including several multi-decadal year droughts prior to 1300 A.D. (Fig. 18).
The case of the Utah juniper shows how climatic controls can influence species distribution, migration,
and establishment in the Central Rockies within the context of millennial-scale climate change and
highlights the importance of recognizing other environmental factors affecting the distribution of
species. While suitable climate can allow for species to become established, a number of other factors
such as substrate, dispersal and competition influence how successfully species can disperse to and
colonize suitable habitat. This provides an important example for considering how ecosystems and
species will respond to changes in the spatial distribution of suitable habitat with changing climatic
conditions. As illustrated by changes in the distribution of Utah juniper, individual species responses to
environmental changes are spatially variable and contingent on a number of factors in addition to
suitable climatic conditions. In this case, landscape structure and climate variability play key roles in
governing the pattern and pace of natural invasions and will be important variables to consider when
anticipating future changes in the distribution of plant species. The high temporal and spatial precision
provided by this study illustrates the fact that vegetation response to future conditions will be more
nuanced than a steady march to newly suitable habitats and better characterized by episodic longdistance colonization events, expansion and backfilling (Lyford et al. 2003). This study suggests that
models predicting plant invasions based on climate model projections may be oversimplified and
encourages a more focused examination of how species dispersal will interact with the spatial
distribution of suitable habitat and climate variability to govern future invasions.
Natural variability in precipitation in Wyoming’s Green River Basin during the last 1100 years
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Figure 18. 1,100 years of drought history in
the Green River Basin region of southwest
Wyoming, as reconstructed from tree rings
(Cook et al. 2004). The plot shows values for
the Palmer Drought Severity Index (see also
PDSI figure on page 10. Positive values (blue)
of the index represent wet conditions,
negative values (red) indicate drought. Values
are plotted so that each point on the graph
represents mean conditions over a 25-year
period. Source: Gray et al. 2009, copyright
reprint permission needed
The drought reconstruction for the Green River Basin suggests that some of the most severe droughts in
recent times (1930s and 1950s) were relatively minor events compared to many dry periods during the
past millennia and that the latter half of the 20th century was relatively wet with no prolonged droughts
(Fig. 18). This study shows conditions that characterize the long-term climate history of the CR-GYA,
providing strong evidence that drought periods have occurred through the past millennia and are a
natural feature of the regional climate. Hence, using the past century as a reference for climate
conditions would lead one to conclude that the area is wet and relatively free of drought, highlighting
the fact that longer-term records are critical for understanding natural variability of climate conditions in
the CR-GYA and the western U.S.
Southern U.S. Rockies
Changing plant distributions in response to climatic variability and in northeastern Utah
In the Southern Rockies, the history of pinyon pine (Pinus edulis Engelm.) distribution in the Dutch John
Mountains (DJM) of northeastern Utah are strongly controlled by variations in multi-decadal
precipitation patterns (Gray et al. 2006). Gray et al. (2006) used woodrat midden and tree-ring data to
track the spatial and temporal patterns of pinyon pine distribution throughout the Holocene. The DJM
population of pinyon pine is an isolated northern outpost of the pinyon pine that established ca. AD
1246. Similar to Utah juniper in the Central Rockies, the pinyon pine probably reached the DJM via longdistance dispersal from the Colorado Plateau (Jackson et al. 2005), during the transition from the
warmer, drier MWP to the cooler, wetter LIA (Fig. 19). Changes in the distribution of the DJM pinyon
pine population were episodic and did not respond smoothly to centennial scale climate warming during
the Holocene but instead was more closely related to episodic multdecadal and decadal scale variability
in moisture (Gray et al. 2006). DJM pinyon pine expansion was stalled in the late 1200s and significant
recruitment did not resume until the 14th century pluvial when regionally mesic conditions promoted
establishment within the DJM (Fig. 19).
