National Park Service

U.S. Department of the Interior

Natural Resource Program Center

Climate and Terrestrial Ecosystem Change in the U.S. 

Rocky Mountains and Upper Columbia Basin

Historical and Future Perspectives for Natural Resource 

Management

Natural Resource Report NPS/GRYN/NRR—2010/260
ON THE COVER

Grinnell Glacier, Glacier National Park 

NPS photo by Doug McMains
Climate and Terrestrial Ecosystem Change in the U.S. 

Rocky Mountains and Upper Columbia Basin

Historical and Future Perspectives for Natural Resource 

Management

Natural Resource Report NPS/GRYN/NRR—2010/260

David B. McWethy

Department of Earth Sciences

Montana State University

Bozeman, MT 59717

Stephen T. Gray

Water Resource Data System

University of Wyoming

Laramie, WY 82071

Philip E. Higuera

Department of Forest Resources

University of Idaho

Moscow, ID 83844

Jeremy S. Littell

Climate Impacts Group

University of Washington

Seattle, WA 98195

Gregory T. Pederson

U.S. Geological Survey

Northern Rocky Mountain Science Center

Bozeman, MT 59715

Andrea J. Ray

NOAA Earth System Research Lab

Boulder, CO 80305

Cathy Whitlock

Department of Earth Sciences

Montana State University

Bozeman, MT 59717

October 2010

U.S. Department of the Interior

National Park Service

Natural Resource Program Center

Fort Collins, Colorado
ii

Title of Report 

The National Park Service, Natural Resource Program Center publishes a range of reports that address natural 

resource topics of interest and applicability to a broad audience in the National Park Service and others in natural resource management, including scientists, conservation and environmental constituencies, and the public. 

The Natural Resource Report Series is used to disseminate high-priority, current natural resource management 

information with managerial application. The series targets a general, diverse audience, and may contain NPS 

policy considerations or address sensitive issues of management applicability.

All manuscripts in the series receive the appropriate level of peer review to ensure that the information is scientifically credible, technically accurate, appropriately written for the intended audience, and designed and published 

in a professional manner. 

This report received formal peer review by subject-matter experts who were not directly involved in the collection, analysis, or reporting of the data, and whose background and expertise put them on par technically and 

scientifically with the authors of the information.

Views, statements, findings, conclusions, recommendations, and data in this report do not necessarily reflect 

views and policies of the National Park Service, U.S. Department of the Interior. Mention of trade names or 

commercial products does not constitute endorsement or recommendation for use by the U.S. Government.

This report is available from Greater Yellowstone Inventory and Monitoring Network (http://science.nature.nps.

gov/im/units/gryn/index.cfm) and the Natural Resource Publications Management Web site (http://www.nature.

nps.gov/publications/NRPM) on the Internet. 

Please cite this publication as:

McWethy D. B., S. T. Gray, P. E. Higuera, J. S. Littell, G. T. Pederson, A. J. Ray, and C. Whitlock. 2010. Climate 

and terrestrial ecosystem change in the U.S. Rocky Mountains and Upper Columbia Basin: Historical and 

future perspectives for natural resource management. Natural Resource Report NPS/GRYN/NRR—2010/260. 

National Park Service, Fort Collins, Colorado.

Climate and Ecosystem Change in the U.S. Rocky Mountains and Upper Columbia Basin

NPS 963/105898, October 2010
iii

Contents Contents

Figures .....................................................................................................................................vi

Tables .....................................................................................................................................viii

Executive summary.....................................................................................................................ix

Acknowledgements....................................................................................................................xi

Chapter 1: Introduction..............................................................................................................1

1.1 Climate controls and variability at different spatio-temporal scales......................................2

1.2 How can understanding climatic variability inform management?.......................................5

Chapter 2: The last 20,000 years of environmental change in the western United 

States.............................................................................................................................................9

2.1 Drivers of millennial-scale climate variation ..........................................................................9

2.2 Glacial–Holocene transition..................................................................................................9

2.3 Mid-Holocene transition.....................................................................................................11

2.3.1 Southern Canadian and Northern U.S. Rocky Mountains............................................11

2.3.2 Central U.S. Rocky Mountains and the Greater Yellowstone Area ..............................12

2.3.3 Cygnet Lake.................................................................................................................14

2.3.4 Southern U.S. Rocky Mountains..................................................................................14

2.3.5 Upper Columbia Basin.................................................................................................15

Chapter 3: What do paleoenvironmental records tell us about millennial scale climate 

variations? ......................................................................................................................................17

Chapter 4: The last 2000 years of environmental change ....................................................19

4.1 Primary drivers of change ...................................................................................................19

4.2 Biophysical conditions.........................................................................................................19

Chapter 5: Case studies .............................................................................................................23

5.1 Northern U.S. Rocky Mountains.........................................................................................23

5.1.1 Drought variability and ecosystem dynamics in Glacier National Park.........................23

5.2 Central U.S. Rocky Mountains and the Greater Yellowstone Area .....................................24

5.2.1 Changing distributions of Utah juniper .......................................................................24

5.2.2 Precipitation variability in Wyoming’s Green River Basin .............................................25

5.3 Southern U.S. Rocky Mountains.........................................................................................26

5.3.1 Changing pinyon pine distribution ..............................................................................26

5.4 Upper Columbia Basin........................................................................................................27

5.4.1 Climate variation and fire-related sedimentation ........................................................27

Chapter 6: What can we learn from the last 2000 years about decadal and centennial 

scale climate change?................................................................................................................29

Chapter 7: 20th century climate change and the instrumental record.................................31
iv

A Framework for Climate Analysis and Reporting for Greater Yellowstone and Rocky Mountain Networks Climate and Ecosystem Change in the U.S. Rocky Mountains and Upper Columbia Basin

7.1.1 Temperature ................................................................................................................31

7.1.2 Precipitation ................................................................................................................32

7.1.3 Surface hydrology........................................................................................................32

7.1.4 Ocean-atmosphere interactions ..................................................................................33

7.1.5 Changes in storm track and circulation patterns.........................................................34

7.1.6 Ecological impacts.......................................................................................................34

7.2 Northern U.S. Rocky Mountains.........................................................................................36

7.2.1 Temperature ................................................................................................................36

7.2.2 Precipitation ................................................................................................................37

7.2.3 Surface hydrology........................................................................................................37

7.2.4 Ocean-atmosphere interactions ..................................................................................37

7.3 Central U.S. Rocky Mountains and the Greater Yellowstone Area .....................................38

7.3.1 Temperature ................................................................................................................38

7.3.2 Precipitation ................................................................................................................38

7.3.3 Ocean-atmosphere interactions ..................................................................................39

7.4 Southern U.S. Rocky Mountains.........................................................................................39

7.4.1 Temperature ................................................................................................................39

7.4.2 Precipitation ................................................................................................................39

7.4.3 Surface hydrology........................................................................................................40

7.4.4 Ocean-atmosphere interactions ..................................................................................40

7.5 Upper Columbia Basin ........................................................................................................40

7.5.1 Temperature ................................................................................................................40

7.5.2 Precipitation ................................................................................................................40

7.5.3 Surface hydrology........................................................................................................41

7.5.4 Ocean-atmosphere interactions ..................................................................................41

Chapter 8: What can we learn from 20th century observations?.........................................43

8.1.1 Small changes can have large impacts ........................................................................43

8.1.2 Shifting distributions and new norms..........................................................................43

Chapter 9: What can we expect in the future?......................................................................47

9.1.1 GCM projections for North America............................................................................47

9.1.2 Westwide climate: Statistically downscaled projections ..............................................49

9.1.3 Climate projections downscaled to specific alpine sites...............................................50

9.1.4 Model projections of future climatic and hydrologic conditions .................................52

9.1.5 Downscaled model methodology................................................................................52

Chapter 10: Climate projections for the western United States..........................................55

10.1.1 Climate conditions.....................................................................................................55
v

Contents

10.1.2 Surface hydrology......................................................................................................55

10.1.3 Extreme conditions: droughts, floods, heat waves....................................................55

10.1.4 Productivity and phenology.......................................................................................55

10.1.5 Disturbance ...............................................................................................................55

Chapter 11: Planning for the future ........................................................................................57

Chapter 12: Summary conclusions...........................................................................................59

Literature cited...........................................................................................................................61
vi

A Framework for Climate Analysis and Reporting for Greater Yellowstone and Rocky Mountain Networks Climate and Ecosystem Change in the U.S. Rocky Mountains and Upper Columbia Basin

Figures

Figure 1. Location of study area. ....................................................................................................................1

Figure 2. Primary drivers of climate and resulting climate variations at millennial, centennial, and 

interannual scales. ......................................................................................................................................3

Figure 3. Area of the Laurentide ice sheet (top panel) and central Greenland temperature reconstruction 

(bottom panel).................................................................................................................................................9

Figure 4. Primary air masses that influence the study area. ............................................................................10

Figure 5. Ecological response to changing climatic conditions following glacial retreat in the southern 

Canadian and northern U.S. Rockies. .............................................................................................................11

Figure 6. Changes in pollen percentages for different plant taxa during the last 3700 years. .........................13

Figure 7. Ecological response to changing climatic conditions following glacial retreat in the Central Rockies 

and the summer-dry region of the Greater Yellowstone Area..........................................................................13

Figure 8. Ecological response to changing climatic conditions following glacial retreat in Yellowstone 

National Park. ................................................................................................................................................14

Figure 9. Ecological response to changing climatic conditions following glacial retreat in the southern 

Rockies. ........................................................................................................................................................15

Figure 10. Ecological response to changing climatic conditions in the western portion of the Upper 

Columbia Basin following deglaciation. .........................................................................................................15

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....................16

Figure 12. Comparison of regional and global temperature reconstructions. ................................................19

Figure 13. Long-term aridity changes in the West as measured by the percent area affected by drought 

(PDSIb_1, thick black line), 95% boostrapped confidence intervals (light-blue dotted lines) and the 

long-term mean (thin horizontal black line)....................................................................................................20

Figure 14. Relationship between Glacier National Park summer drought, inferred winter snowpack, fire 

area burned, and glacial recession 1700 to present. .......................................................................................23

Figure 15. Habitats for Utah juniper in Wyoming and adjacent Montana. Black areas indicate extremely suitable habitats and gray areas mode.................................................................................................................25

Figure 16. 1100 years of drought history in the Green River Basin region of southwest Wyoming. ................26

Figure 17. Establishment and presence/absence of pinyon pine in relation to climate variability and elevation 

during the Holocene. .....................................................................................................................................27

Figure 18. Fire-related sedimentation in the South Fork Payette River (SFP) area in central Idaho and in 

Yellowstone National Park (YNP), Wyoming....................................................................................................28

Figure 19. Observed annual average North American surface temperature, 1950–2007 (left) and the time 

series of the annual values of surface temperature averaged over North America (right). ...............................31

Figure 20. Trends in April 1 SWE from snow course observations, 1960–2002...............................................32

Figure 21. Trends in winter mean wet-day minimum temperatures, 1949–2004. ..........................................33

Figure 22. Annual number of forest fires >1000 acres (total column height) in the northern Rockies 

(black area) and other western states. ...........................................................................................................36
vii

Contents

Figure 23. Comparison of variability and trends in western Montana (blue-green line) and Northern Hemisphere (dark blue line) annual average temperatures......................................................................................37

Figure 24. Average winter (Dec–Feb; top), spring (Mar–May; middle), and annual (bottom) minimum 

temperatures. ...............................................................................................................................................37

Figure 25. Idealized relationship between northern Rockies snowpack and streamflow anomalies with 

associated Pacific SSTs, atmospheric circulation, and surface feedbacks. .........................................................38

Figure 26. (a) Observed annual (previous July through current year June) precipitation for Wyoming Climate 

Division 1 (gray line) compared to precipitation estimates based on the stepwise regression model (black line). 

