PART I - DRAFT 03.12.2010
Climate and Ecological Responses to Climate Change in the U.S. Rocky Mountains and Upper
Columbia Basin: a Synthesis of the Best Available Science
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David B. McWethy*, 2Stephen T. Gray, 3Philip E. Higuera, 4Jeremy S. Littell, 5Greg 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. Ecological response to
climatic change is linked to changing conditions at different temporal scales (i.e., millennial, centennial,
decadal, multi-decadal, annual and inter-annual) and the controls on climate vary at these different time
scales. Paleoecological data show that ecosystems in the western U.S. have undergone significant
changes since the last glacial maximum (ca. 20,000 years ago) and the assemblages that we observe
today are relatively recent phenomena (Whitlock et al. 2002, Pederson et al. 2009, Jackson et al. 2009).
Understanding the drivers and rates of climate change and the attendant ecological response to climate
changes at longer time scales is critical for better understanding how ecological communities and
individual species will respond to rates and magnitude of climatic change in the future.
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The purpose of this report is to synthesize important controls and rates of climatic change at different
scales during the past 20,000 yrs, highlight some of the key ecological responses to climatic variation
and report on the most current predictions for the future based on modeling efforts. Of central
importance is a comparison of the drivers and rates of climate change and the ecological response in the
past with recent observations and predictions for the future. This comparison will provide an historical
context for considering how species and ecosystems might respond to current and future changes in
climate conditions. The synthesis highlights key examples of research from each climate region to
illustrate past, present and future changes and is not intended to be an exhaustive review of research.
Importantly, the level of certainty of our understanding of past, present and future changes in different
climatic conditions and the ecological response to these conditions varies greatly and levels of
uncertainty are particularly important to consider when interpreting past records and anticipating future
changes based on model predictions. For example, lake-sediment records typically provide decadal
scale resolution of past change that extend back into the last glacial (tens of thousands of years)
whereas tree ring records provide inter-annual resolution but are limited in providing data beyond the
last two millennia (Table 1). Similarly, paleoclimatological proxy provide relatively consistent records of
past changes in temperature (at certain scales) but are often limited in providing clear and consistent
reconstructions of past variation in precipitation. This variation in levels of certainty influences our
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ability to predict future changes in climate, and consequently, the extent to which we will be able to
predict trends for certain variables (e.g., temperature) are much greater than for others (e.g.,
precipitation).
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The synthesis is organized into past, present and future sections across four climate regions of the
western US: the northern U.S. Rockies (NR), the central U.S. Rockies and the Greater Yellowstone Area
(CR-GYA) and the southern U.S. Rockies (SR) (Fig. 1). The past section reviews coarse (past 20,000 yrs.),
and medium ( past 2,000 yrs.) resolution paleo environmental data; the present section highlights
recent (20th century) high-resolution data from the instrumental record and the future section outlines
predictions for future climate conditions from downscaled model results for the study area. For the past
and recent climate sections the synthesis proceeds by: (1) considering the primary drivers and controls
of climate and identifying specific changes in biophysical conditions and climate as evidenced by data;
(2) highlighting key periods of climate change and associated broad ecological responses; and (3)
discussing implications for managers tasked with maintaining key resources in the face of changing
conditions. The section on future conditions focuses on showing model predictions for a selection of
important climate variables and a discussion of anticipated biophysical and ecological changes
accompanying these predictions. The synthesis concludes with a discussion of challenges in planning for
future conditions where uncertainty is high and provides an example of an approach for dealing with a
wide range of potential future conditions.
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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 scales
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The influence of variation in climatic conditions on ecosystem change varies across temporal and spatial
scales and the ecological response depends on the rate and magnitude of change (Table 1). Climate
controls operating on centennial to millennial time scales shape the broad characteristics of climate and
vegetation. An important control on millennial time scales is the long-term variations in the seasonal
and annual cycle of insolation, because it influences the persistence and strength of storm tracks,
subtropical high-pressure systems, ocean-land temperature gradients, and El Niño-Southern Oscillation
(ENSO) variability that shape vegetation assemblages on shorter time scales. For example, higher-thanpresent summer insolation in the early Holocene (11,000 to 6,000 cal yr BP) supported long-term
changes in vegetation in the northwestern U.S., directly through its influence on higher-than-present
summer temperatures and decreased effective moisture, and indirectly by enhancing atmospheric
circulations systems that suppressed summer precipitation (Whitlock et al. 2008, Anderson et al. 2008)
and promoted xerophytic vegetation (Long et al. 1998, Walsh et al. 2008).
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Table 1. Drivers of climatic variation at different scales and ecological response
Scale
Variation
Driver
Ecological Response
Interannual
Storms, droughts,
ENSO events
Interannual variation,
volcanism
Disturbance mortality, shifts in
abundance and population
dynamics
Decadal/Centennial
Decadal and
centennial
anomalies
Interannual variation,
volcanism, incoming
solar radiation
shifts in relative abundance of
different taxa
Millennial
Deglacial and
postglacial
variations
Ice sheet size, insolation,
ocean-atmospheric-ice
interactions, orbital
harmonics, CO2
variations
Range expansion and
contraction, community
reorganization and
establishment
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Climate variations occur at different temporal scales ranging in length between seasonal to
multimillennial over the past 20,000 yrs (Fig. 2a-d). Shorter climate variations are superimposed on
longer ones, amplifying or dampening their magnitude and each scale of variation influences ecosystem
dynamics to different degrees. Climate variations on time scales of 10,000-100,000 years are attributed
to slowly varying changes in the earth’s orbit (Kutzbach et al. 1998). Superimposed on these long-term
variations are centennial-scale (i.e., 100 yr.) changes in climate. The causes of these variations are still
debated but millennial-scale climate variability during glacial and interglacial periods are thought to be
linked to a number of internal and external forcings including: ocean-atmosphere-ice interactions
(Broecker and Denton 1989), stochastic processes (Hasselmann 1976), variable solar output (Crowley
2000, Bond et al. 2001, Van Geel et al. 1999), nonlinear feedbacks within the climate system (Rind 1999)
orbital cycles (Pestiaux et al. 1988), and volcanism (Viau et al. 2006, Zielinski 1996, Crowley 2000). The
main drivers of climatic conditions for North America for much of the last glacial are thought to be
ocean-ice-atmosphere interactions, insolation and volcanism (Fig. 2a). As the Laurentide and Cordilleran
ice-sheets began to recede and climate amelioration commenced, paleorecords show widespread
reassembly of vegetation throughout the western U.S. (Pederson et al. 2007, Jackson et al. 2009,
Whitlock 2002, Whitlock 1992).