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Figure 19. (a) Percentages of pinyon-type pollen (black
vertical bars) and presence (solid circles) or absence
(open circles) of pinyon pine macrofossils from 12 000 yr
of woodrat (Neotoma) midden records collected at Dutch
John Mountain (DJM). (b) Profile map of the DJM study
area showing locations of midden sites (open circles) and
sampling units used in the tree-ring age studies (shaded
polygons). Each midden site represents a cave or rock
overhang where one or more of the 60 middens were
collected. The estimated establishment dates (year AD)
are based on the average age of the four oldest pinyon
pines found in each sampling unit. (c) Ages for the oldest
pinyon tree on DJM and the four oldest pinyons in each
of the eight sampling units (black dots) plotted against
reconstructed annual (gray line) and 30-yr smoothed
(black line) precipitation values for the Uinta Basin
Region. (d) Percentage of the western United States
experiencing drought conditions over the past 1200 yr as
reconstructed from a large tree-ring network (Cook et al.
2004). Data are plotted as a 50-yr moving average. The
horizontal line at 37% (dark gray) shows the average or
background level of drought through time. Significant
multidecadal dry and wet periods identified by Cook et al.
(2004) are shaded black and gray, respectively. Source:
Gray et al. 2006, copyright reprint permission needed
The case of the pinyon pine in the DJM demonstrates the importance of episodic, multidecadal climatic
variation in controlling rates of ecologic change in the southern U.S. Rockies over the past millennia.
Records suggest that the development of pinyon population at DJM was not a steady wave-like
movement associated with improving climate conditions but a markedly episodic invasion regulated by
natural fluctuations in precipitation (Gray et al. 2006). The multi-decadal episodic variations in
precipitation patterns over the region significantly impacted and altered rates of ecologic change and
vegetation cover transition over the study area and show that the cumulative effects of shorter multidecadal episodic climatic variation superimposed on long term (centennial to millennial scale) climate
change can significantly affect species’ distributions and patterns of migration and establishment.
As with previous studies of plant invasions at longer time scales (Lyford et al. 2003) this study suggests
that climatic variation can amplify and dampen the probability of survival and reproduction after species
colonize new areas and can also change the density and distribution of favorable habitats across the
landscape and influence competitive interactions and disturbance processes (Gray et al. 2006). This
study demonstrates that decadal to multi-decadal climate variability influences important processes
such as disturbance, dispersal, recruitment, mortality and survival in ways that result in range
expansions characterized by episodic invasions (Gray et al. 2006). For example, even the different
locations within a region may experienced similar changes in precipitation or temperature over a
particular time period differences in the frequency and amplitude of variability in climate conditions
could easily produce different disturbance dynamics and different end states. It is often inferred that
vegetation responds to changing climates with a steady wave-like movement to new locations offering
suitable climate yet the DJM pinyon pine example reveals that species response to changing conditions
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is influenced by other important factors such as characteristics of long-distance dispersal, landscape
structure and the variability of climate conditions at different scales (Gray et al. 2006). Anticipating
ecological response to climate change will require a better understanding of how natural climate
variability regulates species migrations and invasions at smaller time scales.
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Paleoenvironmental records provide us with critical information on past climates and the ecological
response to natural climatic variability and context for anticipating future change. Ecosystems have
responded continuously to climatic variation during the Holocene. At large scales, climate ultimately
governs the distribution of vegetation across the landscape, and as evidenced by paleoenvironmental
records from throughout the four climate regions, climate variability at different spatiotemporal scales
regulates the distribution of vegetation across the landscape, although climate is not the sole factor
governing species distributions. The importance of climate is evident in pollen records documenting
movements of dominant tree species upslope, downslope and to new locations associated with major
climate changes during the Holocene. And while these longer-term records reveal fundamental shifts in
vegetation at millennial and centennial scales, high-resolution data show that the pace and magnitude
at which climatic variation occurs is important. Species respond to changing conditions by dispersing
and colonizing newly suitable sites unless variation in climate is too rapid or extreme. Importantly, plant
invasions and migrations are often better characterized as episodic movements rather than continuous
steady expansions (Gray et al. 2006, Lyford et al. 2003). Predicting plant response to climate change in
the future will require a more complex consideration of how climate variability and landscape structure
interact with species attributes to pace and govern future redistributions (Jackson et al. 2009).