(b) The full stepwise version of reconstructed annual precipitation (black line) for 1173 to 1998. ...................39

Figure 27. Observed time series (1895–2007) of annually averaged precipitation departures area-averaged 

over the Upper Colorado drainage basin (top) and annual Colorado River natural flow departures at Lees 

Ferry in million acre-feet (bottom). ................................................................................................................40

Figure 28. Trends in average annual Pacific Northwest temperature, 1920–2000...........................................40

Figure 29. Trends in average annual precipitation, 1920–2000. .....................................................................41

Figure 30. Trends in April 1 SWE, 1950–2000. ..............................................................................................41

Figure 31. Modeled changes in annual mean precipitation minus evaporation (P–E) over the Southwest 

(125°W to 95°W and 25°N to 40°N, land areas only), averaged over ensemble members for each of the 19 

climate models participating in the Fourth Assessment Report (AR4) of the IPCC............................................43

Figure 32. Relationships between climate change, coping range, vulnerability thresholds, and adaptation. ...45

Figure 33. Schematic for a normal temperature distribution showing the effect on extreme temperatures 

when the mean temperature increases. .........................................................................................................45

Figure 34. Temperature and precipitation changes in North America projected for 2050 (2040–60 

average) by an ensemble of 22 climate models used in the IPCC AR4.............................................................47

Figure 35. Summer observed average temperatures and statistically downscaled projections for the northern 

and central U.S. Rockies and Greater Yellowstone Area (left) and Southern Rockies (right). ...........................49

Figure 36. June–August 20-year temperature projectionscentered on 2025 (left panels), 2050 (right 

panels) for Glacier National Park (top panels, elevation 1866m) and the Gallatin Mountains, Montana 

(elevation 2778m, bottom panels) for a 4-km grid cell (approximately 30 x 40 mile). .....................................50

Figure 37. June–August 20-year temperature projections centered on 2025 (left panels), 2050 (right panels) 

for Niwot Ridge, Colorado (top panels, elevation 3267m) and Clearwater, Idaho (bottom panels, elevation 

2467m) for a 4-km grid cell (approximately 30 x 40 mile). .............................................................................51

Figure 38. Estimate of (a) mean SWE, (b) mean soil moisture (June–August), and (c) mean potential 

evapotranspiration (June–August) for 1916–2006 and 2030–2059. ..............................................................53

Figure 39. Example of a scenario planning matrix. Each axis represents a critical driver of system change 

or a significant trend in the environment. ......................................................................................................57
viii

A Framework for Climate Analysis and Reporting for Greater Yellowstone and Rocky Mountain Networks Climate and Ecosystem Change in the U.S. Rocky Mountains and Upper Columbia Basin

Tables

Table 1. Climatic variation at different time scales and biotic response (modified from Overpeck et al. 2003)...2

Table 2. Examples of paleoenvironmental proxy data, spatio-temporal resolution, and range and type 

of reconstruction. ............................................................................................................................................5
ix

Contents Executive Summary

Paleoenvironmental records provide critical information on past climates and the 

response of ecosystems to climatic variability. Ecosystems have changed in a variety of 

ways as a result of past climate change, and 

they will continue to do so in the future. At 

large scales, climate governs the distribution 

of vegetation across the landscape and acts 

as a strong control of important biophysical 

conditions (e.g., extent of mountain glaciers) and ecosystem processes (e.g., area 

burned by fire). Paleoenvironmental and 

instrumental records from throughout the 

western United States suggest that vegetation 

response to climate change varies along a 

hierarchy of temporal and spatial scales, and 

the responses range from wholesale shifts in 

biomes to small adjustments in forest density 

or structure. Anticipating how ecosystems 

may respond to ongoing and future climate 

change requires an understanding of the climate–ecological linkages on all these scales 

as well as cross-scale interactions that lead to 

abrupt responses and regime shifts. 

Vegetation changes occurring on millennial time scales are related to changes in 

the seasonal cycle of solar radiation and its 

attendant effect on atmospheric circulation 

patterns and surface energy balances. After 

the ice sheets and local glaciers retreated 

17,000–12,000 cal yr BP (calibrated years 

before present, equal to the number of calendar years before 1950), paleoenvironmental records from the Pacific Northwest and 

Rocky Mountains reveal a sequence of vegetation changes as a result of increasing temperatures and effectively wetter conditions. 

Initially, deglaciated regions were colonized 

by tundra communities and the climate was 

colder and probably drier than at present. 

After 14,000 cal yr BP, warmer and wetter 

conditions allowed present-day conifer taxa 

to expand first in open parkland and later 

as closed forest communities. By 11,000 cal 

yr BP, closed subalpine forests were widespread. The early Holocene (11,000–7000 cal 

yr BP) was a time of warmer and drier summer conditions than at present. Warmer temperatures led to an upslope shift in conifer 

ranges and xerothermic shrub communities 

occupied valley bottoms. Summer drought 

led to higher fire frequencies than at present 

in many ecosystems. A gradual cooling and 

increase in effective moisture in the mid-Holocene (7000–4000 cal yr BP) was followed 

by relatively cool, moist conditions (4000 cal 

yr BP–present). 

Embedded with these millennial scale changes are centennial climate variations such 

as the Medieval Climate Anomaly (ca. AD 

950–1250, AD = Anno Domini) and the Little 

Ice Age (ca. AD 1400–1700). These variations 

had less dramatic impacts on vegetation, but 

records describe shifts in ecotone positions, 

including the upper and lower treeline, and 

disturbance regimes. Gridded tree-ring 

networks suggest that within these intervals 

were multidecadal “megadroughts” associated with tree mortality and fire. On annual 

to decadal scales, climate variations have led 

to disturbance events and shorter droughts. 

These events have shaped successional pathways, tree growth and mortality, and community structure. Case studies investigating 

ecological response to these changes provide 

important lessons for understanding how 

ecosystems may respond to ongoing and 

future environmental change: 

● The last century is an inadequate reference period for considering future climate 

change because it does not capture the 

range of natural climate variability that 

vegetation responds to or the magnitude 

of climate change projected for the near 

future. For example, managers often rely 

on the last several decades of fire occurrence as a baseline for managing different 

ecosystems in the West even though, because of fire suppression policies and fire 

elimination, fire activity of the last century 

is atypical of long-term historical patterns 

in much of the West and is unlikely to represent future conditions. Many vegetation 

types have evolved under a wide range 

of fire frequencies and intensities, calling 

into question the value of a static view of 

a fire regime, e.g., characterized by a mean 

fire return interval. To understand the full 

range of conditions that may be important 

for sustaining ecosystems in the future, 

long-term records of fire (e.g., centuries to 

millennia) provided by tree-ring and lake 

sediment data are essential.
x

Climate and Ecosystem Change in the U.S. Rocky Mountains and Upper Columbia Basin A Framework for Climate Analysis and Reporting for Greater Yellowstone and Rocky Mountain Networks

● Rapid climate transitions have occurred in 

the past and will likely occur in the future. 

In the past, the response of vegetation 

has been highly variable, suggesting an 

equally complex response to future climate change. Among the likely outcomes 

will be a highly individualistic response by 

different species, a reorganization of plant 

communities, and the likelihood of differential lags in the ability of species to stay in 

equilibrium with climate change.

● Climate variability at large scales is often 

expressed in complex and asynchronous 

patterns across the U.S. Rocky Mountains and Upper Columbia Basin, largely 

because of interactions with topography 

and other sources of spatial heterogeneity. This can result in nearby communities 

showing different directions of change 

(e.g., precipitation regimes changing in opposite directions at different 

elevations).

● Many western ecosystems represent 

assemblages that formed as a result of a 

specific sequence of climate conditions 

during the last several thousand years. In 

particular, many middle and late-successional communities were established during the colder, wetter Little Ice Age and 

would be unlikely to form under current 

climate and land use conditions. Likewise, 

restoration to historical baselines is at best 

challenging and in most cases impossible.

Temperatures are projected to warm 1–5°C 

for much of the West by 2100, accompanied 

by declines in snowpack, earlier spring 

snowmelt, and reduced late-summer flows. 

While projections for future precipitation 

are less certain, increased precipitation is 

unlikely to offset increased evapotranspiration associated with even modest warming 

(e.g., 1–2°C), particularly during the summer. Consequently, drought is projected 

to increase in frequency and intensity over 

the next several decades, particularly in the 

Southwest and southern U.S. Rockies.

Investigations of the past suggest that we 

should expect dynamic and rapid ecosystem 

response to changing climate conditions, 

but in ways that may be difficult to predict. 

Paleoenvironmental records illustrate that 

while existing biomes have experienced 

distribution shifts, they have been resilient to 

climatic change across multiple time scales 

(e.g., decadal to millennial). They also suggest, however, that we should anticipate increased extreme and unforeseen disturbance 

synergisms; increased tree mortality, shifts 

in treeline position, and non-native plant 

invasions; and ultimately changes in plant 

community composition, structure, and 

function that may constitute novel vegetation 

assemblages. This poses significant challenges for developing management plans, but 

should not deter an adaptive management 

approach that allows for reassessment and 

modification of management strategies in the 

coming decades. By using scenario planning, 

managers can consider a wide range of possible future conditions to examine potential 

trajectories of ecological change. Managing 

for future conditions will, at the least, involve 

a continuation of well-established resource 

management and conservation practices.
xi

Contents Acknowledgements

We would like to acknowledge all the participants of the 2009 Rocky Mountain Inventory 

& Monitoring Technical Committee Meeting. 

A NPS steering committee including Judy 

Visty, Kathy Tonnessen, Lisa Garrett, Tom 

Rodhouse, Tom Olliff, David McWethy, Stacey Ostermann-Kelm, Bruce Bingham, and 

Penny Latham provided useful comments 

throughout the process. An interagency committee organized through the Great Northern Landscape Conservation Cooperative 

also provided guidance and advice throughout the writing and development process. 

This committee included: Yvette Converse 

(USFWS), Mike Britten (NPS), Tom Olliff 

(NPS), Molly Cross (Wildlife Conservation 

Society), Steve Gray (WY state climatologist), 

Beth Hahn (USFS), Virginia Kelly (Greater 

Yellowstone Coordinating Committee), Tim 

Mayer (USFWS), Jim Morrison (USFS), Stacey Ostermann-Kelm (NPS), Greg Pederson 

(USGS), David Wood (BLM), Andrea Ray 

(NOAA), and Lou Pitelka (NEON). 

Funding was provided by the Greater Yellowstone Coordinating Committee and the 

NPS Inventory & Monitoring Program. Conceptual models were developed and created 

by Robert Bennetts. Maps were created by 

Meghan Lonneker.

Any use of trade, product, or firm names is 

for descriptive purposes only and does not 

imply endorsement by the U.S. Government.

1

Introduction

Figure 1. Location 

of study area. Climate region boundaries modified from 

Littell et al. (2009a), 

Kittell et al. 2002, 

and 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. (Figure 

created by M. Lonneker, NPS, 2010.)

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 (Solomon et al. 

2007). The controls of climate vary at different time scales (i.e., millennial, centennial, 

multidecadal, decadal, annual and interannual), and the ecological response to climate 

change varies accordingly. Paleoecological 

data show that ecosystems in the West have 

undergone significant and sometimes rapid 

changes since the Last Glacial Maximum (ca. 

20,000 years ago), and the biotic assemblages 

observed today are relatively recent phenomena (Thompson et al. 1993; Whitlock and 

Brunelle 2007; Jackson et al. 2009a). The future is also likely to be characterized by rapid 

biotic adjustment, including the possibility 

of novel assemblages as species respond 

individualistically to climate change (MacDonald et al. 2008; Williams et al. 2007). 

Understanding the drivers and rates of past 

climate change 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 (MacDonald et al. 2008; Shafer et al. 2005; 

Whitlock et al. 2003; Overpeck et 

al. 2003; Swetnam and Betancourt 

1999). 

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 important 

role over the coming decades. 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) 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, 

snow water equivalent [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 National 

Park Service (NPS), the U.S. Forest Service 

and the U.S. Fish and Wildlife Service, the 

geographic scope of this report encompasses 

the core regions of the Great Northern 

Landscape Conservation Cooperatives and 

the NPS high-elevation parks of the Rocky 

Mountains. For organizational purposes, 

the report divides the study area into four 

regions: the northern U.S. Rocky Mountains, the central U.S. Rocky Mountains and 

Greater Yellowstone Area, the southern U.S. 

Rocky Mountains, and the Upper Columbia 

Basin (fig. 1).

Chapter 1: Introduction 

0 250 500 Kilometers 750

The data was compiled from the best sources available, so various errors from the sources may be inherent on the map. All boundaries and 

features therein should be treated as such. For official LCC and boundary information: http://www.fws.gov/GIS/data/national/index.htm

Map created by NPS- UCBN; February 2010

Glacier 

National Park 

(GLAC)

Nez Perce

National Historical Park 

(NEPE) Big Hole 

National Battlefield 

(BIHO)

Craters of the Moon 

National Monument 

(CRMO)

Yellowstone

National Park 

(YELL)

City of Rocks

National Reserve

(CIRO)

Grand Teton 

National Park 

(GRTE)

John D. Rockefeller 

Jr. Memorial Parkway

(JODR)

Rocky Mountain 

National Park 

(ROMO)

Florissant Fossil Beds

National Monument 

(FLFO)

Great Sand Dunes 

National Park 

(GRSA)

Bighorn Canyon 

National Recreation Area 

(BICA)

Great Northern LCC

NPS High Elevation Parks

Climate Regions

Canadian Rockies

Central U.S. Rockies and GYA

Northern U.S. Rockies

Southern U.S. Rockies

Upper Columbia Basin
2

Climate and Ecosystem Change in the U.S. Rocky Mountains and Upper Columbia Basin A Framework for Climate Analysis and Reporting for Greater Yellowstone and Rocky Mountain Networks

The synthesis is organized into four sections: 

(1) biophysical responses and drivers of climate changes occurring on multi-centennial 

to millennial scales during the last 20,000 

years; (2) biophysical responses and drivers 

of climate change on annual to centennial 

scales over the last 2000 years; (3) the last 

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 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 about climatic change and ecological 

response, and provides a planning approach 

designed to address a wide range of potential 

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 ecosystemrelated research available for the entire study 

region. 

1.1 Climate controls and variability 

at different spatio-temporal scales

The influence of variation in climate on 

ecosystems changes at different spatial and 

temporal scales. Understanding the potential 

influence of climate on biophysical processes 

often requires local to synoptic (regional or 

larger) information on the type and magnitude of forcings (mechanisms driving 

changing conditions), e.g., relative humidityrelated changes versus changes in solar 

radiative output, long-lived greenhouse gasses, and ocean-atmosphere interactions and 

local albedo (low albedo = low reflectance or 

dark ground surface such as pavement; high 

albedo = high reflectance and lighter surface 

such as snow) along with an understanding 

of the sensitivity of the ecosystem property 

that is being measured (Overpeck et al. 2003; 

Webb and Bartlein 1992) (Table 1). It is also 

important to recognize the hierarchical 

nature of climate variation and change, and 

that short-term (i.e., daily to interannual) 

events are superimposed on longer ones, 

amplifying or dampening the magnitude of 

the underlying physical controls that influence ecosystem dynamics. 