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Figure 2. Climatic variation at millennial, centennial and decadal/interannual time scales, (A) GISP2 temperature
reconstruction with forcing mechanisms thought to influence variation during the glacial period (ca. 49-12 ka BP)
and as ice-sheets in North America began to recede and climate amelioration commenced (ca. 12 ka BP – present);
Gray bands in (A) indicate Younger Dryas and 8.2 ka Holocene transition events; (B) shows detrended variation in
temperature anomalies from multiple pollen reconstructions and the periodicity of variation evident from multipass filter analyses (1100-1200 periodicity indicated by red line, Viau 2006); (C) shows multiproxy temperature
anomalies 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 Warm Period and Little Ice Age from Mann et
al. (2009) but varies by region (MacDonald et al. 2008, Bradley et al. 2003, Carrara 1989). (D) Shows recent
temperature anomalies (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).
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During the last millennium, two examples of centennial scale climatic variation include a period of warm
and dry conditions in the western U.S., known as the Medieval Warm Period, ca. ~900-1300 A.D.(Bradley
et al. 2003), and a period of anomalously cool conditions termed the Little Ice Age (LIA), ca. ~1300-1850
A.D. (Carrara 1989, Fig. 2c). The LIA is defined as a period when glaciers in the Northern Hemisphere
reached their greatest extent during the Holocene (past 10,000 yrs) and includes some of the coldest
periods in the Holocene. Records of past climate suggest that anomalously warm or cool conditions
associated with these periods were not uniformly manifest across North America and some locales show
little indication of these events. Additionally, the magnitude of climate changes during these centennialscale events is less than those on longer time scales, but environmental responses were still substantive
and include changes in glacial extent, disturbance frequency and species distributions.
On shorter time scales, evidence from 20th century instrumental records, and proxy climate data for
several past centuries, indicates that changes at decadal- to multi decadal-scales are a well-defined
component of the climate history of western North America. On decadal to multi decadal time scales,
important modes of climate variability include Pacific Decadal Oscillation (PDO), and the Atlantic
Monthly Oscillation (AMO) and decadal shifts in climate are particularly well expressed in precipitation
reconstructions, and many records indicate decadal-scale droughts or wet periods that surpass events of
the 20th century in magnitude, intensity and duration (Fig. 2d). These events are often regional to
subcontinental in scale and initiated and terminated rapidly (decades to years). The timing of rapid
climate changes (or regime shifts), and their spatial expression suggests that such events are related to
changes in sea-surface temperature and pressure anomalies in both the Atlantic and Pacific Oceans
(e.g., Gray et al. 2003). These decadal shifts in moisture and temperature cause regional ecological
responses. For example, a recurring multi-decadal shift in Pacific sea surface temperatures (e.g. the
Pacific Decadal Oscillation; Mantua et al. 1997) has driven changes in distribution and abundance of fish
populations along the northeast Pacific Coast. Decadal-scale droughts in the semi-arid woodlands of the
American Southwest have resulted in widespread bark beetle outbreaks, forest fires, and tree mortality
(e.g., Allen and Breshears 1998). Likewise, decadal shifts in moisture and temperature have led to rapid
changes in glacier mass balance in both coastal and continental glacier systems (e.g. Bitz and Battisti
1999, Pederson et al. 2004, Watson and Luckman 2004).
Inter-annual climate variability is a shorter, well known mode of climate change largely attributed to the
El Niño Southern Oscillation (Fig. 2d, 3). ENSO is driven by ocean-atmosphere teleconnections related
to patterns of warming (El Niño) and cooling (La Niña) in the central and eastern equatorial Pacific that
take place approximately every 2 to 7 yrs (Chang and Battisti 1998). The strength of these oceanatmosphere teleconnections appears to vary through time and typically results in different regional
moisture conditions. For example, El Niño events have caused the northwestern U.S. to experience
warm dry conditions and the southwestern U.S. to be wet. La Niña, on the other hand, generally
coincided with cool moist conditions throughout the northwestern U.S. while triggering droughts in the
southwestern U.S. This mode of climate variability appears strongest in the southwestern U.S. (e.g.
Swetnam and Betancourt 1998) and is relatively weak and spatially variable in the northwestern U.S.
(e.g., Dettinger et al. 1998).
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Figure 3. Recent drivers of global temperature
changes which are strongly linked to
volcanism, ENSO, solar output and human
contributions to greenhouse gas emissions.
Source: McCarthy 2009 Science, originally
from Lean and Rind 2009
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How do we know about past conditions? Reconstructing past environments.
Reconstructing past climatic conditions and the 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
proxy typically come from organisms (e.g., trees, corals, plankton, diatoms, chironomids and other organisms)
respond to changes in climate through changes in growth rates or population numbers. 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 and tree-ring cores are two of the primary proxy used for reconstructing
past conditions for the western U.S.
Lake-sediment cores often provide some of the longest records of vegetation through analysis of pollen
and charcoal at regular intervals throughout the core. Most lakes in the northwestern U.S. were formed
following late-Pleistocene glaciation, 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 2003). 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.
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 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”.
Crossdating 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.
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Table 2. Example of Indirect proxy and scale of resolution-temporal range.
Proxy
Tree growth
Pollen
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Oxygen
Isotopes
Source
Tree-cores
Lake-sediment
cores
Corals, tree, lake,
ocean or ice cores
Temporal Resolution
High (annual)
Moderate (decadalcentennial)
High (annual) to
moderate (decadal)
Spatial Resolution
High
Moderate (several
km2)
Moderate to low
Temporal range
100-1000’s yrs
High (many
millennia)
High to very high
(100,000 +)
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In summary, modern, historical, and paleoecological data suggest that vegetation-climate linkages are
evident across many time scales, from individual records as well as regional and global compilations.
Shifts in vegetation and species distributions respond to climate variations occurring over multiple time
scales ranging from seasons to millennia. On longer time scales, anomalies in atmospheric circulation
patterns may influence the establishment and organization of dominance and expansion/contraction of
vegetation species and assemblages. On shorter time scales, anomalies in atmospheric circulation
patterns, storm events, volcanism and droughts may influence the relative abundance of individual
species and associated vegetative communities and the frequency and intensity of disturbances linked
to climatic conditions such as fire. On these shorter time scales, climate anomalies are often embedded
in large-scale teleconnection patterns manifest on inter-annual and inter-decadal time scales (i.e., slowly
varying atmosphere-ocean interactions and surface-energy feedbacks). It is on the longer time scales
that climate variations create, maintain, and change major distributions of vegetation and the species
they support. 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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What are the implications of climatic variation at different scales for management of ecosystems?