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What does the long-term paleoenvironmental record tell us about past climate change?
Additionally, proxy records from throughout the four climate regions demonstrate that biophysical and
ecosystem processes are strongly influenced by climate variability at decadal and longer timescales. This
is important because management planning is often based on climate “regimes” that are defined on
decadal scales. For example, fuel management goals are often developed from patterns of fire
associated with different forest stand types over the past several decades even though
paleoenvironmental data show that this will only capture a single mode of climate variability (i.e., an
extended regime of warm and dry or cool and wet conditions, e.g., Pierce et al. 2004, Whitlock et al.
2003). Threshold like changes from one climate regime to the next can occur quickly and, thus, longer
term consideration of the spectrum of conditions that occurred on longer time scales should be used as
a context for resource management. Lastly, a consideration of may not be robust under subsequent
climates. Overall, greater awareness of regime-like behavior must shape our expectations for
ecosystem services and, in the face of global change, paleoenvironmental history relevant to the climate
regions in this study indicates:



Using the past century as a reference for baseline climate condition does not capture important
scales of natural climate variability and is an inadequate scale of reference for considering
future climate change.
Climatic forcings are manifested climatically in different ways in different places as a result of
interactions among controls and large scale climate events can be synchronous across regions
but not always uniform
Many terrestrial ecosystems are young, dynamic and historically contingent, and as a result,
restoration to historical baselines may be impossible. 4) Rapid climate transitions and associated
ecosystem transitions have occurred in the past and will likely occur in the future.
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20th Century Climatic Change and the Instrumental Record
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Throughout much of the western U.S., the expression of natural variations in the climate system can
differ greatly across elevational, latitudinal and longitudinal gradients and at different spatial and
temporal scales. Likewise, it is important to keep in mind that the footprint of any broad-scale climatic
changes will vary across finer spatial and temporal scales, and from one area to another. Additionally,
as a direct measurement of climate conditions for the past century, instrumental observations provide
the most accurate and reliable data available. Yet these records also include local biases (e.g. influence
of changing slope or aspect), along with the signature of finer-scale processes. Thus, individual
instrumental records should not be interpreted as representing the region as a whole, but rather as an
indication of local conditions or, at most, conditions at similar locations (e.g. comparable elevation and
vegetation cover) across a wider area.
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. 20, 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 individual sites in the northwestern U.S., CIG 2010). 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 change in the early 20th century, the rate of increase for much of the
study area doubled from the mid-20th century to present, with the largest part of this rapid warming
occurring since 1975.
Figure 20. 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
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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
Temperature increases are more pronounced during the cool season (Hamlet and Lettenmaier 2007). In
the northern U.S. Rockies annual rates of increase are roughly 2-3 times that of the global average (Vose
et al. 2005, Bonfils et al. 2008, Pederson et al. 2010, Hall and Fagre 2003), a pattern that is evident at
northern latitudes and higher elevation sites throughout the western U.S. (Diaz and Eischeid 2007,
USGCRP 2005). Additionally, nighttime minimum temperatures are increasing faster than daytime
maximums resulting in a decreased Diurnal Temperature Range (DTR) (Pederson et al. submitted). This
decreased DTR has important implications for species like the mountain pine beetle (Dendroctonus
ponderosae), whose population dynamics are governed by minimum temperatures (Carroll et al. 2004).
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) find that recent warming observed in the mountainous areas across the
west cannot be explained by natural forcings (e.g. solar, volcanic, ocean-atmosphere interactions) alone,
rather a major portion of recently observed temperature increase is attributable to human influenced
changes in greenhouse gas (GHG) and aerosol concentrations.