At the longer temporal range of major 

changes in Earth’s climate system, variations 

on scales of 10,000 to 100,000 years are attributed to slowly varying changes in Earth’s 

orbit known as Milankovitch Cycles (i.e., 

changes in Earth’s precession, tilt, and obliquity [Hays et al. 1976, Berger 1978, Berger 

Table 1. Climatic variation at different time scales and biotic response (modified from Overpeck et al. 2003).

Frequency 

and Scale 

of Variation 

(years)

Kind of 

Variation Forcing Mechanism Biotic Response

Millennial

(1000–10,000) 

Deglacial and 

postglacial 

variations

Ice sheet size, insolation, 

trace gases, regional oceanatmosphere-ice interactions

Species migration, range expansion 

and contraction; community reorganization and establishment, species 

extirpation/extinction

Interdecadal–

centennial

(10–100)

Decadal and 

centennial 

anomalies

Internal variations in the 

climate system, solar variability, volcanism

Shifts in relative abundance and 

composition of different taxa through 

recruitment, mortality, and succession

Annual–interannual

(<10)

Storms, 

droughts, 

ENSO events

Internal variations in the 

climate system, solar variability, volcanism

Adjustments in physiology, life history 

strategy, and succession following 

disturbance
3

The Protocol Framework

and Loutre 1991], fig. 2). The net result of 

slowly varying changes in Earth’s orbit have 

included multiple glacial and inter-glacial 

cycles driven by changes in global average 

temperatures over the past several million 

years. These temperature changes are initiated by changes in the seasonal and annual 

cycle of insolation (incident solar radiation; 

the amount of solar radiation received on a 

given surface area during a given time) over 

the high latitudes of the Northern Hemisphere that result in substantial positive 

feedbacks from changing concentrations of 

atmospheric greenhouse gasses (Vettoretti 

Introduction

Figure 2. Primary drivers of climate and resulting climate variations at millennial, centennial, 

and interannual scales. (A) Temperature reconstruction from the central Greenland (GISP2) 

ice core record and the forcing mechanisms thought to influence variation during the glacial 

period (ca. 49,000–12,000 cal yr BP). The Laurentide ice sheet in North America began to recede and climate warming commenced ca. 17,000 cal yr BP. Gray bands indicate two abrupt 

climate changes: the Younger Dryas Chronozone ca. 12,900–11,600 BP, and the 8200 cal yr 

BP cool event (8.2 ka event); (B) 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 the instrumental record (blue) for the past 2000 years (Viau et al. 2006). Continental patterns of drought and interannual and decadal climate variability are associated 

with the Medieval Climate Anomaly (ca. 950–1250) and fewer fires during the Little Ice Age 

(ca. 1400–1700). (Time spans 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 global temperature 

anomalies (black line, based on 1900–2000 mean) from HadCRUT3v instrumental reconstruction (Brohan et al. 2006) and ocean-atmosphere variability. Magenta line represents 

the Multivariate El Niño–Southern Oscillation (ENSO) Index, which is based on six observed 

ocean-atmosphere variables. Positive values of the index depict El Niño events (Wolter and 

Timlin 1993, 1998). Source: NOAA Earth System Research Laboratory.

45 40 35 30 25 20 15 10 5 0

-50

-40

-30

-20

0 200 400 600 800 1000 1200 1400 1600 1800 2000

-1

-0.5

0

0.5

1

ka yrs. BP

yrs. AD

T ( ̊C)

(A)

(B)

Younger Dryas

8.2 ka event

Ice-sheet dynamics/ocean circulation/insolation Insolation/volcanism/

ocean-atmosphere interactions

Medieval Warm Period

Little Ice Age

Insolation/volcanism/ocean-atmosphere interactions anthropogenic

influence

1950 1960 1970 1980 1990 2000 2010

0 0

0.5

-0.5

2

-2

Standardized Departure

Multivariate ENSO Index

(C) Temp. anomaly ( ̊C) Temp. anomaly ( ̊C)

yrs. AD

ka

tio

-a

Temperature Anomaly
4

Climate and Ecosystem Change in the U.S. Rocky Mountains and Upper Columbia Basin A Framework for Climate Analysis and Reporting for Greater Yellowstone and Rocky Mountain Networks

and Peltier 2004; Bond et al. 2001; Kutzbach 

and Guetter 1986; Kutzbach et al. 1998, 

1993).

On multi-millennial 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 

results in ocean-ice-atmosphere interactions 

that drive changes in atmospheric circulation 

patterns (i.e., the position of westerlies and 

preferential positioning of storm tracks [fig. 

2a]), resulting in major changes in ecosystem 

distribution and structure. For example, as 

Northern Hemisphere summer insolation 

increased and the ice sheets and glaciers began to retreat, seasonal storm tracks shifted 

north and paleoecological records show 

widespread reorganization of plant communities throughout the West (e.g., MacDonald 

et al. 2008; Thompson et al. 1993; Bartlein 

et al. 1998; Whitlock and Brunelle 2007; 

Jackson et al. 2005; Betancourt et al. 1990). 

These large-scale and long-term changes 

in insolation and ice cover are important 

features of Earth’s climate dynamics because 

they influence the persistence and strength 

of storm tracks, subtropical high-pressure 

systems, ocean-land temperature gradients, 

and consequently interannual 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 (ca. 11,000–7000 

cal yr BP) led directly to increased summer 

temperatures and indirectly to a strengthening of the Pacific subtropical high-pressure 

system off the northwestern United States, 

effectively intensifying summer drought in 

the region (Bartlein et al. 1998). Records of 

early Holocene glacier dynamics, lake levels, 

aeolian activity (blowing dust), vegetation, 

and fire show the ecological effects of this 

increased summer insolation 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).

On shorter and perhaps more managementrelevant time scales, climate variations at 

interdecadal to centennial scales are related 

to changes in solar activity, volcanism, seasurface temperature, and pressure anomalies in both the Atlantic and Pacific oceans. 

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 centennial-scale climate variation due to 

the relatively warm and dry conditions that 

prevailed across the western United States 

from approximately 900 to 1300. The West 

experienced substantially reduced streamflows (Meko et al. 2007), shifts in the upper 

treeline (Rochefort et al. 1994; Fall 1997), 

and increased fire activity (Cook et el. 2004). 

Though the exact causes of the MCA are still 

debated, the prevailing evidence suggests 

that it was driven by changes in solar activity, 

volcanism, and perhaps sustained La Niñalike conditions in the tropical Pacific (Mann 

et al. 2009). At decadal to interdecadal 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 modes 

of interdecadal variability include the Pacific 

Decadal Oscillation (PDO; see Mantua et al. 

1997) and the Atlantic Multidecadal Oscillation (AMO; see Enfield et al. 2001, Regonda 

et al. 2005). Decadal climate shifts 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 in proxy-based reconstructions of 

precipitation and streamflow (fig. 2c; e.g., 

Gray et al. 2003). These events are often regional to subcontinental in scale and initiate 

and terminate 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.

The El Niño–Southern Oscillation 

(ENSO) is a major global control of both 
5

The Protocol Framework

temperature and moisture patterns (fig. 2c). 

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 

which typically last 6–18 months and reoccur every 2–7 years (Bjerknes 1969; Cane 

and Zebiak 1985; Graham and White 1988; 

Horel and Wallace 1981; Philander 1990; 

Chang and Battisti 1998). The magnitude 

of these sea surface temperature anomalies 

varies, but they typically exert 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 dry conditions across the northwestern United States 

and a southerly displacement of the winter 

storm track (i.e., the jet stream), resulting in 

cool wet conditions across the Southwest 

(McCabe et al. 2004; Enfield 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 Southwest (e.g., Swetnam 

and Betancourt 1998), with an important 

but somewhat attenuated spatial signature 

across the Northwest (e.g., Dettinger and 

Ghil 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 seasons to 

millennia. Knowledge of climate drivers on 

all time scales is necessary to identify temporal and spatial dimensions of future changes 

and possible ecological responses to those 

changes.

1.2 How can understanding 

climatic variability inform 

management? 

Knowledge of natural variations in climate at 

different scales provides a context for understanding how communities and individual 

species might respond to current and future 

rates of change (Table 2). For example, past 

changes in fire regimes have been largely 

driven by large-scale climate changes on 

millennial, centennial, decadal and annual 

scales (Whitlock et al. 2003; Kitzberger et 

al. 2007; Littell et al. 2009a; 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 conditions as opposed to maintaining 

current or past conditions (Whitlock et al. 

2010). 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. 2006, 2004; 

MacDonald et al. 2008; Lyford et al. 2003).

Table 2. Examples of paleoenvironmental proxy data, spatio-temporal resolution, and range and type of reconstruction.

Proxy Source Temporal Resolution Spatial Resolution Temporal Range Type of 

Reconstruction

Tree 

growth

Tree cores High (seasonal/annual) High 100–1000s yrs Temperature, 

moisture

Charcoal Lake/ peat 

sediments

High to moderate 

(annual–decadal)

High to moderate (1 to 

several km2

)

High (many millennia) Fire

Pollen Lake/peat 

sediments

Moderate (multidecadal–centennial)Moderate (several km2

) High (many millennia) Vegetation

Oxygen 

Isotopes

Corals, tree, 

lake, ocean 

or ice cores

High to moderate 

(annual–decadal)

Moderate to low High to very high 

(100,000+)

Temperature, ice 

volume

Introduction
6

Climate and Ecosystem Change in the U.S. Rocky Mountains and Upper Columbia Basin A Framework for Climate Analysis and Reporting for Greater Yellowstone and Rocky Mountain Networks

Obtaining lake 

sediment core, Rocky 

Mountain National 

Park. Photo courtesy 

P. Higuera.

Reconstructing past climatic conditions 

and the associated ecological response 

involves a number of direct and indirect 

measurements. 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 that respond to changes in climate through changes in their growth rate, 

abundance, or distribution as recorded in 

living or fossil specimens or assemblages 

of organisms. Each proxy indicates past 

change at different spatial and temporal 

scales and resolutions. 

Lake sediment, tree-ring cores, and packrat 

middens are three of the primary proxies 

used for reconstructing past conditions for 

the western United States (Whitlock and 

Larsen 2001; Fritts and Swetnam 1989; 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 

intervals throughout the core. Most lakes in 

the Northwest were formed during deglaciation, and therefore 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 intervals (e.g., every 0.5–1 cm) 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 

plotted as a diagram.

 The reconstruction of past vegetation and 

climate from pollen percentages 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 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 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 (particles/cm3

), 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, which provide records of past 

change at centennial and millennial scales, 

have several features that make them well 

suited for climate reconstruction, such 

as ease of replication, wide geographic 

availability, annual to seasonal resolution, 

and accurate, internally consistent dating. 

Reconstructing past environments using proxy data
7

The Protocol Framework

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 

centennial to millennial change. Tree growth 

is highly sensitive to environmental changes 

and therefore tree-ring records are powerful tools for the investigation of annual to 

centennial variations. Tree-ring chronologies are used to reconstruct past 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 on 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. This cross-dating also allows 

tree-ring series from ancient dead wood 

(found in 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 temperature.

Middens left by woodrats of the genus 

Neotoma also provide long-term records 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 

vegetation and climate. Middens located 

in caves or under rock ledges that provide 

protection from water are especially well 

preserved. The plant and animal parts from 

an excavated midden are dissected and 

identified, and then dated using radiocarbon 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 period. A reconstruction of vegetation typically includes the area within 30 to 

100 meters (33–109 yd) surrounding a site 

(Betancourt et al. 1990).

Obtaining tree-ring 

core, Crazy Mountains, Montana. 

Photo D. McWethy.

Packrat midden. Photo courtesy USGS.

Reconstructing past environments

9

2.1 Drivers of millennial-scale 

climate variation

The climate variations of the last 20,000 

years occurring on millennial scales are best 

understood through model-based simulations that look at the regional response to 

large-scale climate changes and paleoenvironmental data that measure specific 

components of climate change. Broad-scale 

climate variations were described by Bartlein 

et al. (1998) using the NCAR Community 

Climate Model (CCM1; 4.4o latitude by 7.5o 

longitude spatial resolution, mixed-layer 

ocean, crude depiction of western cordillera topography). The simulation provided 

estimates of climatic conditions over six 

discrete periods (21,000 cal yr BP=full glacial 

period with full-sized ice sheets; 16,000 to 

14,000 cal yr BP =late Glacial period with 

shrinking ice sheets; 11,000 cal yr BP = early 

Holocene insolation maximum; 6000 cal yr 

BP= mid-Holocene transition; and present). Other recent regional-scale modeling 

studies have provided better temporal and 

spatial resolution with more realistic topography (Hostetler 2009; Hostetler et al. 2003). 

Results from these efforts are summarized 

below. 

At the time of the Last Glacial Maximum (ca. 