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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. For example, paleoenvironmental data suggest that past changes in fire regimes at large spatial
and temporal scales have been largely driven by climate conditions. Consequently, in ecosystems where
fire regimes are expected to change with future climate conditions, management efforts might focus on
adaptation to 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 conditions can have large
consequences and provide an important context for anticipating ecosystem response to future climate
change.
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Past 20,000 Years of Paleoenvironmental Change
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Biophysical conditions
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A number of drivers contribute to an understanding of millennial scale variation during the last 20,000
years including internal forcings: ocean-atmosphere-ice interactions, nonlinear feedbacks within the
climate system, stochastic processes, and external forcings such as variable solar output, orbital
harmonics and volcanism. Ocean-atmosphere-ice interactions explain some of the variation in
temperature changes and changes in ice-sheet dynamics and ocean circulation are strongly linked to
climate variation during the glacial period but became less influential as ice-sheets diminished (Fig. 4).
As ice-sheets receded, insolation variations, volcanic forcing and synoptic climate patterns gained
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influence over the western U.S. (Shuman et al. 2002) and weak periodic solar forcing may explain
millennial scale variation occurring on 1100-1200 yr. cycles as seen in Fig. 2b (Viau 2006).
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Figure 4. 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). Source: adapted
from Shuman et al. 2002
At the time of the Last Glacial Maximum, the large Laurentide and Cordilleran ice-sheets strongly
influenced climatic conditions in the western U.S. (Pederson et al. 2007, Whitlock et al. 2002).
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 (Fig. 5). Another element of the full-glacial
climate was the stronger-than-present surface easterlies related to a strong anticyclone that formed
over the ice sheets. This reduced the amount of precipitation and airflow from the west, effectively
drying out much of the western U.S. Unglaciated areas south of the ice sheets experienced climates that
were cold, dry, and windy (Bartlein et al. 1998).
Large-scale controls of climate changed during the glacial/Holocene transition causing the period from
16,000 to 11,000 yrs BP to be a period of great environmental and biotic adjustment (Pederson et al.
2007, Whitlock et al. 2002). Summer insolation in the Northern Hemisphere 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 in the northwestern US (45˚ latitude). The
regional climate influence imposed by the glacial maximum ice sheets also waned. One consequence
was a return of the position of winter storm tracks shifted northward, bringing increased winter
moisture, but warmer conditions than before from 16,000 to 11,000 cal yr BP (Bartlein et al. 1998).
During this transitional period, alpine glaciers and ice sheets retreated rapidly allowing for the
colonization of new vegetation. Gradually low- and middle-elevation areas were colonized by tundra,
then forest parkland, and finally forest and steppe. During this general ice recession, a brief period of
alpine glacier advance is registered between ca. 12,200 and 12,900 cal yr BP in the Northern Rockies.
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Figure 5. (left panel) Schematic of displacement and shift south of jet stream during LGM and (right panel) general
location of modern jet stream although path varies seasonally and is influenced by a number of factors such as
variation in ocean-atmospheric interactions (e.g., ENSO events). Ice-sheets also influenced atmospheric
conditions in the U.S. Rockies and northwestern U.S. creating generally dry and cool conditions (Whitlock et al.
2002). Right panel indicates primary air masses that influence study area today (Ahrens 2008).
Greater-than-present levels of summer insolation in the early Holocene (11,000 to 6,000 cal yr BP)
caused most areas to experience warmer- and drier-than-present conditions (Barnosky et al. 1987).
Modeling efforts show intensified summer insolation likely led to a strengthening of the eastern Pacific
subtropical high-pressure system and decreased precipitation in much of the northwestern US in the
summer (Thompson et al. 1993) while increasing inland flow of moisture from the Gulf of California to
the southwestern U.S. and parts of the southern U.S. Rockies (Whitlock and Bartlein 1993). During the
middle and late Holocene, summer insolation decreased and winter insolation increased gradually to
present levels. The cool wet conditions that characterized the late Holocene were established about
5,000 years ago, and coincident with this change in climate was the onset of modern plant communities
and renewed glacial advances (Pederson et al. 2007, Whitlock et al. 2002, Millspaugh et al. 2000).
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The Younger Dryas: a temporary return to glacial climate conditions ca. 12-13 ka yrs BP.
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An important event known as the Younger Dryas (YD) interrupted the transition from the last glacial
period to warmer and wetter climate and returned glacial conditions (Alley 2000, Alley et al. 1993). As
deglaciation commenced ca. 14 ka yrs BP, climatic conditions throughout much of the Rocky Mountains
and Upper Columbia Basin were becoming warmer and wetter, but an abrupt cooling event marked a
return to near glacial conditions for almost 1500 years ca. 12,900 – 11,500 yrs BP (Figs. 2, 4).
Temperatures decreased as much as 15˚ C in only a few decades and glacial advances are recorded
throughout higher latitudes in the western U.S. and Canada (Reasoner et al. 1994, Gosse et al. 1995,
Menounos and Reasoner 1997). The ecological response to the YD is indicated by a lowering in treeline
in Colorado (Reasoner and Jodry 2000) and is attributed to cooler temperatures rather than changes in
precipitation (Whitlock et al. 2002). The termination of this cooling period occurred rapidly, with
temperatures increasing 10˚C in only a few years (Patterson et al. 2009). The cause of the YD is still
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debated but evidence suggests that an influx of fresh water into the North Atlantic resulted in the
reduction or shut down of the North Atlantic thermohaline ocean circulation (Broecker 2006). An
alternative hypothesis suggests an abrupt cessation of ENSO events in response to orbital changes
caused this rapid reversal in climate amelioration (Broecker 2006). A similar event occurred ca. 8.2 ka
BP (Figs. 2, 4), where a large influx of freshwater disrupted circulation in the North Atlantic, causing
another period of cooling lasting several centuries (Le Grande et al. 2006, Alley and Agustsdotti 2005).
The Younger Dryas and the 8.2 ka event illustrate how rapidly climate changes due to oceanatmosphere-ice interactions can occur and how quickly these events can influence vegetation.
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A number of paleo records provide evidence for several periods of anomalously dry/wet conditions
during the Holocene (Shuman et al. 2009). Lake level data and other direct evidence for changes
millennial scale hydrologic change indicate much of the western U.S. experienced drought conditions
(drier than today) ca. 7000-4500 AD for (Shuman et al. 2009). Contrary to other periods in the Holocene
when changes in climate conditions were highly spatially variable, paleoenvironmental records
throughout the western U.S. suggest that conditions drier than today were widespread. Lake level
variation and pollen data from a large network of sites suggest that moisture balance was drier than
today during the mid-Holocene which was then followed by generally wetter than modern conditions
throughout much of the interior west. Anomalously dry conditions for much of the U.S. Rockies were
coupled with wetter 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).