Precipitation
Trends in precipitation for the study area are far less clear. Instrumental data for much of the
northwestern U.S. show modest increases in precipitation during the past century (Fig. 21, 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 during the past
century impacted large areas of the study region. Though 20th century droughts had substantial
socioeconomic and ecosystem impacts, there is ample evidence they were not as severe, in terms of
duration and magnitude, as a number of drought events that occurred during the past millennium (Cook
et al. 2007, 2004; Meko et al. 2007). 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 and central 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). Research suggests that
climatic conditions that influenced the nature and location of these droughts is likely linked to lowfrequency oscillations in ocean-atmospheric interactions (McCabe et al. 2004, Gray et al. 2007, Hidalgo
2004, Graumlich et al. 2003), with evidence for substantial surface-feedbacks during the 1930s “Dust
Bowl” drought (Cook et al. 2009).
Surface Hydrology
Generally speaking, since 1950 snowmelt and peak runoff has tended to occur earlier, and river flows in
many locations are decreasing during late summer months (Fig. 22, Pederson et al. 2010, 2009, Mote et
al. 2006, Barnett et al. 2008, Stewart et al. 2004, McCabe and Clarke 2005). Recent impacts on
snowpack and surface hydrology are strongly associated with more precipitation falling as rain than
snow, and warming temperatures driving the earlier snowmelt, and snowmelt driven runoff (Pederson
et al. 2010, 2009, Mote et al. 2006, Barnett et al. 2008, Stewart et al. 2004, McCabe and Clarke 2005).
The earlier snowmelt driven runoff, consequently results in reduced surface storage of moisture,
resulting in increasingly low baseflows during the dry summer months (Luce and Holden 2009).
Examination of both paleoclimate data and instrumental records (e.g. stream gages, snow course
records, valley meteorological stations) suggests, however, that the total amount cool season
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precipitation received across a particular region is more strongly associated with natural multi-decadal,
decadal and inter- and intra-annual variability in ocean-atmosphere conditions (e.g. PDO, AMO, ENSO).
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Figure 21. 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
Despite modest increases in precipitation in parts
of the central and northern study area (modest
declines for parts of the southern U.S. Rockies),
significant declines in snowpack are evident
throughout much of the northwestern U.S. This
decline is especially prevalent in the northern U.S.
Rockies and parts of the Upper Columbia Basin
(Hamlet and Lettenmaier 2007, Pederson et al.
2004, 2010, Selkowitz et al. 2002, Mote et al. 2005,
Mote et al. 2008). Warmer temperatures and
declining snowpack have, in turn, contributed to
significant declines in the region’s glaciers. In
Glacier National Park, glaciers have decreased in
area by over 60% since 1900 (Hall and Fagre 2003,
Key et al. 2002, Fountain et al. 2007); and of the
150 glaciers and snow and ice-fields present in
1910, only 26 remain (Fagre 2008, Pederson et al
2010). Changes in glacier mass and area are emblematic of the recent changes in surface hydrology
across the western U.S., with the recent substantial declines attributable to human-caused climate
change related to increases in greenhouse gas emissions and aerosols (Barnett et al. 2008, Bonfils et al.
2008, Pierce et al. 2008, Hidalgo et al. 2009). 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 over the coming decades (Hoerling and Eischeid 2007).
Figure 22. Trends in winter-mean wet-day
minimum temperatures for 1949–2004.
Source: Knowles et al. 2006, copyright
reprint permission needed
Ocean-Atmosphere Interactions
Ocean-atmosphere interactions are
important drivers of inter-annual to
multi-decadal, variability in
temperature and precipitation
(Pederson et al. 2010, 2009, Gray et al.