21,000 cal yr BP), the large Laurentide and 

Cordilleran ice sheets strongly influenced 

climatic conditions in the western United 

States (Bartlein et al. 1998; fig. 3), depressing 

temperatures approximately 10°C in areas 

south of the ice sheets and steepening the 

latitudinal temperature gradient. The presence of the large ice sheets also displaced the 

jet stream south of its present position, greatly reducing winter precipitation in the northwestern United States and Canada while increasing precipitation across the Southwest. 

Another element of the full-glacial climate 

was stronger than present surface easterlies 

related to a strong high-pressure system that 

persisted over the ice sheets. The presence of 

this strong high-pressure system steepened 

the latitudinal temperature gradient and 

weakened the westerly storm tracks, resulting in colder and effectively drier conditions 

across the Northwest and cold wet conditions in the Southwest (Hostetler et al. 

2000; Bartlein et al. 1998; Thompson et al. 

1993).

2.2 Glacial–Holocene transition

During the Glacial–Holocene transition 

(16,000–11,000 yr BP), solar insolation over 

high-latitude Northern Hemisphere landmasses increased, peaking ca. 11,000 yr BP 

when summer insolation was 8.5% higher 

and winter insolation was 10% lower than at 

present at 45oN latitude. One consequence 

was a northward shift of winter storm tracks 

from their full-glacial position, bringing 

wetter winter conditions to the Northwest 

while the Southwest became increasingly 

dry (Bartlein et al. 1998). Increasing summer insolation resulted in warmer growing 

season temperatures, causing alpine glaciers 

and ice sheets to melt rapidly. (See fig. 4 for 

an example of modern air mass circulation.)

12000 10000 8000 6000 4000 2000 0

-42

-40

-38

-36

-34

-32

0

0.2

0.4

0.6

0.8

8.2 ka event

Area of the Laurentide Ice Sheet

Fraction of LGM area

GISP2 δ18O

(per mille V-SMOW)

Age (cal yr BP)

Younger

Dryas

Chapter 2: The last 20,000 years of 

environmental change in the western United States

Figure 3. Area of 

the Laurentide ice 

sheet (top panel) 

and central Greenland temperature 

reconstruction 

(bottom panel). Ice 

sheet area estimated from Dyke and 

Prest (1987) and 

Barber et al. (1999); 

oxygen isotope record (bottom panel, 

black line) associated with variations 

in Northern Hemisphere temperature 

(higher isotope 

values represent 

warmer temperatures) from GISP2 

(Stuiver et al. 

1995). The Younger 

Dryas Chronozone 

was an abrupt 

cooling ca. 12,900–

11,600 BP which 

represented a 

temporary reversal 

in warming during 

the Pleistocene–

Holocene transition 

(Alley et al. 1993). 

(Figure modified 

from Shuman et al. 

2002).
10

Climate and Ecosystem Change in the U.S. Rocky Mountains and Upper Columbia Basin A Framework for Climate Analysis and Reporting for Greater Yellowstone and Rocky Mountain Networks

At the end of the Pleistocene (the last 2.5 

several million years of repeated glaciations 

prior to the Holocene), much of the Northern Hemisphere experienced an abrupt 

cooling known as the Younger Dryas Chronozone (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 YDC is less obvious 

in the western United States, where most 

paleoenvironmental data show little or no 

response 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). In 

some areas, however, there is evidence of a 

re-advance of mountain glaciers (Osborn 

and Gerloff 1997; Reasoner and Huber 1999; 

Friele and Clague 2002; Menounos and 

Reasoner 1997) and vegetation changes that 

indicate cooling, including a lower treeline 

in central Colorado (Reasoner and Jodry 

2000) and changes in the isotopic signature 

in speleothem data. Temperatures decreased 

as much as 5–10°C (9–18°F) in Greenland, 

but pollen data for the northwestern United 

States indicate much more moderate cooling (0.4–0.9°C [0.7–1.6°F]; 

Reasoner and Jodry 2000). The 

YDC ended rapidly with warming 

of ~7°C (13°F) in Greenland occurring within one to several decades 

(Alley 2000). Consequently, the 

period is closely scrutinized as an 

example of abrupt climate change 

(Alley et al. 2003).

A similar abrupt cooling occurred 

ca. 8200 cal yr BP (figs. 2–3) when a 

large influx of fresh water disrupted 

circulation in the North Atlantic, 

causing cooling that lasted several 

centuries in the North Atlantic 

region (Le Grande et al. 2006; Alley 

and Ágústsdóttir 2005). Like the 

YDC, the 8.2 ka event is not registered at many sites in the western 

United States, 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, Utah (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. They also show a variable 

signal in regions distal to the North Atlantic origin, so that some sites in the western 

United States show a response while others 

do not.

Greater summer insolation (8% above present) and lower winter insolation (8% lower) 

in the early Holocene (11,000–7000 cal yr 

BP) profoundly affected the climate and 

ecosystems of the western United States. 

Increased summer insolation led to warmer 

temperatures throughout the region during 

the growing seasons, while winters were likely colder than at present. Model simulations 

show that increased insolation led indirectly 

to a strengthening of the eastern Pacific 

subtropical high-pressure system which suppressed summer precipitation in much of the 

Northwest. At the same time, it strengthened 

inflow of moisture from the Gulf of California to the Southwest and the southern and 

central Rocky Mountain region (Whitlock 

and Bartlein 1993), resulting in greater than 

Figure 4. Primary 

air masses that 

influence the study 

area. Ice-sheet 

dynamics and 

ocean-atmosphere 

interactions have 

altered large-scale 

air mass circulation 

patterns throughout 

the late Glacial and 

Holocene, influencing temperature and 

precipitation. The 

U.S. Rocky Mountains lie in a transition zone where 

many of these air 

masses interact, 

strongly influencing the local and 

regional climate. 

Changes in the position and strength 

of these air masses, 

driven by largescale changes in 

the climate system, 

are responsible for 

climatic variations 

on different time 

scales and act as an 

important control on 

vegetation. (Modified from Ahrens 

2008 and Aguado 

and Burt 2010.)
11

Contents

present summer precipitation. East of the 

Rockies, increased summer precipitation 

was likely offset by increased temperatures 

and rates of evapotranspiration, making 

conditions effectively drier than at present, 

which is consistent with low lake levels and 

dune activation (Shuman et al. 2009; Stokes 

and Gaylord 1993). 

Today the West is characterized by summer-dry areas under the influence of the 

subtropical high (i.e., the Northwest) and 

summer-wet areas where summer precipitation reflects monsoon activity (i.e., the 

Southwest). 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 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 mid-Holocene (ca. 7000–4000 

cal yr BP) and the late Holocene (4000 cal yr 

BP–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.

2.3 Mid-Holocene transition

In the Pacific Northwest, including the 

Columbia Basin, the mid-Holocene was a 

transition period between the warm, dry, 

early Holocene and the cooler, wetter, late 

Holocene. In British Columbia, Hebda and 

Mathewes (1984) call this period the mesothermal and trace the expansion of hemlock (Tsuga heterophylla) and Douglas fir 

(Pseudotsuga menziesii) during this period. 

This signal is also evident in the northern 

Rockies and perhaps as far south as the 

Greater Yellowstone Area. 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 

Rockies experienced drier conditions during 

the mid-Holocene (ca. 6000 cal yr BP) than 

at present (Shuman et al. 2009), similar to 

climatic conditions across the Great Plains. 

Anomalously dry conditions for the central 

and southern Rockies also contrast with wetter-than-present conditions in the 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; variations 

in the seasonal insolation cycle imposed local 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).

2.3.1 Southern Canadian and 

Northern U.S. Rocky Mountains

Prior to 14,000 cal yr BP, much of the 

southern Canadian and northern U.S. 

Rocky Mountains were glaciated (fig. 5). 

The first plant communities to colonize 

Figure 5. Ecological 

response to changing climatic conditions following 

glacial retreat in the 

southern Canadian 

and northern U.S. 

Rockies. Derived 

from MacDonald 

1989 and Reasoner 

and Hickman 1989 

(Whitlock unpublished).

Southern Canadian and Northern U.S. Rocky Mountains

Region

glaciated

Elevation (m)

3000

2000

1000

Ice

Picea-Pinus contorta-Populas

forest

Picea-Abies

forest

Alpine tundra

3000

2000

1000

Grassland with Pinus flexilus

Open PiceaPinus contorta forest

Pinus-Alnus forest

Shrub-herb

tundra

Ice

3000

2000

1000

Open Pinus forest

Shrub-herb tundra

Ice

Late Holocene

ca. ~4 ka yrs BP to present

Early Holocene

ca. ~11-7 ka yrs BP

Late Glacial

ca. ~14-11 ka yrs BP

> 14,000 cal yr BP

The last 20,000 years of environmental change
12

Climate and Ecosystem Change in the U.S. Rocky Mountains and Upper Columbia Basin A Framework for Climate Analysis and Reporting for Greater Yellowstone and Rocky Mountain Networks

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-thanpresent 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 at present. Low-elevation forests were characterized by spruce and 

lodgepole pine (Pinus contorta), and valley 

floors supported open grassland with limber 

pine. The upper treeline was at least 90 meters above present during the early- and midHolocene (ca. 8500–4000 cal yr BP). Low 

and middle elevations experienced the greatest fire activity during the early Holocene 

(ca. 9000–8000 cal yr BP) (MacDonald 1989) 

and forest composition resembled modern 

subalpine spruce-fir (Picea-Abies) forest. 

In the late Holocene, summer (growing) 

conditions became cooler, glacial activity 

increased, and the upper treeline decreased 

substantially (Reasoner and Hickman 1989). 

Cooler and moister conditions led to tundra 

at high elevations, spruce and fir forests at 

mid-elevations, and spruce, lodgepole pine, 

and aspen at low elevations. 

Ecosystem response to centennial-scale 

climatic variations is evident from a 3800-

year history of climate, 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 kilometers 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 cal yr BP, an abrupt 

rise in lake levels coincided with a transition 

from steppe and pine forest to pine forestwoodland 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 from 2200 to 1200 cal yr BP, 

and increases in grass, pine, and sagebrush 

and declines in Douglas fir and larch led 

to the development of a steppe/parkland/

forest mosaic ca. 700 cal yr BP (Power et al. 

2006; Stevens et al. 2006). Increases in grass 

and sagebrush in the late 19th and early 20th

centuries coincided with human activities. 

Notable climatic events during this period 

include a long, intense drought ca. 1140 

following a wetter period from 1050 to 1100 

(Stevens et al. 2006).

These shifts in vegetation were accompanied 

by pronounced changes in fire patterns as 

evident in the charcoal record (fig. 6). Intervals dominated by forests coincide with high 

magnitude and frequent fires (e.g., stand replacing fires), periods dominated by steppeparkland vegetation are associated with 

smaller and less frequent fires, and a decline 

in charcoal deposition in the last century 

likely reflects the impact of fire suppression 

(fig. 6). The Foy Lake record demonstrates 

the impacts of centennial-scale climate variations and their associated ecosystem response during the last two millennia. While 

relatively modest changes in vegetation cover 

occurred after the conifer forests were established ca. 2700 cal yr BP, multidecadal shifts 

in climate are evident in the fire reconstruction for the last several millennia.

2.3.2 Central U.S. Rocky Mountains 

and the Greater Yellowstone Area

Pollen records from Yellowstone and Grand 

Teton national parks show the nature of the 

biotic change that occurred in conjunction 

with broad climatic changes at different elevations (fig. 7, Whitlock 1993). Deglaciation 

(ca. 17,000–14,000 BP) was followed by colonization of tundra vegetation which, during 

the late Glacial and into the mid-Holocene, 

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 early Holocene, pine, juniper 
13

Contents

Figure 7. Ecological response to changing climatic conditions following glacial retreat in the 

Central Rockies and the summer-dry region of the Greater Yellowstone Area (derived from 

Whitlock 1993). The figure illustrates that tundra occupied most of the area >14 ka BP, but is 

only present at high elevations in the late Glacial and late Holocene.

Figure 6. Changes in pollen percentages for different plant taxa during the last 3700 years. 

Black and gray silhouettes (5x exaggeration) represent changes in the proportion of pollen for individual taxa over time as a percentage of all terrestrial pollen. The arboreal pollen 

(AP) and non-arboreal pollen (NAP) ratio includes all upland tree pollen divided by all nonaquatic shrubs and herbs. Fire frequency is recorded as number of episodes per 1000 years. 

Examples of vegetation response to climate change include a distinct change from steppe 

vegetation to forest accompanied by increased fire frequency ca. 2125 cal yr BP and a return 

to more open vegetation ca. 750 cal yr BP. Source: Power et al. 2006; reprinted with permission. Elevation (m)

Late Holocene

ca. ~4 ka yrs BP to present

Early Holocene

ca. ~11-7 ka yrs BP

Late Glacial

ca. ~14-11 ka yrs BP

> 14,000 cal yr BP

Central U.S. Rocky Mountains and Greater Yellowstone Area

1750

2000

2250

2500

2750

3000 Tundra

Picea-AbiesPinus forest

Pinus contorta forest

Pseudotsuga forest

Pinus flexilus forest

Steppe

Central Plateau

1750

2000

2250

2500

2750

3000

Tundra

with Betula,

Juniperus, Salix

1750

2000

2250

2500

2750

3000

Pinus-Juniper-Betula forest

Pinus contorta forest

with Pseudotsuga

and Populus

Subalpine

forest

1750

2000

2250

2500

2750

3000 Tundra

Grassland

Picea

parkland

The last 20,000 years of environmental change
14

Climate and Ecosystem Change in the U.S. Rocky Mountains and Upper Columbia Basin A Framework for Climate Analysis and Reporting for Greater Yellowstone and Rocky Mountain Networks

(Juniperus), and birch (Betula) were present 

at low elevations, whereas lodgepole pine, 

Douglas fir, and aspen (Populus) established 

and characterized mid-elevation forests. 