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Mid-Holocene Drought in the U.S. Rocky Mountains and Upper Columbia Basin
While a number of records indicate widespread drought during the mid-Holocene, the mechanisms
explaining synoptic conditions that led to this drought are still not fully understood. Some propose that
a strengthening of the jet stream in the northwestern U.S. increased westerly flow, exacerbating
orographic loss of moisture in the west-slope of the Rockies, enhancing atmospheric aridity east of the
Rockies. This “zonal flow” hypothesis attributes much of the mid-continent aridity to enhanced jet
stream flow across the northwestern U.S. and mid-continent (Dean et al. 1996, Bartlein et al. 1984).
Alternatively, others argue that orbital changes and associated changes in seasonal insolation anomalies
best explain widespread mid-Holocene aridity (Shinker et al. 2006, Shuman et al. 2009, Shuman et al.
2010, Harrison et al. 2003). This second explanation suggests that variation in seasonal insolation
affects the dynamic between surface conditions and atmospheric processes in ways that enhanced
aridity during this period. Moisture availability (atmospheric flux and soil moisture recycling) and
suppressed precipitation via subsidence are thought to have played important roles in decreased aridity.
Specifically, paleo data from a network of sites throughout the Rockies suggest that orbital changes
influenced broad synoptic changes that: 1) enhanced atmospheric subsidence, 2) reduced winter
moisture advection linked to low winter insolation,, 3) increased evapotranspiration, and 4) altered
surface energy budgets and soil moisture which may have resulted in an important feedback driving
increased aridity, especially at higher elevations (Shuman et al. 2010, 2009).
Shuman et al. (2009) found that regional scale model simulations with fine-scale topography best
predict mid-Holocene aridity and, hence, conclude that insolation and associated surface feedbacks
(e.g., reduced snowpack and soil moisture) were likely more important drivers of millennial and
centennial scale changes in precipitation and moisture than changes in the strength of zonal flow. This
is supported by models showing strengthened summer monsoonal moisture existed in the American
Southwest which may have increased subsidence and aridity around the area of increased
convergence/convection and could have interacted with increased seasonality/direct influences of
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radiation to cause the widespread aridity evident in records from much of the U.S. Rockies during this
time period. The implications of this period of drier than normal conditions for the future are important
because they illustrate how increased temperatures will likely increase rates of evapotranspiration,
altering surface energy budgets related to snowpack and soil moisture, and increasing aridity at local
levels (Hoerling M.P. and Eischeid 2007). Hence, conditions that led to the mid-Holocene drought might
also influence much of the western U.S. with future warming.
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The last 20,000 years have seen climate changes that have driven extensive biological change and
reorganization beginning with the establishment of tundra parkland vegetation and followed by forest
cover. Climate amelioration following deglaciation led to increased moisture availability and effectively
warmer and wetter conditions that facilitated dramatic biotic change, allowing in-migration in many
areas once covered by snow and ice and the establishment of vegetation assemblages across biophysical
gradients. Warmer temperatures over the period from 11,000 to 6,000 cal yr BP favored drought
adapted species and resulted in higher-than-present treeline and frequent fire activity. Long-term solar
and orbital variations caused cooler and wetter conditions between 4,000-5,000 cal yrs BP resulting in
more extensive glacier advances (the Neoglacial period), a downslope shift of treelines, more closed
forests, increases in subalpine taxa, and less severe fire regimes. These cool and wet conditions during
the late Holocene (past 2-3 millennia) facilitated the development of modern vegetation assemblages
across much of the western U.S. (Whitlock et al. 2002, 1992).
Biotic change
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Southern Canadian and Northern U.S. Rockies
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During the full glacial period prior to 14,000 years BP, much of the Northern Canadian and U.S. Rockies
were glaciated. The postglacial colonizing vegetation prior to ca. 10,000 yr B.P. was comprised of cold
steppe vegetation of shrubs and herbs dominated by Artemisia, Gramineae and Alnus (Whitlock et al.
2002, Whitlock 1992, MacDonald 1989, Reasoner and Hickman 1989). By the early-Holocene, climate
amelioration but still cool and relatively dry conditions led to the establishment of shrub and herb
communities at higher elevations, and pioneering forests of Pinus albicaulis/flexilis and Alnus were
evident at mid elevations with lesser abundances of Picea and Pinus contorta. Lower elevations were
characterized by open parklands of pine and spruce and limber pine and grasslands at the lowest
elevations.
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Figure 6. Ecological response to changing climatic conditions following the retreat of glaciers in the Northern U.S.
and Southern Canadian Rockies (derived from MacDonald 1989 and Reasoner and Hickman 1989).
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Timberline remained at least 90 m above modern timberline elevation, during the period ca. 8500 yr
B.P. to ca. 3000 yr B.P. in response to warmer climatic conditions and limited glacial activity. Low/midelevations experienced the greatest fire activity during the Holocene ca. 9-8 ka BP (MacDonald 1989)
and forest compositions resembling the modern subalpine Picea-Abies forest had developed by the end
of this period. The late-Holocene, ca. 3000 yr B.P. to present, was marked by deteriorating climatic
conditions associated, increased glacial activity and declining timberlines (Reasoner and Hickman 1989).
Cooler and moister conditions that prevail today led to modern vegetation of tundra at high-elevations,
spruce and fir forests at mid-elevations and spruce, lodgepole pine and aspen at lower elevations.
Paleorecords show dynamic vegetation response to deglaciation, warmer-wetter conditions during the
mid-Holocene and the establishment of the cool/moist conditions that characterized much of the
southern Canadian and northern U.S. Rockies today.
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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 at different elevations throughout the Parks (Fig. 7, Whitlock 1993). Deglaciation (ca. 17-14 ka
BP) was first followed by colonization of ice-free areas by tundra vegetation (Betula, salix, Juniperus,
then sagebrush and grasses (Artemisia and Poaceae) at mid elevations and spruce (Picea) at lower
elevations (ca. 14-10 ka BP; Fig. 7). With continued climate amelioration, and effectively warmer and
wet conditions, pine, juniper and birch trees were present at low elevations, lodgepole pine (Pinus
contorta), Douglas fir (Pseudotsuga menziesii) and aspen (Populus) trees became established at midelevations and subalpine forests (Picea and Abies) occupied higher elevations. A number of
paleorecords indicate widespread drought during intervals of the mid-Holocene but warmer and drier
conditions resulted in only subtle shifts in elevational ranges of dominant vegetation in the central U.S.