2007, 2004, 2003, McCabe et al. 2004,
Hidalgo 2004, Cayan 1998, Dettinger et
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al. 2000) but their impacts vary greatly across latitudinal, elevational and longitudinal gradients. The El
Niño-Southern Oscillation (ENSO), and Pacific Decadal Oscillation (PDO) measures of high (2-7 yr period)
and low frequency (20-30 yr period) variability in sea surface temperatures (SSTs) in the equatorial and
North Pacific ocean are, respectively, the predominant sources of inter-annual and inter-decadal
climate variability for much of the study area (Mantua et al. 2002, 1997). 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 warmer 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 (Bjerknes 1986, Cane and Zebiak 1985, Graham
and White 1988, Horel and Wallace 1981, Philander 1990, Chang and Battisti 1998). 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. Warm and cool
phases typically alternate on 2-7 year cycles (Philander 1990, Chang and Battisti 1998). The PDO
exhibits alternate cool and warm phases with a spatial pattern similar to ENSO, but these phases
typically last for 20-30 years (Mantua et al. 1997). During the 20th century several switches occurred
between warm and cool PDO phases and the magnitude of PDO phases increased in the latter half of
the past century (McCabe et al. 2004, Mantua et al. 2002, 1997). The Atlantic Multi-decadal Oscillation
(AMO), representing low frequency (50-80 yr period) oscillations in North Atlantic SSTs have also been
linked to multi-decadal variability in temperature and precipitation in the western U.S. through complex
interactions with PDO and ENSO, although the magnitude of influence of AMO is debated (Kerr 2000,
McCabe et al. 2004, Enfield et al. 2001).
Changes in ENSO and PDO impact precipitation differently across the western U.S., with winter
precipitation in the Upper Columbia Basin and the northern U.S. Rockies being negatively correlated
with warm conditions in the equatorial Pacific (i.e., during El Ninos). Parts of the central U.S. Rockies
and the southern U.S. Rockies tend to be wetter than average during El Ninos, and dry during La Nina
(Mote et al. 2005). Likewise, the northwestern U.S. generally receives more winter precipitation during
the cool-phase PDO, and the southwestern U.S. often receives more precipitation when the PDO is
warm. Further evidence for spatial heterogeneity in ENSO and PDO impacts can be seen in area such as
Greater Yellowstone. Here windward aspects and the high mountains and plateaus to the west tend to
follow more of an Upper Columbia Basin and Northern Rockies pattern, whereas locations to the
leeward and east often behave like the central and southern Rockies (Gray et al. 2004).
Multi-year droughts and extended dry regimes appear to be linked to complex interactions between
PDO, AMO and to a lesser extent, variations in ENSO. 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, Hidalgo 2004, Fye et al. 2003). Drought conditions in the interior
western U.S. are associated with low-frequency variations in PDO and AMO (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
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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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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 and studies indicate that the PDO was in a cool
phase from approximately 1890 to 1925 and 1945 to 1977 (Mantua 1997, 2002). 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”).
Ocean-atmosphere interactions - PDO and ENSO.
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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.
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 (Cayan et al. 1998, Swetnam and Betancourt
1998, Dettinger 1998, Mote et al. 2006)
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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Changes in storm track and circulation patterns
Simulations of 21st century climate suggest a northward movement of the storm-tracks that influence
precipitation over much of the western U.S. (Yin 2005, Lorenz and DeWeaver 2007), and this 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 lead to a longer duration of warm season-like conditions (i.e., predominately warm
and dry) for the Upper Columbia Basin, northern U.S. Rockies and parts of the central U.S. Rockies.
Changes in storm-track position and circulation patterns will be superimposed on a background of
natural variability. Thus, various combinations of ENSO, PDO, etc. and broad-scale trends could lead to
local impacts that vary greatly in their magnitude over time.
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
observed in the past century include: earlier spring bloom and leaf out of plant species, forest infilling at
and near treeline, and increased severity of disturbances such as wildfire and insect outbreaks, all of
which are likely to continue with additional 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, Schwartz et al. 2006). 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, Millar et al. 2009). 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 other variables (e.g., aspect, soils, and micro-topography) that control
treeline position (Graumlich et al. 2005, Lloyd and Graumlich 1997, Bunn et al. 2007, 2005).
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., with the most evident impacts 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. 23-24,
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
diurnal temperature range (DTR) (Pederson et al. 2009). The magnitude of minimum temperature
increases also appears seasonally variable: in area mid-elevation snow telemetry (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. 24). The magnitude
of changes varies locally but there are few exceptions to this general trend in warming.