Subalpine forests (Picea and Abies) expanded 

their ranges to higher elevations, and the upper treeline was at an elevation similar to that 

of the present. Decreased summer insolation 

in the late Holocene (ca. 3000–4000 BP) led 

to cooler, wetter conditions. Sagebrushsteppe 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 the upper treeline to an elevation close to 

its present position.

2.3.3 Cygnet Lake

An examination of charcoal, pollen, and 

climate conditions from central Yellowstone 

provides an example of the linkage between 

fire and climate even in the absence of vegetation change (fig. 8; Whitlock 1993). Pollen 

from Cygnet Lake indicates the area was 

dominated by a tundra community of sagebrush and grass (Poaceae) when cool and 

wet conditions prevailed prior to 12,000 BP, 

after which the vegetation was dominated by 

lodgepole pine, which is favored in infertile 

rhyolite soils. The last 12,000 years reveal 

little change in the vegetation at Cygnet Lake 

compared to 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 Cygnet Lake fire record 

shows the effect of early Holocene drought 

on fire regimes; the charcoal data suggests 

rising fire activity in response to increasing summer insolation. Fire return intervals 

were 75–100 years between 11,000 and 7000 

cal yr BP, and cooler wetter conditions after 

7000 cal yr BP coincide with decreasing fire 

frequency. The present fire return interval 

of 200 to 400 years was reached in the late 

Holocene.

2.3.4 Southern U.S. Rocky Mountains

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 and wetter conditions (2–5°C [4–9°F] cooler and 

7–16 cm [3–6"] wetter than at present) 

supported tundra vegetation at high elevations (~300–700 m [984–2,297'] below the 

present treeline) and spruce parkland at low 

elevations (fig. 9, Fall 1997). During the late 

Glacial, fir (Abies) increased in abundance 

and subalpine forest was established at low 

and middle elevations (Fall 1997). Summer 

insolation increased during the early Holocene, and warmer temperatures allowed 

subalpine forests to expand 

above their present elevation. 

While the central and northern 

U.S. Rockies generally experienced warm dry conditions 

during the early Holocene, the 

southern Rockies experienced 

warmer, wetter growing seasons 

driven by a more intense North 

American monsoon (Thompson 

et al. 1993). Markgraf and Scott 

(1981) recorded an upslope 

advance of subalpine forests due 

to warmer conditions and an 

expansion of pine forests at both 

the lower and upper treeline facilitated by warmer and still wet 

conditions. Similarly, Fall (1997) 

found that from 9–4 ka BP, warm 

summers (mean temperatures 

Figure 8. Ecological response to 

changing climatic 

conditions following glacial retreat 

in Yellowstone National Park. Pollen 

and charcoal diagram from Cygnet 

Lake, central Yellowstone (modified 

from Millspaugh et 

al. 2000).
15

Contents

1.9°C [3.4°F] above present) facilitated the 

expansion of forests of spruce and fir (Picea 

engelmannii and Abies lasiocarpa) upward to 

an elevation of almost 4,000 meters (13,123'), 

at least 300 meters (948') higher than today 

(Fall 1997), and downward to elevations 

below 3,000 meters (9,483').

In the late Holocene, low-elevation montane 

forests mixed with steppe vegetation and 

low-elevation subalpine forests were defined 

by an increasingly open stand structure. 

Summer temperatures declined to preindustrial levels ca. 1850, and spruce and fir 

dominated subalpine forests and krummholz 

vegetation. Montane taxa retreated upslope, 

sagebrush steppe expanded at lower elevations, and alpine tundra dominated a larger 

range of high elevation areas, suggesting 

that drier conditions increased for parts of 

the southern Rockies (Markgraf and Scott 

1981; Fall 1997). Conditions similar to the 

present were established approximately two 

millennia ago, with modest treeline elevation 

fluctuations during the Medieval Climate 

Anomaly. 

2.3.5 Upper Columbia Basin

Paleoenvironmental data are available from 

the Upper Columbia Basin (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).

During glacial times, tundra-steppe communities dominated by Artemisia and Poaceae 

were widespread in the basins, reflecting 

cold dry conditions (figs. 10–11). As the climate warmed in the late Glacial period, pine, 

spruce, and fir 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, 

sagebrush, and Chenopodiaceae. Summer 

drought was more pronounced throughout 

much of the region during the early Holocene as the amplification of the seasonal 

insolation cycle resulted in warmer and 

Figure 9. Ecological 

response to changing climatic conditions following 

glacial retreat in the 

southern Rockies. 

Derived from Fall 

1997.

Figure 10. Ecological response to 

changing climatic 

conditions in the 

western portion of 

the Upper Columbia 

Basin following deglaciation. Derived 

from Carp Lake, 

Washington (Whitlock et al. 2000).

Elevation (m)

Southern U.S. Rocky Mountains

Tundra

Picea-Abies forest

Pinus-Pseudotsuga-Populus forest

Steppe 2500

3000

3250

3500

3750

4000

Tundra Picea-Abies forest

Tundra

Picea parkland Picea-Abies-Pinus parkland

2500

3000

3250

3500

3750

4000

Picea-Abies parkland

Pinus forest 2500

3000

3250

3500

3750

4000 Tundra

2500

3000

3250

3500

3750

4000

Krummholz

Late Holocene

ca. ~4 ka yrs BP to present

Early Holocene

ca. ~11-7 ka yrs BP

Late Glacial

ca. ~14-11 ka yrs BP

> 14,000 cal yr BP

Elevation (m)

Upper Columbia Basin

Pinus-Picea

Abies forest

Artemisia steppe

500

1000

1500

2000

Pinus-Artemisia

parkland

Pinus-Pseudotsuga

forest

Tundra-steppe

 with Artemisia

Poaceae

Picea

parkland

Pinus-Picea

Abies parkland

Pinus-Pseudotsuga

-Abies forest

Pinus-Artemisia

parkland

snow/ice

Pinus-Artemisia-Poaceae

parkland-steppe

Artemisia-Poaceae tundra-steppe 

Pinus parkland

alpine

500

1000

1500

2000

500

1000

1500

2000

500

1000

1500

2000

alpine alpine

Late Holocene

ca. ~4 ka yrs BP to present

Early Holocene

ca. ~11-7 ka yrs BP

Late Glacial

ca. ~14-11 ka yrs BP

> 14,000 cal yr BP

The last 20,000 years of environmental change
16

Climate and Ecosystem Change in the U.S. Rocky Mountains and Upper Columbia Basin A Framework for Climate Analysis and Reporting for Greater Yellowstone and Rocky Mountain Networks

effectively drier conditions at low and midelevations (Whitlock et al. 2000). Cool dry 

conditions in the Okanogan Highlands of 

northern Washington continued to support 

grasses and sagebrush, and the forest-steppe 

ecotone was north of its present location by 

as much as 100 kilometers (Mack et al. 1978). 

Pine, spruce, and fir were present in areas 

with greater precipitation (e.g., Waits Lake, 

eastern Washington) (Whitlock and Brunelle 

2007, fig. 11). In the mid-Holocene, increased effective moisture allowed the establishment and expansion of pine woodland at 

middle and higher elevations, and the upper 

treeline was higher than at present. In the 

western Columbia Basin, the expansion of 

pine woodland (primarily Pinus ponderosa) 

was followed by an invasion of mixed forest 

in the late Holocene (Douglas fir, larch, fir, 

western hemlock, and oak) (Whitlock and 

Brunelle 2007). 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 and fir pollen increased 

(Whitlock and Brunelle 2007) 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 period 

(Whitlock and Brunelle 2007).

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; 

Picea, spruce; Pinus 

contorta, lodgepole pine; Poaceae, 

grass; Pseudotsuga, Douglas 

fir; Shepherdia, 

buffaloberry; 

Tsuga heterophylla, 

western hemlock. 

Source: Whitlock 

and Brunelle 2007; 

reprinted with permission.
17

The western United States has experienced 

large-scale changes in climate, vegetation, 

and disturbance regime since the last glaciation 20,000 years ago. With the initial recession of glacial ice more than 14,000 years 

ago, the climate was colder and generally 

drier than at present and most regions were 

colonized by tundra communities. As the 

climate warmed and precipitation increased 

from 14,000 to 11,000 years ago, these tundra 

communities were replaced by subalpine 

parkland and then closed subalpine forest. 

During the early Holocene (11,000–7,000 

cal yr BP), the development of warmer and 

drier than present conditions led to more 

xerophytic vegetation and more fires in 

most areas. After 7000 cal yr BP, the climate 

became cooler and effectively wetter. In most 

regions, the modern vegetation and climate 

were established during the late Holocene 

(the last 4000 cal yrs). These changes highlight several important lessons for understanding the impacts of climate change on 

vegetation:

● Millennial-scale variations in climate 

over the last 20,000 years were caused by 

changes in the latitudinal and seasonal 

distribution of incoming solar radiation, 

the size and extent of the continental ice 

sheets, and attendant shifts in atmospheric 

circulation (e.g., southward displacement 

of the jet stream, the strength of the northeastern Pacific subtropical high-pressure 

system, and the intensity of monsoonal 

circulation). These slowly varying changes 

determined the distribution and composition of plant communities.

● Abrupt climate variation during the last 

20,000 years led to rapid changes in the 

assemblages and distribution of vegetation 

across the U.S. Rocky Mountains and the 

Upper Columbia Basin and influenced 

ecosystem processes such as fire.

 The Medieval Climate Anomaly 

(MCA, ca. 950–1250) and Little 

Ice Age (ca. 1400–1700) resulted in shifts in plant distributions and disturbance regimes 

in some locations, but were not 

uniformly manifested across 

the study area. For example, 

warm dry conditions in Yellowstone during the MCA led to 

increased fires in the summerdry areas while the summerwet areas of Yellowstone were 

effectively wetter (Whitlock and 

Bartlein 1993; Whitlock and et 

al. 2003).

 The ecological impact of these 

shorter climatic events (lasting centuries) is variable either 

because the climate signal was 

regionally heterogeneous or 

because plant communities 

were not responsive to climate 

change on this relatively short 

time scale.

● The western United States is influenced by 

two precipitation regimes—a summer-dry 

area under the influence of the northeastern Pacific subtropical high-pressure system and a summer-wet region strongly influenced by summer monsoon circulation. 

In the northern U.S. Rocky Mountains, 

the location of these regimes is sharply 

delimited and constrained by topography. 

These two regimes were enhanced during 

the early Holocene, when summer solar 

radiation was higher than at present. As a 

result, summer-wet areas became wetter 

and experienced fewer fires than at present, and summer-dry areas became drier 

with more fires and xerophytic vegetation 

than at present (Whitlock and Bartlein 

1993; Huerta et al. 2009; Millspaugh et al. 

2004). The contrast between summer-wet 

and summer-dry regions was greatest in 

the early Holocene. This past response 

suggests that future changes in precipitation regimes will also likely be spatially 

heterogeneous, and that the boundary 

between precipitation regimes will likely 

be quite sharp in mountainous regions.

● Although climate exerts strong controls 

on the distribution of vegetation at large 

spatial and long temporal scales, edaphic 

factors can amplify or minimize the response at smaller scales, as illustrated by 

the persistence of lodgepole pine forests 

Chapter 3: What do paleoenvironmental records tell 

us about millennial scale climate variations?
18

Climate and Ecosystem Change in the U.S. Rocky Mountains and Upper Columbia Basin A Framework for Climate Analysis and Reporting for Greater Yellowstone and Rocky Mountain Networks

on rhyolitic soils in Yellowstone’s Central 

Plateau throughout the Holocene (Whitlock et al. 1993; Millspaugh et al. 2000) 

and the influence of ultramafic soils on 

Holocene vegetation and fire regimes 

(Briles et al. in review). 

● Climate change can influence the distribution of vegetation via direct climate 

constraints (e.g., temperature and precipitation) or indirectly by influencing 

key ecosystem processes such as fire and 

nutrient cycling. Feedbacks related to 

vegetation changes can also influence fire 

by changing fuel availability.

The superposition of climate changes occurring on multiple time scales means that 

no period in the last 20,000 years is an exact 

analogue for the future. Nonetheless, the 

paleoenvironmental records show the resilience of vegetation to periods of extreme 

drought, changes in disturbance regimes, 

and rapid climate change. These examples 

provide insights about the sensitivity and 

pathways by which ecosystems respond to 

climate changes of different duration and 

intensity.

Key findings:

1. Natural variations in climate and the 

accompanying ecological responses 

occur at multiple temporal and 

spatial scales, all of which must be 

understood to explain modern plant 

communities and their distributions. Paleoecological data provide 

evidence of a range of responses 

that are not adequately represented 

in the last two centuries. A baseline 

of natural variability for restoration 

efforts must therefore consider a 

longer time scale.

2. Many terrestrial ecosystems in 

the study region were established 

during the last 3000 to 4000 years 

in response to gradual cooling and 

increased effective moisture in the 

late Holocene. More subtle changes 

occurred during the Medieval 

Climate Anomaly and Little Ice Age 

(as discussed in the next section). 