Rockies and GYA. Decreased summer insolation in the late-Holocene (ca. 3-4 ka BP) led to cooler and
wetter conditions similar to those experienced today in much of the CR-GYA and vegetation
assemblages further differentiated across elevational gradients with a sagebrush-steppe present at low
elevations, and limber pine (Pinus flexilus), Douglas fir and lodgepole pine dominating mid-elevations
(Whitlock et al. 1993). Spruce, fir and undifferentiated pine forests were present at higher elevations
and cool, dry conditions at the highest elevations limited vegetation to tundra grasses and herbs.
Figure 7. Ecological response to changing climatic conditions following the retreat of glaciers in the Central U.S.
Rockies and Greater Yellowstone Area (derived from Whitlock 1993).
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Following deglaciation, the CR-GYA experienced rapid and unprecedented biotic reassembly as new
areas were available for vegetation to colonize. The late glacial and early-Holocene saw the invasion of
tundra vegetation and eventual establishment of subalpine forests at higher elevations, lodgepole pine
forest dominated much of the mid-elevations and pine, juniper and birch forests were present at lower
elevations. With the onset of modern climatic conditions (i.e., cooler and wetter) lodgepole pine
persisted throughout much of its range but subalpine forests were confined to a lower and narrower
range of elevations and forests and steppe communities further differentiated at low elevations. By the
onset of the Holocene, much of the vegetation that exists today was present in the CR-GYA, and
following a period of warm and dry conditions, the late-Holocene was marked by continued re-assembly
and differentiation of vegetation communities.
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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, Whitlock et al.
2009). Pollen and charcoal diagram from
Cygnet Lake, central Yellowstone Source:
Millspaugh et al. 2000
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Cygnet Lake
An examination of charcoal, pollen and
climate conditions from central
Yellowstone shows a strong link
between changes in climate and
patterns of fire at millennial time scales
over the past 17,000 yrs. Pollen from
Cygnet Lake in the Central Plateau of
Yellowstone indicates the area was
dominated by tundra community of
sagebrush (Artemisia) and grass (Poaceae) between ca. 17,000 and 12800 cal yr B.P. when cool and wet
conditions prevailed (Millspaugh et al. 2000). Increased summer insolation (8.5% greater than present)
during the late-glacial/early Holocene (ca. 12,000-11,300 yrs. B.P.) led to warmer and drier conditions in
parts of central and northwestern Yellowstone facilitating increases in Pinus (diploxylon-type) pollen and
establishment of lodgepole pine forests (Millspaugh et al. 2000). A decrease in insolation ca. 7,500 led
to cooler and wetter conditions throughout Yellowstone yet vegetation changed little throughout the
rest of the Holocene. The transition from tundra vegetation to pine forest at the onset of the Holocene
marked a fundamental shift in vegetation that persisted for thousands of years despite changes in
climatic conditions including cooler temperatures and increased effective moisture (Millspaugh et al.
2000).
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This study provides important lessons for understanding climate-vegetation linkages. First, unlike
vegetation in other parts of the CR-GYA that were more sensitive to changes in Holocene climate
conditions, lodgepole pine forests on rhyolitic soils in the Central Plateau responded little to changes in
climate and suggest that susbtrate can act as an important long-term control on vegetation in some
settings (Whitlock et al. 1993, Millspaugh et al. 2000). Second, the interval of increased summer
insolation that led to warm and dry conditions in the Central Plateau of Yellowstone resulted in
effectively wetter conditions in other parts of southern and eastern Yellowstone and the southwestern
U.S. where increased seasonality was followed by intensified monsoonal moisture (Whitlock and
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Bartlein 1993). Hence, the response of climate conditions to changes in seasonal radiation were highly
spatially variable and while the contrast between these different regions weakened in the late-Holocene
(Whitlock et al. 1995), evidence suggests that this transition zone between Pacific Northwest and
Southwest U.S. atmospheric circulation patterns persists today and continues to influence climate and
vegetation in the western U.S. Increased seasonality resulting from ocean-atmosphere and land-surface
feedbacks may intensify the influence of these circulation types in the future (Shuman et al. 2010, 2009,
Gray et al. 2003, 2007).
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A network of vegetation reconstructions from pollen and macrofossil data provide a history of climate
and vegetation change in western Colorado during the late glacial and Holocene (Fall 1997). Prior to 14
ka BP, cooler (2-5˚ C cooler than present) and wetter (7-16 cm wetter than present) conditions
supported tundra vegetation (e.g., Artemisia, Poaceae, Acomastylis, Dryas) at high-elevations (~300-700
m below present treeline) and spruce parklands at lower elevations (Fig. 9, Fall 1997). During the lateGlacial period, (ca. 11 ka BP), Abies increased in abundance in both the subalpine forest which became
established at mid-elevations and in lower elevation mixed spruce, pine, fire parkland as climatic
conditions became cooler winter precipitation increased (Fall 1997). Summer insolation increased
during the early Holocene, and warmer temperatures allowed subalpine forests to expand upslope.
Whereas 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 conditions.
Paleoclimatological data suggest that increased summer insolation resulted in an enhanced summer
monsoon and consequent greater summer precipitation during this period (Thompson et al. 1993).
Pollen data from the region support greater summer than winter precipitation by the dominance of
Picea engelmannii throughout the mid and early Holocene. From 9-4 ka BP, warm summer
temperatures (1.9˚ C mean summer temperature above present) facilitated the upslope expansion of
subalpine forests of Picea engelmannii and Abies lasiocarpa, at least 300 m greater than today to almost
4,000 meters (Fall 1997). Subalpine conifers, Picea engelmannii and Abies lasiocarpa, also expanded to
lower elevations below 3000 m, along with montane conifers (Pinus sp.) and these conditions lasted
until the mid-Holocene ca. 6 ka BP. Fall (1997) estimates summer precipitation increased by 8 -11 cm as
indicated by the lowering of these forests.