Figure 23. 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, copyright reprint permission
needed
Temperature trends within the NR generally track Northern Hemisphere trends across temporal scales
(Fig. 23). 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 24. 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:
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 yaxis 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. 25, Pederson et al. submitted) Ecological responses are also evident in
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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 (Peterson and Peterson 2001). The
response is opposite at low elevation sites 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).
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. 25). 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.
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.
25). 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. 25).
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).
Figure 25. Idealized relationship between NRM
snowpack and streamflow anomalies with
associated Pacific SSTs, atmospheric
circulation, and surface feedbacks. Source:
Pederson et al. submitted
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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 (Vose et al. 2005 GRL,
Bonfiles et al. 2008, Mote et al. 2003, 2006). 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 some cases, lower summer flows for
major basins in the climate region (Mote et al 2006, McCabe and Clark 2005, Stewart et al 2004, Hidalgo
et al. 2008).
Precipitation
Records of precipitation for the CR-GYA show highly variable patterns across gradients in elevation,
latitude and longitude. No long-term trends are evident over the past century, although patterns in
inter-annual, decadal and multi-decadal variation are evident in reconstructions of past hydrology from
tree-ring records (Fig. 26, Watson et al. 2009, Gray et al. 2007, 2004, Graumlich et al. 2003). A greater
proportion of precipitation is likely 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 CRGYA, the 1930s and 1950s were significantly drier and the 1940s were wetter than average although
sub-regional 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).
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Figure 26. (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, lessreplicated years. Source: Gray et al. 2007
Quaternary Research, copyright reprint
permission needed
Ocean-atmosphere interactions
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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,
latitude and longitude often result in opposite trends in climatic conditions at sites within the same
region (Gray et al. 2004).
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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).
High elevation snow basins within the western portions of the GYA typically respond to large-scale
climate forcing in a manner 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 warm-phase (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 a variable response to ENSO that depends heavily on the strength of
event, as well as interactions with non-ENSO climate drivers (Gray et al. 2004, McCabe et al. 2007).
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
(McCabe et al. 2007).
Southern U.S. Rockies
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. 27, Dettinger et al. 2005).
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Fig 27. 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
4km-gridded PRISM data. Colorado
River natural flow data from the
Bureau of Reclamation. Source: Ray
et al. 2008. copyright reprint
permission needed
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. 28, 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 (Mote et al. 2003, 2006). 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 28. 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. 29, 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 29. 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.
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 (Fig. 30, Mote et al. 2003, Hamlet et
al. 2005, Mote 2006). Declines in snowpack and SWE are associated with increased temperatures and
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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 30. 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 2010). 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 may already be having dramatic impacts on surface hydrology as they
contribute to earlier melt-off 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), lower summer flows (Barnett et al. 2008). 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 will likely
contribute to increased drought severity and duration, and the potential for more frequent droughts
(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 evaporation and plant water-use that will be driven by rising temperatures
(Gray and McCabe 2010, Hoerling and Eischeid 2007, Seager et al. 2006).
Shifting distributions and new norms
Many parts of the study region are vulnerable to small changes in temperature because the overall
climate is arid to semi-arid to begin with, and what water is available in these areas depend heavily on
the dynamics of mountain snowpack (Gray and Anderson 2010). While many ecosystems are adapted to
natural variations in water availability, a shift in drought frequency and magnitude, or even the
occurrence of an especially severe and prolonged dry event, could result in regional ecosystems
reaching a tipping point whereby a major redistribution of vegetative communities ensues (Fig. 31, Gray
et al. 2006, Jackson et al. 2009). Rapid changes in surface hydrology related to altered snowpack
dynamics could bring similar impacts to aquatic ecosystems that have developed under distinctly
different hydroclimatic regimes (Fig. 31).
Figure 31. 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.