The sequence of events that led to 

present vegetation is different from 

that projected for the future, which 

argues against strategies to restore 

to a reference condition. Instead, we 

need process-based approaches and 

flexible management responses. 

3. Key ecosystem processes such as fire 

are driven by climate at large spatiotemporal scales. Patterns of the fire 

in the 20th century poorly represent 

the potential range of fire regimes 

that have occurred in the past and 

may occur in the future.
19

4.1 Primary drivers of change

The primary drivers of climate during the 

last 2000 years include ocean-atmosphere 

interactions, volcanic eruptions, changes 

in incoming solar radiation, and increases 

in atmospheric greenhouse gases (GHGs) 

and aerosols due to human activities (fig. 2). 

Climate model simulations indicate that during the pre-industrial portion of the last 2000 

years, solar fluctuations and volcanic eruptions were likely the most strongly varying 

forcings, and in combination with ocean-atmosphere interactions they likely resulted in 

periods of relative warmth and cold (Amman 

et al. 2007; Mann and Jones 2003; Mann 

2007; Jones et al. 2009; Esper et al. 2002). 

GHGs and tropospheric aerosols varied little 

until around AD 1700 when they began to 

be significantly impacted by human activities (Solomon et al. 2007; Keeling 1976).The 

rapid rise in 20th century global temperatures 

is best explained by the combination of natural and anthropogenic GHG forcings, with 

GHGs playing an increasingly dominant role 

during recent decades.

4.2 Biophysical conditions 

Compared to the conditions driving continental deglaciation and the Pleistocene/

Holocene transition, orbital and radiative 

forcings over the last two millennia have 

remained relatively constant and can be considered more analogous to modern conditions. However, the major centennial-scale 

climate variations evident during this period 

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). Because the rates of climatic 

change during the last two millennia were 

much smaller in magnitude than those 

associated with the late Glacial and early 

Holocene, the ecological response to climatic variation 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 species found in modern vegetation assemblages but rarely to widespread 

changes in dominant vegetation. Vegetation 

types since the late Holocene are considered 

similar to present conditions and, overall, 

the magnitude and duration of the changes 

are not comparable to those of the Pleistocene–Holocene transition. Rather, the last 

2000 years have had smaller magnitude and 

shorter duration (centennial, decadal, interannual) climatic controls on ecosystems that 

have nevertheless resulted 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, 

the North Atlantic Oscillation, the Atlantic 

Multidecadal Oscillation, and the El Niño–

Southern Oscillation interact to influence 

temperature, precipitation, and atmospheric 

circulation and help explain droughts and 

wet periods at interannual to interdecadal 

scales (e.g., Cook et al. 2004; Mann et al. 

2009; Gray et al. 2004; Hidalgo 2004; Fye et 

al. 2003; McCabe et al. 2004; Graumlich et 

al. 2003; Enfield et al. 2001). Past records of 

temperature illustrate centennial, decadal, 

and interannual variation, providing a context for understanding ecosystem changes 

that occurred in different regions (fig. 12).

Tree-ring records from throughout the 

western United States show natural variation 

in temperature, precipitation, and available 

moisture during the last two millennia, some 

of which is synchronous across large areas of 

the four climate regions in this study, while 

other variations are more representative of 

local phenomena (figs. 12–13, Pederson et al. 

2006; Cook et al. 2004). These records show 

that decadal and multidecadal fluctuations in 

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).

Chapter 4: The last 2000 years of environmental change

0 500 1000 1500 2000

Temperature Anomaly

(Degree C)

-3

-2

-1

0

1

2

Luckman & Wilson 2005 Northern Rockies

Salzer & Kipfmueller 2005 Central Rockies

Mann et al. 2008 Global

Mann et al. 2008 Northen Hemisphere

Regional to Global Temperature Reconstructions
20

Climate and Ecosystem Change in the U.S. Rocky Mountains and Upper Columbia Basin A Framework for Climate Analysis and Reporting for Greater Yellowstone and Rocky Mountain Networks

precipitation are a defining characteristic of 

the climate during 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).

Two major, well-documented examples of 

centennial-scale climate change during the 

last two millennia are the Medieval Climate 

Anomaly (950–1250) and the Little Ice Age 

(1400–1700). As indicated by figure 13 (Cook 

et al. 2007), a number of anomalous warm 

dry periods and cool wet periods occurred 

(see Biondi et al. 1999; Meko et al. 2007; 

Salzer and Kipfmuller 2005; Cook et al. 

2007), and resulted in extensive hydrological 

and ecological impacts. During the MCA, 

dry precipitation 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 variable. Elevated aridity and “mega-droughts” (lasting 

at least ~50 years) were common across the 

western United States, with more areas experiencing drought simultaneously than during 

the LIA or most of the 20th century (fig. 13, 

Cook et al. 2007).

Data from a number of sites suggest that 

regionally synchronous drought events 

occurred regularly during the MCA, with 

durations and extents unmatched in the late 

Holocene (Cook et al. 2007, 2004). Glacial 

retreat occurred in mountainous areas of 

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 central 

U.S. Rockies, Greater Yellowstone Area, and 

parts of the southern Rockies experienced 

increased summer moisture (Whitlock and 

Bartlein 1993). Upper treelines in some areas 

increased 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 lowfrequency variation in ocean-atmosphere 

interactions was an important factor (Mann 

et. al 2009).

The LIA was a period of anomalous Northern Hemisphere cooling when mountain 

glaciers throughout the western United 

States expanded, many reaching their Holocene maximum (e.g., Pederson et al. 2004; 

Luckman 2000; Watson and Luckman 2006). 

While temperatures across the study area 

were persistently cooler than the long-term 

average (>1°C [2°F] cooler), some data suggest 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 

Figure 13. Long-term aridity changes in the West as measured by 

the percent area affected by drought (PDSIb_1, thick black line), 

95% boostrapped confidence intervals (light-blue dotted lines) 

and the long-term mean (thin horizontal black line). The four most 

significant (p<0.05) dry and wet epochs since 800 are indicated by 

arrows. The 1900–2003 interval is highlighted by the yellow box. 

The average drought area during that period and for the 900–1300 

interval is 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; reprinted with permission.
21

Contents

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 glacier advance during the LIA (Watson and Luckman 

2004; Pederson et al. 2004). Advances of this 

magnitude did not occur in the southern and 

central Rockies, demonstrating the spatial 

variability in LIA climate 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 which led to decreased incoming solar 

radiation from ca. 1645 to 1715 is considered 

one of the main variables explaining cooler 

conditions for this interval of the LIA (Eddy 

1976; Luckman and Wilson 2005).

The last 2000 years of environmental change
Chapter 6: 
23

The following case studies from the four climate regions highlight examples of biophysical and biotic response to climate change 

during the last two millennia and provide 

clues to the timing and extent of future 

biotic changes.

5.1 Northern U.S. Rocky Mountains

5.1.1 Drought variability and 

ecosystem dynamics in Glacier National 

Park

Records of past hydroclimatic changes, 

glacier dynamics, and fire activity in Glacier 

and Waterton national parks show how 

interdecadal and longer-term changes in 

climate interact to alter ecosystem processes 

in the northern Rockies (fig. 14). Long-term 

changes in regional temperature (e.g., the 

relatively cool conditions of the LIA in contrast to warmer temperatures in the 1350s to 

1450s and especially over the last half of the 

20th century, fig. 12) combined with persistent shifts in summer and winter moisture 

regimes over decadal to multidecadal time 

scales have a particularly strong impact on 

fire regimes in the Glacier National Park region (fig. 14a–c). Regional records spanning 

the last three centuries show that periods 

from the 1780s to the 1840s and the 1940s to 

the 1980s had generally cool wet summers 

coupled with high winter snowpack, resulting in extended (>20 years) burn regimes 

characterized by small infrequent fires with 

relatively little area burned. Conversely, 

decadal and longer combinations of low 

snowpack and warm dry summers resulted 

in burn regimes characterized by frequent 

fires and large total area burned (e.g., 1910–

1940, 1980–present).

Figure 14. Relationship between Glacier National Park summer drought, inferred winter 

snowpack, fire area burned, and glacial recession 1700 to present. (a) Instrumental and reconstructed summer drought (MSD) normalized by converting to units of standard deviation 

and smoothed using a 5-year running mean. (b) Measured spring snowpack (May 1 SWE) 

anomalies (1922–present) and average annual instrumental and reconstructed PDO anomalies. Each time series was normalized and smoothed using a 5-year running mean. For ease 

of comparison, the instrumental and reconstructed PDO index was inverted due to the strong 

negative relationship between PDO anomalies and May 1 snowpack. (c) Fire area burned 

timeline for the GNP region. Timeline is presented with maps of fire activity during periods of 

interesting winter and summer precipitation regimes. (d) Maps showing the decrease in area 

of the Sperry Glacier at critical points from 1850 to 2003. The retreat patterns of the Sperry 

Glacier are representative of regional patterns of recession for glaciers sensitive to regional 

climate variability. Source: Pederson et al. 2006, American Meteorological Society; reprinted 

with permission.

Chapter 5: Case studies 
24

Climate and Ecosystem Change in the U.S. Rocky Mountains and Upper Columbia Basin A Framework for Climate Analysis and Reporting for Greater Yellowstone and Rocky Mountain Networks

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 

(ca. 1850), four centuries of generally cool 

summers prevailed (fig. 12). What drove the 

glaciers to reach their greatest extent since 

the Last Glacial Maximum was cool wet 

summers coupled with generally high snowpack conditions from 1770 to 1840 (Watson 

and Luckman 2004; Pederson et al. 2004). 

During subsequent periods when summer 

drought and snowpack were generally in opposing phases (e.g., 1850–1910) and summer 

temperatures remained relatively cool (fig. 

14), glaciers experienced moderate retreat 

rates (1–7 m/yr [3–23']). From 1917 to 1941, 

however, sustained low snowpack, extreme 

summer drought conditions, and high summer temperatures drove rapid glacial retreat. 

The Sperry Glacier retreated at 15–22 meters 

(49–72') per year and lost approximately 

68% of its area (fig. 14d). Other glaciers such 

as the Jackson and Agassiz glaciers at times 

retreated at rates ≥100 meters (328') per year 

(Carrara and McGimsey 1981; Key et al. 

2002; Pederson et al. 2004). Climatic conditions from 1945 to 1977 became generally 

favorable (i.e., level to cooling summer temperatures, high snowpack with variable summer drought conditions) for stabilization and 

even accumulation of mass at some glaciers 

(fig. 14a,b). Since the late 1970s, however, exceptionally high summer temperatures (Pederson et al. 2010) combined with low winter 

snowpack and multiple periods of severe and 

sustained summer drought have resulted in a 

continuation of rapid glacier retreat. 

As demonstrated by these records from 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 time scales. This relationship between biophysical and ecosystem 

processes with hydroclimatic changes that 

can appear to be static on shorter (e.g., decadal) scales poses challenges to management 

and sustainability in three ways. First, longduration proxy reconstructions call into 

question the conventional strategy of defining reference conditions or management 

targets based solely on < 100-year records. 

For example, the use of 30-year climatology 

for the allocation of natural resources and 

development of resource management goals 

is flawed because a 30-year climatic mean 

may only capture a single mode of climate 

variability (e.g., an extended regime of wet 

or dry conditions). Second, abrupt, highmagnitude changes from one climate regime to the next can onset rapidly and have 

prolonged impacts on ecosystem processes. 

These persistent and frequent climate-related shifts may amplify or dampen the effects 

of management activities. Lastly, decadal 

and longer persistence of either deficits or 

abundances in 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 how ecosystems 

respond to climate change at longer temporal 

scales provides important context for future 

management.

5.2 Central U.S. Rocky Mountains 

and the Greater Yellowstone Area

5.2.1 Changing distributions of Utah 

juniper

A long-term reconstruction of changing distributions of Utah juniper (Juniperus osteosperma) in the CR-GYA provides an example 

of how climate variation, complex topography, and spatial distribution of suitable habitat and biotic factors interact to govern plant 

invasions. It also provides evidence contrary 

to the popular idea that plant invasions are 

characterized by a steady and continuous 

march across landscapes. The distribution of 

Utah juniper in the mountains of Wyoming, 

southern Montana, and Utah during the late 

Holocene has been tracked by radiocarbondating fossilized woodrat middens (Lyford et 

al. 2003). During a dry period in the midHolocene (ca. 7500–5400 BP), Utah juniper 

migrated north into the central Rockies via 

a series of long-distance dispersal events. 

Further range expansion and backfilling 

of suitable habitat was stalled during a wet 

period from 5400 to 2800 BP (Lyford et al. 

2003). In response to warmer, drier conditions that developed after 2800 years BP, 
25

Utah juniper populations rapidly expanded 

within the Bighorn Basin, especially 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 

the distributions of species with centennialscale life expectancies. In the case of the 

Utah juniper, Lyford et al. (2003) noted that 

establishment rates are significantly more 

affected by adverse climatic conditions than 

by individual or population survival where 

the species is already established. This could 

explain, in part, the tendency for Utah juniper populations to remain static instead of 

contracting during the Holocene.