Southern U.S. Rockies
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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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By the end of the mid-Holocene, lower montane forests mixed with steppe vegetation, and lower
elevation subalpine forests became more open even though summer monsoons continued to provide
precipitation that was still several centimeters greater than today. By 4 ka BP summer temperatures
had declined with the decreased summer radiation, spruce forests consequently moved lower and fir
(Abies lasiocarpa) dominated treeline forests and krummholz vegetation for the first time during the
Holocene, likely in response to a reduction in the intensity of summer monsoonal moisture and a return
to winter dominated precipitation regime. During the late-Holocene, montane taxa retreated upslope,
sagebrush steppe expanded at lower elevations and alpine tundra dominated larger range of highelevations suggesting drier conditions increased for parts of the SR. Conditions similar to the present
were established two millennia ago although treeline and krummholz levels have fluctuated during the
Medieval Warm Period. As in the northern and central U.S. Rockies, millennial scale variation in climatic
conditions led to significant redistribution of vegetation at across elevation gradients in the Southern
U.S. Rocky Mountains. Unlike much of the Rocky Mountains at higher latitudes, increased summer
insolation and warmer conditions were accompanied by greater precipitation associated with an
intensified summer monsoons. This combination of warmer and wetter conditions allowed for the
expansion of forests at both their upper and lower elevational limits. As heightened seasonal insolation
extremes abated, modern climatic conditions became established along with slightly cooler and drier
conditions that prevail today.
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Upper Columbia Basin
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Paleoenvironmental data from the Upper Columbia Basin (UCB) (Barnosky 1985), the Snake River Plain
(Davis 1986) and the mountains of southeastern Idaho (Beiswenger 1991, Bright 1966) provide a history
of climatic conditions during the past 20 ka and the ecological response to millennial scale climatic
variation. Prior to the Holocene, cool and dry conditions across much of the UCB and Snake River Plain
led to the establishment of tundra-steppe communities dominated by Artemisia and Poaceae. At
higher elevations, greater available precipitation allowed for the establishment of Pinus, Picea and Abies
forests which expanded with climate warming during the late glacial and early Holocene. For much of
the UCB, cool dry conditions that persisted until the Holocene were followed by warmer and drier
conditions. Increased seasonal radiation led to dry conditions and the expansion of Juniperus, Artemisia,
Chenopodiaceae, Amaranthaceae and other temperate taxa where precipitation limited the
establishment of closed forests. At higher elevations increased abundance of Pinus suggest that
summer droughts were likely more common (Whitlock 1992, Beiswenger 1991). As summer insolation
decreased ca. 8.5 ka BP, temperatures cooled and more available moisture allowed for the
establishment and expansion of Pinus woodlands at mid and higher elevations. Further cooling and the
onset of modern climatic conditions during the mid-Holocene led to the expansion of Pinus parklands ca.
4 ka BP and the establishment of mixed forests of Pinus, Pseudotsuga and Abies at wetter sites
(Whitlock 1992, Whitlock 1985). 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 Picea and Abies pollen increase (Whitlock
1992).
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Figure 10. Ecological response to changing climatic conditions following the retreat of glaciers in the Upper
Columbia Basin (derived from Whitlock 1992, 1985, Beiswenger 1991).
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For much of the Upper Columbia Basin, cool and dry conditions that persisted through the late-glacial
gave way to warm and dry conditions associated with the amplification of seasonal radiation during the
early and mid Holocene. The late Holocene gave way to a return to cool and dry conditions that
characterize modern day climate for much of the Basin. Low levels of precipitation and available
moisture throughout most of the last 20 ka BP have limited the establishment of closed forests across
low and mid-elevations of the Upper Columbia Basin.
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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 1850 until human activities began to significantly impact levels of greenhouse
gases and aerosol concentrations (Barnett et al. 2008, Bonfils et al. 2008, Keeling 1976). Volcanic
eruptions and solar fluctuations were likely the most strongly varying external forcings during this
period, but it is currently estimated that temperature variations caused by these forcings were much
less pronounced than the warming due to greenhouse gas forcing since the mid-19th century. Climate
model simulations indicate that solar and volcanic forcings together likely produced periods of relative
warmth and cold during the preindustrial portion of the last 1,000 years. However, anthropogenic
greenhouse gas increases are needed to simulate late 20th century increases in temperature.
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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
variation is evident during this time period and can be linked to fluctuating radiation resulting from
changes in solar output, volcanic forcing, and ocean atmosphere interactions (Jackson et al. 2009).
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 dominant
Biophysical Conditions
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vegetation. Since the late-Holocene, vegetation types throughout this period are considered similar to
today and, overall, the magnitude and duration of the changes are not comparable to those of the
Pleistocene/Holocene transition reflecting shorter (centennial, decadal, inter-annual) scale climatic
controls on ecosystems. At these shorter time scales ocean-atmosphere interactions such as the Pacific
Decadal Oscillation (PDO), the North Atlantic Oscillation (NAO), and the El Nino Southern Oscillation
(ENSO) interact to influence temperature, precipitation, and atmospheric circulation and help explain
droughts, wet periods and decadal and annual climatic events.
Tree-ring records from throughout the western U.S. show natural variation in wet and dry cycles 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. 11-12, 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 (Pederson et al. 2006, Cook et al. 2004). 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).
Figure 11. Location of tree-ring-based precipitation and drought reconstructions used in comparison of moisture
conditions along a north–south Rocky Mountain transect. (right) Tree-ring-based reconstructions of moisture
anomalies. Each series has been normalized and smoothed using a 25-yr cubic spline to highlight the prominent
20–30-yr frequencies identified by MTM spectral analysis (Mann and Lees 1996). Source: Pederson et al. 2006
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Figure 12. (A) Smoothed DAI reconstruction (solid black curve) for the West, showing two-tailed 95% bootstrap
confidence intervals (dashed black curves) and the long-term mean (thin horizontal black line). Sixty-year
smoothing was applied to highlight the multidecadal to centennial changes in aridity. The four driest epochs (P <
0.05, those with confidence limits above the long-term mean in Fig. 11A) are before 1300 AD, whereas the four
wettest (P < 0.05) epochs occur after that date. The difference between the means of the 900 to 1300 AD period
(redline, 42.4%) and 1900 to 2003 AD period (blue line, 30%) are also apparent. The 12.4% difference between the
two periods translates into an average drought area (PDSI < –1) increase of 41.3% in the West during the earlier
period. This difference is statistically significant (P < 0.001) given an equality-of-means t test with degrees of
freedom corrected for first-order autocorrelation. Even so, some of the 900 to 1300 AD period PDSI estimates are
extrapolations, because they fall outside the range of the instrumental PDSI data in the 1928 to 1978 AD
calibration period. As regression-based estimates, these extrapolations have greater uncertainty compared to
those that fall within the range of the calibration period. However, they are still based on the actual growth
histories of highly drought-sensitive trees. Therefore, we argue that our DAI reconstruction is indicative of what
really happened in the West, even during the AD 900 to 1300 period of elevated aridity. (B) The annually resolved
AD 1900 to 2003 portion, which more clearly reveals the severity of the current drought relative to others in the
20th century and an irregular trend (red smoothed curve) toward increasing aridity since 1900. Source: Cook et al.