Seemingly small increases in temperature will result in greater evaporative losses from lakes, streams,
wetlands and terrestrial ecosystems, and it is likely that this enhanced evaporation will lead to
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significant ecosystem and water management impacts (Arnell 2009, Gray and Andersen 2010). In the
central and southern portions of the study area, the increases in winter precipitation predicted by some
models are not expected to offset this increased evaporation and transpiration. For example, models
from Gray and McCabe (2010) estimate an average 15-25% decrease in average Yellowstone River flows
given 1.5-3° C temperature increase, and it would require the equivalent of the wettest years in the last
millennium to offset the impacts of increased evapotranspiration on this system (Gray and McCabe
2010). As seen in many recent observational records, seemingly small changes in mean conditions will
lead to an increased frequency of hot weather (relative to historical conditions) and extreme
precipitation events (Karl et al. 2009, IPCC 2007, Groisman et al. 2005, Kunkel et al. 2003, Madsen and
Figdor 2007, Fig. 32). Regarding precipitation in particular, multiple assessments point to a potential
shift in precipitation where storms become more intense, but less frequent (Groisman et al. 2005,
Kunkel et al. 2003, Madsen and Figdor 2007). This, in turn, would increase the number of dry days
between precipitation events, while also altering runoff, infiltration and erosion rates.
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Figure 32. 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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1900
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. One such example of short-term variations in climate would be the unusually cool Northern
Hemisphere winter of 2009-2010. This cooling is likely linked to a well-known natural variation in
climate called the North Atlantic Oscillation (NAO, Hurrell 1995). In short, the NAO is a large-scale
circulation feature that can alternately block or allow cold air from the artic to enter the mid-latitudes of
North America, Europe and Asia. In ’09-10, the NAO was perfectly positioned to spawn cold winters, as
well as strong storms and heavy snowfalls. Interestingly, when considering the global average, the
December 2009-February 2010 period still ranks as the 13th warmest in the past 131 years
(http://www.ncdc.noaa.gov/sotc/). In addition, rates of warming in the Northern Hemisphere have
slowed somewhat in the past decade, even though the 9 of the 10 warmest years on record occurred
during this time (http://www.ncdc.noaa.gov/sotc/). This phenomenon has been largely attributed to a
temporary decrease in water vapor—essentially another type of “greenhouse gas”--contained in the
lower stratosphere (Solomon et al. 2010).
1901
1902
1903
Changes in the amount and spatial distribution of precipitation are still poorly understood (IPCC 2007).
This is largely because precipitation is controlled by complex interactions between global- and
hemispheric-scale circulation features and ocean- and land surface -atmosphere interactions that occur
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across a range of spatial and temporal scales. Moreover, the complex terrain of the study region will
likely alter the impact of any broad-scale shifts in precipitation pattern. Predicting future precipitation is
further complicated by the fact that human activities may also be altering natural controls (e.g., ENSO
and PDO) on atmospheric circulation patterns and stormtracks (Barnett et al. 2008, Bonfils et al. 2008).
While models generally predict increased winter precipitation for parts of the northern Rockies and the
Upper Columbia Basin, with decreases for the southwestern U.S., overall uncertainty remains high (IPCC
2007). In the central Rockies and GYA model results vary widely, and there is very little clear guidance
for how conditions might in coming decades. In addition, all of these areas feature pronounced natural
variability at multi-year to multi-decadal timescales, and this natural variation may alternately mask or
enhance the regional expression of any broad-scale precipitation trends.
1914
Model projections of future conditions
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1920
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
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.
1921
1922
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.
1923
1. Northern U.S. Rockies
1924
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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1931
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.
1932
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1936
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Trends that will likely continue to impact large parts of the study area
Climate conditions


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
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. 33, Hoerling and Eischeid 2007, Seager et al.
2006).
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1944
1945
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1947
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1981
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1988
Fig. 33. 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, copyright reprint
permission needed
Surface Hydrology

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
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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)
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)
Circulation patterns
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1989
1990
1991
1992
1993
1994
1995
1996
1997
1998
1999
2000
2001
2002
2003
2004
2005
2006
2007
2008
2009
2010
2011
2012
2013
2014
2015


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)
Productivity – Phenology


Earlier blooming dates for many plant species (Cayan et al. 2001, Schwartz and
Reiter 2000)
Longer growing season (Bates et al. 2006)
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)
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)
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2016
Planning for the future
2017
2018
2019
2020
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Planning for future conditions that are highly uncertain presents a significant challenge for land
managers. However, techniques developed for business, finance and military applications give us a
roadmap for planning in the face of large uncertainties. “Scenario planning” is one such approach for
preparing for climatic variability and change that uses a combination of scientific input, expert opinion
and forecast data to develop alternative scenarios for the future (Schwartz 1991, van der Heijden 1996).