In general, Utah juniper range expanded 

during periods characterized by warmer, 

drier conditions, and expansion and establishment ceased during cool, wet periods 

(Lyford et al. 2003). Thus, the migration of 

Utah juniper into the central Rockies was 

at least partly controlled by millennial-scale 

climatic variations during the Holocene. Although Utah juniper distribution is severely 

limited by cool temperatures and high precipitation in higher elevations of the central 

Rockies, the species inhabits only a fraction 

of the suitable climate space in the region 

(fig. 15) because it is limited by suitable substrate; present distributions cover more than 

90% of the substrate in the region deemed 

highly suitable for Utah juniper survival.

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. It also highlights the importance of recognizing other 

environmental factors that affect species distribution. While suitable climate can allow 

a species to become established, spatially 

variable factors such as substrate, dispersal, 

and competition influence how successfully 

it can disperse to and colonize areas with 

suitable climate. 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. 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 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 species dispersal will interact with the 

spatial distribution of suitable habitat and 

climate variability to govern future invasions.

5.2.2 Precipitation variability in 

Wyoming’s Green River Basin

Tree-ring records from Wyoming’s Green 

River basin provide a reconstruction of 

Figure 15. Habitats for Utah juniper in Wyoming and adjacent Montana. Black areas indicate extremely suitable habitats and gray areas 

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 (Knight et al. 1987, 

Driese et al. 1997). Note that climate (a) over-predicts the modern 

distribution, which is strongly constrained by substrate variables (b), 

and that favorable habitat is patchily distributed (c). Source: Lyford et 

al. 2003; reprinted with permission.

Case studies
26

Climate and Ecosystem Change in the U.S. Rocky Mountains and Upper Columbia Basin A Framework for Climate Analysis and Reporting for Greater Yellowstone and Rocky Mountain Networks

drought conditions for the last millennium 

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 1100-year record of the Palmer Drought 

Severity Index, a measure of drought that 

includes precipitation and temperature 

trends for southwestern Wyoming (Cook et 

al. 2004). This record is typical of many areas 

of the CR-GYA, showing above-average 

effective moisture in the early 20th century 

(Woodhouse et al. 2006; Gray et al. 2004, 

2007; Meko et al. 2007) and the potential 

for severe, sustained droughts far outside 

the range of 20th century records, including 

several multidecadal droughts prior to 1300 

(fig. 16).

Although using the last century as a reference for climate conditions would suggest 

that the Green River Basin is wet and relatively free of drought, this longer-term reconstruction indicates that some of the most 

severe droughts in the 1930s and 1950s were 

relatively minor compared to many dry periods in past millennia, and that the second 

half of the 20th century was relatively wet 

with no prolonged droughts (fig. 16). This 

study provides strong evidence that drought 

periods are a natural feature of the regional 

climate and that long-term records are critical for understanding natural variability of 

climate conditions in the western United 

States.

5.3 Southern U.S. Rocky Mountains

5.3.1 Changing pinyon pine 

distribution

Gray et al. (2006), who used woodrat midden and tree-ring data to track the spatial 

and temporal patterns of pinyon pine (Pinus 

edulis Engelm.) distribution in the Dutch 

John Mountains (DJM) of northeastern 

Utah, showed that the distribution during 

the Holocene has been strongly controlled 

by multidecadal precipitation patterns. The 

DJM population is an isolated northern 

outpost of pinyon pine that established ca. 

1246. Similar to Utah juniper in the central 

Rockies, the pinyon pine probably reached 

the DJM via long-distance dispersal from 

the Colorado Plateau (Jackson et al. 2005) 

during the transition from the warmer, drier 

MCA to the cooler, wetter LIA (fig. 17). DJM 

pinyon pine expansion stalled in the late 

1200s and significant recruitment did not 

resume until the 14th century pluvial, when 

regionally mesic conditions promoted establishment (fig. 17).

The case of DJM pinyon pine demonstrates 

the importance of episodic, multidecadal 

climatic variation in controlling rates of ecological change in the southern Rockies during past millennia. Records suggest that the 

development of the DJM population was not 

a steady movement associated with improving climate conditions but rather a markedly 

episodic invasion regulated by fluctuations 

in precipitation (Gray et al. 2006). In particular, this example highlights the consequences 

of having short-term, episodic climatic variation superimposed on centennial to millennial scale climate change, a pattern that can 

significantly affect species migration and 

establishment in ways that are more complex 

than a simple wave-like expansion.

As with previous studies of plant invasions 

at longer time scales (Lyford et al. 2003), this 

study suggests that climatic variation can amplify or dampen the probability of survival 

and reproduction after a species colonizes 

new areas. Climate can also modify the 

Figure 16. 1100 years of drought history in the Green River Basin 

region of southwest Wyoming (reconstructed from tree rings, Cook 

et al. 2004). The plot shows values for the Palmer Drought Severity 

Index (see also fig. 13). Positive values (blue) of the index indicate wet 

conditions and negative values (red) drought. Values are plotted so 

that each point represents mean conditions over a 25-year period. 

Source: Gray and Andersen 2010.
27

density and distribution of favorable habitats 

across the landscape and influence competitive interactions and disturbance processes 

(Gray et al. 2006). For example, different 

locations within a region may experience 

similar changes in average precipitation or 

temperature during a particular period, yet 

differences in the variability of precipitation 

could easily produce different disturbance 

dynamics and different end states. It is often 

assumed that vegetation responds to climate 

change with a steady wave-like movement 

to better growing conditions, yet the DJM 

example reveals that species are influenced 

by other factors, including the dynamics of 

long-distance dispersal and climate variability 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 

spatial and temporal scales.

5.4 Upper Columbia Basin

5.4.1 Climate variation and firerelated sedimentation

Fire-related sediment deposits in central 

Idaho reveal millennial to centennial scale 

climate variation and its control on fire 

regimes. Pierce et al. (2004) dated charcoal 

in alluvial fan deposits to reconstruct firerelated erosion events 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 late Holocene, 

especially during the LIA, and suggested 

that these were associated with frequent 

fires of low to moderate severity. Large 

Figure 17. Establishment and presence/absence of pinyon pine in relation to climate variability and elevation during the Holocene. (a) Percentages of pinyon-type pollen (black vertical 

bars) and presence (solid circles) or absence (open circles) of pinyon pine macrofossils from 

12,000 years of woodrat midden records collected in the DJM. (b) Map of the study area 

showing locations of midden sites (open circles) and sampling units used in the tree-ring 

age studies (shaded polygons). Each midden site is a cave or rock overhang where one or 

more of the 60 middens were collected. The estimated establishment dates are based on the 

average age of the four oldest pinyons found in each sampling unit. (c) Ages for the oldest 

pinyon on DJM and the four oldest pinyons in each of the eight sampling units (black dots) 

plotted against reconstructed annual (gray line) and 30-year smoothed (black line) precipitation values for the Uinta Basin Region. (d) Percentage of the western United States experiencing drought conditions during the last 1200 years as reconstructed from a large tree-ring 

network (Cook et al. 2004). Data are plotted as a 50-year 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; reprinted with permission.

Case studies
28

Climate and Ecosystem Change in the U.S. Rocky Mountains and Upper Columbia Basin A Framework for Climate Analysis and Reporting for Greater Yellowstone and Rocky Mountain Networks

sedimentation events were associated with 

prolonged periods of drought and severe 

fires that were most pronounced during the 

MCA (Pierce et al. 2004, fig. 18).

These results were 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, fig. 18). Changes in 

inferred fire occurrence were synchronous 

between central Idaho and northern Yellowstone during the warmer, drier MCA. 

In central Idaho, longer intervals of warm 

dry conditions allowed for drying of large 

fuels to increase the frequency of large 

stand-replacing canopy fires and fire-related 

erosion events. Hence, strong climatic 

controls changed a fuel-limited, infrequent, 

low-severity fire regime to a fuel-rich, highseverity, stand-replacing fire regime. Large 

stand-replacing fires also increased in mixed 

conifer forests of northern Yellowstone during the MCA (fig. 18), coinciding with large 

pulses of fire-related sedimentation ca. AD 

1150. Meyer et al. (1995) inferred that this 

activity resulted from increased intensity and 

interannual variability in summer precipitation. During the LIA, ca. AD 1400-1700, 

cooler wetter conditions in both northern 

Yellowstone and central Idaho are inferred 

to have maintained high-canopy moisture 

that inhibited canopy fires and facilitated the 

growth of understory grass and fuels required to sustain frequent low-severity fires 

in central Idaho (Pierce et al. 2004).

Paleoenvironmental investigations like these 

provide further evidence that millennial and 

centennial scale variations in temperature 

and precipitation have influenced biophysical conditions and important disturbance 

processes across the Upper Columbia Basin 

and CR-GYA. The results of these and similar studies (e.g., Whitlock et al. 2003) suggest 

that natural climatic variability acts as a primary control of ecosystem processes, which 

has important implications for management. 

Efforts to manage fuels in different 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 20th century 

structural and disturbance characteristics 

may have limited impacts under future climatic conditions. 

Figure 18. Fire-related sedimentation in the South Fork Payette River (SFP) area in central 

Idaho and in Yellowstone National Park (YNP), Wyoming. Probability distributions were 

smoothed using a 100-year running mean to reduce the influence of short-term variations 

in atmospheric 14C but retain peaks representing the most probable age ranges. The trend 

of decreasing probability probability before 4000 cal yr BP reflects fewer sites and decreased 

preservation of older deposits. (a) SFP small events are thin deposits probably related to low 

or moderate burns. Note correspondence of peaks with minima in YNP fire-related sedimentation and major peak during the LIA. Fewer near-surface deposits since 400 cal yr BP 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 peak in large-event probability during the MCA. Source: Pierce 

et al. 2004; reprinted with permission.
29

Paleoenvironmental records of the last 2000 

years provide information on decadal and 

centennial scale climate variability, the associated ecological responses, and the context 

for anticipating future change. These records 

suggest that ecosystems have responded to 

climatic variation primarily through shifts in 

ranges, ecotonal position, and community 

composition and structure. High-resolution 

records indicate that the rate, magnitude, 

and duration of climate change strongly 

govern the ecological response. Species 

respond to changing environmental conditions by moving up or downslope and by 

increasing or decreasing in density (e.g., 

treeline “infill”) and abundance. In contrast 

to the widely held assumption that changes 

in plant distribution are characterized by 

steady advances across landscapes, paleoenvironmental evidence suggests that such 

range adjustments are episodic in response 

to climatic conditions, occurring rapidly 

when conditions are suitable and slowly or 

not at all otherwise (Gray et al. 2006: Lyford 

et al. 2003). Predicting plant response to 

future climate change will require consideration of the rate and magnitude of climate 

change, spatial heterogeneity in biophysical 

conditions, the catalytic and synergistic role 

of ecosystem responses (e.g., fire, nutrient cycling, insect outbreaks), and intrinsic 

biotic limitations (Whitlock and Brunelle 

2007; Jackson et al. 2009a).

Paleoenvironmental records from throughout the western United States demonstrate 

that ecosystem processes are strongly 

influenced by climate variability occurring 

at decadal and longer timescales whereas 

management planning is often based on 

climate “regimes” that are determined using 

a few decades of climate data. For example, 

fuel management goals are often developed 

from fire patterns over the last century even 

though paleoenvironmental data show that 

this only captures 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). Records of change in the U.S. Rocky Mountains 

and Upper Columbia Basin provide context 

for better understanding current changes 

and the character of ecosystem response 

we might expect in the future. In particular, 

these records highlight:

● Decadal to centennial scale climate variability are related to ocean-atmosphere 

interactions such as ENSO, volcanic eruptions, and solar variability. These events 

include the Medieval Climate Anomaly 

(ca. 950–1250), the Little Ice Age (ca. 

1400–1700), and numerous decadal and 

multidecadal droughts and pluvials.

● Prolonged droughts in the last millennia, 

which rival 20th century droughts in duration and magnitude, influenced vegetation 

in many ways, including:

 Changes in treeline position.

 Increases and decreases in 

treeline density (e.g., the “infill” 

phenomenon experienced in 

some areas during the MCA).

 Multidecadal to centennial 

scale climate variability that 

results in episodic species 

dispersal, colonization, and 

establishment.

● At small temporal (decadal to centennial) 

and spatial (m2 to km2) scales, local factors 

(e.g., substrate, disturbance, and human 

activities) interact with climatic variation 

to influence the distribution of vegetation.

● Human activity was superimposed on 

climatic variability and altered prehistoric 

fire regimes in some areas of the western United States, but the impacts were 

mostly local (Whitlock and Knox 2002). 

As evidenced by changes in fire regimes 

of the late 19th and early 20th centuries, 

human activities left a strong imprint on 

the landscape and can rapidly influence 

ecosystems and ecosystem processes such 

as fire.

Key findings:

1. Using the last century as a baseline 

for climate conditions does not 

capture important scales of natural 

Chapter 6: What can we learn from the last 2000 

years about decadal and centennial scale climate change? 
30

Climate and Ecosystem Change in the U.S. Rocky Mountains and Upper Columbia Basin A Framework for Climate Analysis and Reporting for Greater Yellowstone and Rocky Mountain Networks

climate variability and is often an 

inadequate reference for considering 

future climate change.

2. Even large-scale climatic changes 

can have spatially variable impacts 

due to interactions among a number 

of controlling factors. For example, 

cooling during the Little Ice Age, 

evident at the scale of the Northern 

Hemisphere occurred asynchronously or may not have registered at 

some locations in the study area (fig. 