2007
Two major, well documented examples of centennial scale climate change within the two millennia are
the Medieval Warm Period (MWP; 950-1250 AD) and the Little Ice Age (LIA; 1400-1700 AD). As
indicated by Figures 12-13 (Cook et al. 2007, Mann et al. 2009) a number of warm (dry) and cool (wet)
periods during this time frame exist (see Biondi et al. 1999 and Cook et al. 2007) yet these two periods
of anomalous climate conditions were unusual in duration, extent of influence and ecological impacts.
During the MWP, positive temperature anomalies persisted across western North America (Fig. 13).
While the general climatic trend during this time was towards warmer, drier conditions, the local effects
were highly spatially variable. Elevated aridity and mega-droughts were commonplace during this
period, with synchronous paleoenvironmental records of aridity occurring throughout the study area
(Fig. 12, Cook et al. 2007).
Data from a number of sites suggest regionally synchronous drought events occurring regularly during
the Medieval Warm Period, with durations and extents unmatched in the late Holocene (Cook et al.
2007). Glacial retreat in mountainous locations occurred in Colorado, Wyoming, Montana, and the
Cascades and lake levels generally dropped throughout the study area (Millspaugh et al. 2000, Brunelle
and Whitlock 2003). However, spatial and temporal variations in this general trend were widespread, as
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the CR-GYA and parts of the SR experienced increased summer monsoonal activity (Davis 1994, Whitlock
and Bartlein 1993, Thompson et al. 1993, Petersen 1988) and generally wetter conditions. Treelines in
some areas advanced upwards in elevation and areas now covered by krummholz were occupied by
arborescent trees (Whitlock et al. 2002). Additionally, alpine larch expanded 90 kilometers north of its
current range (Whitlock et al. 2002) ca. 950 1100 BP. The mechanisms and drivers leading to the MWP
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. 13, Mann et. al 2009).
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Throughout the study area, modern species assemblages were generally present throughout the past
two millennia and show that many of the ecosystems we see today in each of the four climate regions
are relatively young (Jackson et al. 2009). Accordingly, no massive climate-driven ecologic
Fig. 13. 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
In contrast to the MWP, 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 (Whitlock et. al 2002, Pedersen et al. 2007). While temperatures across the study
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 (Whitlock et al. 2002). Here paleo data document major glacier advances and increased
winter precipitation (Pederson et al. 2007). The Southern and Central U.S. Rockies did not see advances
of this magnitude and regional signals of centennial scale cooling are less apparent, demonstrating the
continental variability in LIA climatic impacts. As with the MWP, 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).
Biotic change
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reorganizations analogous to those associated with continental deglaciation or the
Pleistocene/Holocene transition occurred during this time period. Instead, existing species changed in
abundance and extent in response to changing climatic conditions and associated disturbance regimes.
For example, during the last two millennia, tree lines expanded to higher elevations during warm and
moist periods and decreased in elevation during severe droughts and cooler conditions (Graumlich et al.
2005, Rochefort et al. 1994, Winter 1984). Additionally, increased temperatures and droughts
associated with MWP conditions likely resulted in increased fire activity throughout the study area (Fig.
14, Gray et al. 2007, Cook et al. 2004, Whitlock et al. 2002). With the onset of the LIA, decreased
temperatures and glacial advanced resulted in alpine tree mortality, downslope movement of treeline,
and species distributions shifted again, and fire activity decreased with respect to the MWP. In
comparing natural variations in precipitation during the past millennium Gray et al. (2007) assert that
major disturbances (e.g, fire) and fundamental transitions in vegetation coincide with multi-decadal and
decadal scale variations in precipitation and that this variation is a primary driver of ecosystem dynamics
for the region (Fig. 14).
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Figure 14. Major disturbance
events and landscape transitions
in the greater Yellowstone area
(after Romme et al., 1995; Meyer
and Pierce, 2003; Larsen and
Ripple, 2003) related to
precipitation regimes captured in
this reconstruction. Precipitation
data are the 60-yr smoothed
series. Source: Gray et al. 2007
Quat. Res.
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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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A 3,800 year history of climatic, vegetation, and ecosystem change is preserved in the 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). Lake sediment 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 record (Jackson et al. 2009). Following the establishment of mixed conifer forests,
lake levels decreased from 2200 to 1200 years BP and increases in grass, pine and sagebrush and
Northern U.S. Rockies
Climate, vegetation and ecosystem change in northwestern Montana during the late Holocene
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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).
These shifts in vegetation were accompanied by pronounced changes in patterns of fire as evidenced by
charcoal deposition in the lake sediment cores. Intervals dominated by forests coincide with highmagnitude 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. The Foy Lake
lake-sediment 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) centennial and
decadal shifts in climatic conditions are evident in changes in the charcoal record and associated
patterns of fire for the past several millennia. The Foy Lake record illustrates how climate change can
influence vegetation via direct climate constraints or by influencing key ecosystem processes such as fire
and how feedbacks related to changes in vegetation can also influence ecosystem processes such as fire
by increasing fuel availability. The Foy Lake record also allows for a consideration of recent ecosystem
dynamics in the larger historical context. 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. A few important lessons emerge from the Foy Lake
record: 1) 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, 2) many
terrestrial ecosystems like those in the vicinity of Foy Lake are relatively young, dynamic and have only
recently established in response to the interaction of a number of climate drivers operating at different
scales and, 3) as evidenced by changes in fire regimes of the later 19th and eary 20th centuries, human
activities have left a strong imprint on the landscape and can rapidly influence ecosystems and key
ecosystem processes such as fire.
Could also summarize key points from Pederson et al. 2006 Earth Interactions here.
Upper Columbia Basin
Natural variations in temperature in east-central Idaho during the last millennium
A study of tree-rings from the Sawtooth – Salmon River region in east-central Idaho provides a highresolution regional temperature proxy for the past 858 years (Biondi et al. 1999). July temperatures for
the 858 year period were reconstructed from whitebark pine (Pinus albicaulis) and Douglas fir
(Pseudotsuga menziesii) tree-cores and show periodic multi-decadal temperature fluctuations within
centennial temperature trends. Cold and warm periods lasting decades to centuries alternate
throughout the period, illustrating regional temperature fluctuations during the past millennium (Fig.
15). The tree-ring records indicate a prolonged cold period from ca. 1200 – 1350 AD was followed by
alternating multi-decadal cold and warm periods up to the present (Biondi et al. 1999). Over the entire
858 year period, the coldest modeled July temperatures occurred just after 1600 AD (Biondi et al. 1999).