This contrasts with more traditional attempts at developing precise, quantitative assessments of future
conditions, which are often of limited value for understanding climate change because of compounded
uncertainties. In scenario planning, a suite of alternative scenarios can be used as a starting point for
exploring species or ecosystem vulnerabilities under a wide 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 case, 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. 34). These quadrants each provide a
contrasting scenario or “storyline” 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.
Figure 34. 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, the axes and the resulting storylines may be
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defined quantitatively, or they may be based on qualitative assessments alone. Source: Jackson et al. 2009
Paleontological Society Papers, copyright reprint permission needed
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2100
2101
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The influence of human activities in the climate system, primarily changes in GHG and aerosol
concentrations, will be superimposed on natural drivers of climatic change. As records of the past,
present and model projections for the future suggest, the certainty with which we can expect future
conditions to change varies widely, and have thus far demonstrated that we can be more certain that
temperatures will increase over large spatiotemporal scales than other climatic variables. Investigating
past change also instructs us that, at smaller spatiotemporal scales, changes in future climatic conditions
will likely vary greatly, and especially in mountain environments characterized by high levels of
biophysical and environmental heterogeneity. Temperatures are predicted to increase in the ensuing
decades, meaning much of the western United States will be vulnerable to increased frequency and
duration of drought conditions (Barnett et al. 2008, Seager et al. 2007, Diffenbaugh et al. 2005). Small
increases in temperature will require large increases in precipitation to offset increased
evapotranspiration, especially in settings where increased temperatures significantly alter available
moisture and land-surface-atmospheric feedbacks (Hoerling and Eischeid 2007). High levels of
uncertainty about how ecosystems will respond to changing conditions should not paralyze managers
planning for the future. Scenario planning can be an effective approach for considering a wide range of
possible future conditions that oftentimes indicate similar management actions. In many ways, the
outcomes of scenario planning reinforce fundamental principles of resource management.
Paleoenvironmental records from the past suggest that many ecosystems are resilient to climatic change
across many (millennial to decadal) time scales but also suggest that we should anticipate novel
vegetation assemblages to occur in the future (Whitlock et al. 2003, MacDonald et al. 2008, Williams et
al. 2007, Jackson et al. 2009). Managers should anticipate dynamic and heterogenous redistribution of
vegetation across landscapes in the West and, as a result, might best plan for the future by supporting
this dynamic redistribution through by continuing to employ sound resource management and
conservation practices such as management of areas surrounding protected areas in ways that work in
concert with management objectives. Examples include increasing connectivity, providing buffer habitat
around protected areas and mediating human impacts both within and around public lands.
Management actions that work to accommodate dynamic redistribution of vegetation that will be
difficult to predict may best protect the attributes that are valued on these lands.
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 context to
scenarios. It also helps to determine the likelihood of any of the four quadrants. For example, transition
to “novel ecosystems” is analogous to the transition observed 11,000 yr BP at Yellowstone when open
tundra vegetation was replaced by closed forests (Millspaugh et al. 2000), while the transitions to
“inevitable surprises” are analogous to the late Holocene changes in fire regimes (Whitlock et al. 2003,
Romme and Despain 1989). The greatest value in scenario planning comes from uncovering
vulnerabilities and potential responses, particularly those common to the range of conditions
characterizing a set of scenarios. Hence, despite high levels of uncertainty, scenario planning may reveal
that management response may be similar for a wide range of possible outcomes. Managers can then
move beyond the challenges presented with not knowing how conditions may change to identifying
management responses that address ecological responses to a range of climatic conditions.
Conclusions
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