These anomalously cold temperatures correlate well with other northern hemisphere temperature
reconstructions and coincide with the eruption of Huaynaputina in Peru (de Silva and Zielinksi 1998).
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Temperatures increased by mid-17th century, dropped during the LIA (ca. 1850) and recover again in the
20th century. The Biondi et al. (1999) July temperature reconstruction demonstrates multi-decadal and
annual temperature fluctuations and patterns that characterize climatic variations over the past
millennia.
Figure 15. Inter-annual and inter-decadal (thick line) variability of July temperature during the second millennium,
A.D. 1135-1992. Warm and cold intervals of decadal to centennial duration alternate in the temperature
chronology. Source: Biondi et al. 1999
Centennial scale cooling is notably absent during the LIA, illustrating the how an event that is evident at
sites throughout the western U.S. is not uniformly manifest across the landscape. Multi-decadal
variation in temperatures throughout the last two millennia are also strongly linked to volcanic
eruptions and consequent cooling, highlighting the significance of volcanism as an important driver of
climatic conditions on annual to decadal time scales. Records from this area of the Upper Columbia
Basin show a strong pattern of multi-decadal variation in wet and dry periods and highlight the fact that
recent warm and dry intervals are unremarkable when considered in comparison to past droughts and
pluvials (ca. AD 1300 and 1500). Data from the Sawtooth – Salmon River reconstruction suggests that
climate events that are synchronous across regions are not always uniformly manifest at small scales
and also that climatic forcings are represented climatically in different ways in different places as a
result of interactions among controls (McCabe et al. 2008, Jackson et al. 2009).
Central U.S. Rockies and GYA
Changing distributions and invasion of Utah Juniper in Wyoming, Montana and Utah during the late
Holocene
Changes in the distribution of Utah juniper (Juniperus osteosperma) in the mountains of Wyoming,
Montana, and Utah during the late Holocene demonstrate how ecosystems may respond to future
climates (Lyford et al. 2003). Analyzing radiocarbon-dating fossilized woodrat middens in the Bighorn
Basin of north-central Wyoming and adjacent Montana, Lyford et al. (2003) reconstructed
spatiotemporal patterns of Utah juniper migration and invasion during the Holocene. Utah juniper
thrives in warmer, arid conditions in the Great Basin and Colorado Plateau. During a dry period in the
mid-Holocene (ca. 7500 – 5400 BP), Utah juniper migrated into the Central Rockies of Wyoming,
Montana, and eastern Utah 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 where it had pre-established and within the Bighorn Basin,
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Utah juniper establishment was particularly high from 2800 to 1000 years BP. The notable absence of
significant Utah juniper establishment and expansion during the MWP suggests that longer scale climate
variations may be of increased importance in determining 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. 16). 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.
Figure 16. 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
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 millennium 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 long-distance 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
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species dispersal will interact with the spatial distribution of suitable habitat and climate variability to
govern future invasions.
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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. 17).
Natural variability in precipitation in Wyoming’s Green River Basin during the last 1100 years
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Figure 17. 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
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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. 17). 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.
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Spatio-temporal patterns of pinyon pine distributions and multi-decadal precipitation variability during
the late Holocene
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In the Southern Rockies, the spatiotemporal patterns of pinyon pine (Pinus edulis Engelm.) distribution
in the Dutch John Mountains (DJM) of northeastern Utah are strongly controlled by variations in multidecadal precipitation patterns (Gray et al. 2006). Gray et al. (2006) used woodrat midden and tree-ring
data to track the spatiotemporal patterns of pinyon pine distribution over the ensuing millennium. The
DJM population of pinyon pine is an isolated northern outpost of the pinyon pine, and was initially
colonized by AD 1246. Similar to Utah juniper in the Central Rockies, the pinyon pine reached the DJM
via long-distance dispersal from the Colorado Plateau to the south (Jackson et al. 2005). The period of
DJM pinyon pine colonization coincides with the transition from the warmer, drier MWP to the cooler,
Southern U.S. Rockies
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wetter LIA (Fig. 18). Other studies of paleo vegetation patterns in the region suggest that the end of the
MWP was a time of significant vegetation adjustment throughout the Southern Rockies (Gray et al.
2006). However, changes in the distribution of the DJM pinyon pine population was episodic and did
not respond smoothly to centennial scale climate amelioration (Gray et al. 2006). Although evidence of
recruitment is present in the 13th century, DJM pinyon pine establishment was stalled during a severe
multi-decadal drought late in the 1200s and significant recruitment did not resume until the 14th century
pluvial when regionally mesic conditions promoted establishment within the DJM (Fig. 18).
Figure 18. (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 Ecology
The case of the pinyon pine in the DJM demonstrates the importance of episodic, multi-decadal 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
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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
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 reconstructions of past climates at different temporal and
spatial scales and the ecological response to natural variability in climate at different scales provides
important context for anticipating future change. Historical records from the past 20,000 years show
vegetation has responded continuously to climatic variation operating on millennial to inter-annual
scales. At large scales, climate ultimately governs the distribution of vegetation across landscapes of the
northwestern U.S. and this is evident from paleorecords. As evidenced by research from different
regions and time periods, climate variability regulates the distribution of vegetation across the
landscape, although it is not the sole factor determining species distributions and movements. 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 magnitude and temporal scale at which climatic
variation occurs is important. Species can tolerate variation and respond by dispersing and colonizing to
newly suitable sites unless variation in climate is too extreme for an extended length of time or dispersal
to newly suitable habitat is not possible. Gray et al. (2006) and Lyford et al. (2003) use high-resolution
data to demonstrate that plant invasions and migrations are better characterized as episodic
movements rather than continuous steady expansions. 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).
In summary, the long-term paleoenvironmental history of the climate regions included in this study
indicates:
What does the long-term paleoenvironmental record tell us about past climate change?
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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.
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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 Warm
Period and the Little Ice Age) but were not uniformly manifested across the study area.
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.
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Anticipating future climatic changes and the ecological response to these changes will require
consideration of the interaction between natural climate variability and anthropogenic drivers of
change. It is already evident that the increased rate of change related to anthropogenic influences will
add novel stresses to many ecosystems throughout the western U.S., making the ecological response to
these new conditions difficult to predict. Consequently, past change suggests that the combination of
natural variability and relatively extreme and rapid anthropogenic contributions to high levels of
greenhouse gasses and land-use change will likely lead to novel climatic conditions and ecosystems and
ecological responses that are surprising. A number of important lessons emerge from longer-term
records of past climate and the ecological response.
1012
Ideas for summarizing this information better?
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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.
Rapid climate transitions and associated ecosystem transitions have occurred in the past and
will likely occur in the future.
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