14).

3. Rapid climate changes and associated ecosystem responses have 

occurred in the past and will likely 

occur in the future.
31

Throughout much of the western United 

States, 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, the footprint of 

broad-scale climate changes will vary across 

finer spatial and temporal scales, and from 

one area to another. As a direct measurement 

of climate conditions for the last century, 

instrumental observations provide the most 

accurate and reliable data available, but they 

are influenced by local biases (e.g., differing 

slopes or aspects) 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., 

with comparable elevation and vegetation 

cover). 

7.1.1 Temperature

For all climate regions in this study, 20th

century climate change is characterized 

by high spatial and temporal variability. At 

broad spatial and temporal 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 United States from 0.5°C to 2°C 

(Pederson et al. 2010; Mote 2003; Ray et al. 

2008), although cooling has occurred in a 

few areas, e.g., southeastern Colorado (Ray 

et al. 2008) and individual sites in the Northwest (CIG 2010). Where temperatures are 

trending higher, the rate of change varies 

by location and elevation, but is typically 

1°C since the early 20th century (Hamlet et 

al. 2007). In most of the northern portions 

of the study area, temperatures generally 

increased from 1900 to 1940, declined from 

1940 to 1975, and have increased since then 

(Parson 2001). Similarly, in the southern 

Rockies, temperatures generally increased in 

the 1930s and 1950s, cooled in the 1960s and 

1970s, and have consistently increased since 

then (Ray et al. 2008). The rate of increase 

for much of the study area doubled since the 

mid-20th century, with most of this warming 

occurring since 1975 (fig. 19).

Temperature increases are more pronounced 

during the cool season (Hamlet and Lettenmaier 2007). In the northern U.S. Rocky 

Mountains, 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 

West (Diaz and Eischeid 2007). In addition, 

nighttime minimum temperatures are increasing faster than are daytime maximums, 

resulting in a decreased diurnal temperature 

range (Pederson et al. submitted). This has 

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 were 0.87°C higher in 1987 

to 2003 than in 1970 to 1986, and were the 

warmest since 1895 (Westerling et al. 2006). 

Bonfils et al. (2008) and Barnett et al. (2008) 

Figure 19. Observed annual average North American 

surface temperature, 1950–2007 

(left) and the time 

series of the annual 

values of surface 

temperature averaged over North 

America (right). 

Annual anomalies 

are with respect to a 

1971–2000 reference. The smoothed 

curve (black line) indicates low frequency variations. (Data 

source: UK Hadley 

Center’s CRUv3 

global monthly gridded temperatures.) 

Source: Ray et al. 

2008; reprinted with 

permission.

Chapter 7: 20th century climate change 

and the instrumental record 
32

Climate and Ecosystem Change in the U.S. Rocky Mountains and Upper Columbia Basin A Framework for Climate Analysis and Reporting for Greater Yellowstone and Rocky Mountain Networks

found that the recent warming observed in 

mountainous areas across the West cannot 

be entirely explained by natural forcings 

(e.g., solar, volcanic, and ocean-atmosphere 

interactions); a major portion is attributable 

to human-influenced changes in greenhouse 

gas and aerosol concentrations.

7.1.2 Precipitation

Trends in precipitation for the study area 

are far less clear. Instrumental data from 

the last century show modest increases in 

precipitation for much of the northwestern United States (fig. 20; Mote et al. 1999, 

2003, 2005), but no trends in parts of the 

southern Rockies (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 last century impacted large 

areas. Though 20th century droughts had 

substantial socioeconomic and ecosystem 

impacts, there is ample evidence they were 

not as severe in duration and magnitude as 

a number of drought events that occurred 

during the last millennium (Cook et al. 2007, 

2004; Meko et al. 2007). For example, two 

droughts in the 1930s and 1950s impacted 

much of the study area. The 1930s drought 

was more widespread and pronounced in the 

northern and central climate regions while 

the 1950s drought was centered more on 

the south-central and southwestern United 

States (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 are likely linked 

to low-frequency oscillations in ocean-atmosphere 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).

7.1.3 Surface hydrology

Generally speaking, snowmelt and peak 

runoff has tended to occur earlier since 

1950, and river flows in many locations are 

decreasing during late summer (Pederson et 

al. 2010, 2009; Mote 2006; Barnett et al. 2008; 

Stewart et al. 2004; McCabe and Clark 2005). 

Recent impacts on snowpack and surface 

hydrology are strongly associated with more 

precipitation falling as rain than snow, and 

warming temperatures driving earlier snowmelt and snowmelt-driven runoff (fig. 21; 

Pederson et al. 2010, 2009; Mote 2006; Barnett et al. 2008; Stewart et al. 2004; McCabe 

and Clark 2005), leading to reduced surface 

storage of moisture and increasingly low 

baseflows during the dry summer months 

(Luce and Holden 2009). Examination of 

paleoclimate data and instrumental records 

(e.g., stream gages, snow course records, valley meteorological stations) suggests, however, that the total amount of cool season 

precipitation received across a particular 

region is more strongly associated with natural multidecadal, decadal, interannual, and 

intra-annual variability in ocean-atmosphere 

conditions (e.g., PDO, AMO, and ENSO).

Modest increases in precipitation have occurred in parts of the central and northern 

study area, but modest declines have taken 

place in parts of the southern Rockies, and 

significant declines in snowpack are evident 

throughout much of the Northwest. This decline is especially prevalent in the northern 

U.S. Rocky Mountains and parts of the Upper Columbia Basin (Hamlet and Lettenmaier 2007; Pederson et al. 2004, 2010; Selkowitz 

et al. 2002; Mote et al. 2005, 2008). Warmer 

temperatures and declining snowpack have, 

in turn, contributed to significant declines 

in the region’s glaciers. In Glacier National 

Figure 20. Trends 

in April 1 SWE 

from snow course 

observations, 

1960–2002. SWE is 

an important indicator for forecasting 

summer water 

supplies. Positive 

trends are shown 

in blue and negative in red by the 

scale indicated in 

the legend. Source: 

Mote 2006, American Meteorological 

Society; reprinted 

with permission.
33

Park, glaciers have decreased in area by 

more than 60% since 1900 (Hall and Fagre 

2003; Key et al. 2002), and only 26 of the 150 

glaciers and snow and ice fields present in 

1910 remain (Pederson et al. 2010). Changes 

in glacier mass and area are emblematic of 

changes in surface hydrology across the 

West, with the recent substantial declines 

related to increases in greenhouse gases and 

aerosols (Barnett et al. 2008; Bonfils et al. 

2008; Pierce et al. 2008; Hidalgo et al. 2009). 

During the last century, drought conditions 

have been increasing in the central and 

southern parts of the study area and decreasing in the northern parts (Andreadis and 

Lettenmaier 2006). Drought conditions are 

expected to increase for much of the study 

area over the coming decades (Hoerling and 

Eischeid 2007).

7.1.4 Ocean-atmosphere interactions

Ocean-atmosphere interactions are important drivers of interannual to multidecadal 

variability in temperature and precipitation 

(Pederson et al. 2010, 2009; Gray et al. 2007, 

2004, 2003; McCabe et al. 2004; Hidalgo 

2004; Cayan et al. 1998; Dettinger et al. 

2000), but their impacts vary greatly across 

latitudinal, elevational, and longitudinal 

gradients. The El Niño–Southern Oscillation 

and Pacific Decadal Oscillation measures 

of high and low frequency variability in sea 

surface temperatures (SSTs) are, respectively, 

the predominant sources of interannual and 

interdecadal 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) or La Niña (the 

cool phase). An El Niño event 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 1969; Cane and 

Zebiak 1985; Graham and White 1988; Horel 

and Wallace 1981; Philander 1990; Chang 

and Battisti 1998). A La Niña event 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 to 7 year cycles (Bjerknes 1969; Cane and 

Zebiak 1985; Graham and White 1988; Horel 

and Wallace 1981; Philander 1990; Chang 

and Battisti 1998). 

The PDO exhibits alternate cool and warm 

phases with a spatial pattern similar to that 

of ENSO, but these phases typically last for 

20 to 30 years (Mantua et al. 1997). Several 

switches occurred between warm and cool 

PDO phases during the 20th century and the 

magnitude of PDO phases increased in the 

latter half (McCabe et al. 2004; Mantua et 

al. 2002, 1997). The Atlantic Multidecadal 

Oscillation (AMO), representing low frequency (50–80 yr) oscillations in North Atlantic SSTs, has been linked to multidecadal 

variability in temperature and precipitation 

in the western United States through complex interactions with the PDO and ENSO, 

but the magnitude of the AMO influence 

is debated (Kerr 2000; McCabe et al. 2004; 

Enfield et al. 2001).

Changes in ENSO and PDO impact precipitation differently across the West, with 

winter precipitation in the Upper Columbia 

Basin and the northern Rockies being negatively correlated with warm conditions in 

the equatorial Pacific (i.e., during El Niños). 

Parts of the central and southern Rockies 

tend to be wetter than average during El Niños and dryer during La Niñas (Mote et al. 

2005). Likewise, the Northwest generally receives more winter precipitation during the 

cool-phase PDO, and the Southwest often 

receives more precipitation during the warm 

phase. Further evidence for spatial heterogeneity in ENSO and PDO impacts can be seen 

in Greater Yellowstone, where windward 

aspects and the high mountains and plateaus 

Figure 21. Trends 

in winter mean 

wet-day minimum 

temperatures, 

1949–2004. 

Source: Knowles et 

al. 2006, American 

Meteorological Society; reprinted with 

permission.

20th century climate change and the instrumental record
34

Climate and Ecosystem Change in the U.S. Rocky Mountains and Upper Columbia Basin A Framework for Climate Analysis and Reporting for Greater Yellowstone and Rocky Mountain Networks

to the west tend to follow more of an Upper 

Columbia Basin and northern Rockies pattern, while 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 the PDO, the AMO, 

and variations in ENSO. For example, the 

Dust Bowl drought, which was associated 

with a positive AMO and PDO, was centered 

primarily over the Southwest, whereas the 

1950s 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 West are associated with low-frequency variations in the 

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 

Rockies than in parts of the Southwest 

(Hidalgo 2004). In the Southwest and coastal 

Northwest, variations in precipitation and 

warm-season water availability appear more 

sensitive to low-frequency ENSO variations 

than to 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-atmosphere interactions such 

as ENSO and PDO are partially responsible 

for variations in climatic conditions across 

this study area, research suggests that since 

the late 20th century increased greenhouse 

gas and aerosol concentrations have been 

amplifying, dampening, and, in some cases, 

overriding the influence of these phenomena 

(Barnett et al. 2008; Bonfils et al. 2008; McCabe et al. 2008; Gray et al. 2006, 2003).

7.1.5 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 United States (Yin 2005; Lorenz 

and DeWeaver 2007), which could 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 20oN), which results in a poleward 

displacement of the Pacific Northwest storm 

track, increased west to east flow, reduced 

spring precipitation west of the Rockies, and 

increased spring precipitation east of the 

Rockies (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, it could lead to a 

longer duration of warm-season conditions 

(i.e., predominately warm and dry) for the 

Upper Columbia Basin, northern Rockies, 

and parts of the central Rockies. Changes in 

storm-track position and circulation patterns 

will be superimposed on a background of 

natural variability. Thus, various combinations of ENSO, PDO, and broad-scale trends 

could lead to local impacts that vary greatly 

in their magnitude over time.

7.1.6 Ecological impacts

Recent changes such as warming temperatures and associated declines in snowpack 

and surface-hydrology are already influencing ecosystem dynamics. Examples observed 

in the last century include: earlier spring 

blooming and leaf-out, forest infilling at 

and near the treeline, and increased severity 

of disturbances such as wildfire and insect 

outbreaks, all of which are likely to continue 

with additional warming. Spring blooming 

of a number of plant species has occurred 

earlier throughout much of the western 

United States, in some cases by as much as 

several weeks (Cayan et al. 2001; Schwartz 

and Reiter 2000; Schwartz et al. 2006). In 

the northern Rockies, increased density of 

trees at or near treeline has been observed 

at some sites (Butler et al. 2009). This “infill” phenomenon is not uncommon in the 

West 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; Rochefort et al. 1994; 

Millar et al. 2009). While evidence for infill 

is widespread, upslope movement in treeline 
35

The PDO has two phases: warm (positive 

index value) and cool (negative index value). 

Figure 1 shows the sea surface temperature 

anomalies associated with the warm phase 

of the PDO and ENSO, both of which 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 

(not shown) have the opposite patterns: cool 

along the equator and the coast of North 

America and warm in the central north Pacific. Each PDO phase typically lasted for 20 

to 30 years during the 20th century, 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 to 1946 and 

from 1977 to at least 1998. Pacific climate 

changes in the late 1990s, in many respects, 

suggested another reversal from warm to 

cool phase and possibly back to warm.

Natural variation in the strength of PDO 

and ENSO events impacts climate regions in 

different ways. In the Northwest and parts 

of the central 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. 1998). In the southern Rockies and the 

Southwest, 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 and Ghil 1998; 

Mote 2006).

Figure 2. Multivariate ENSO index, 

1950–2009. Positive 

(red) values indicate an 

El Niño event; negative (blue) values a La 

Niña event (Wolter and 

Timlin 1998, 1993).

Figure 1. Warm 

phase PDO and ENSO. 

The spatial pattern 

of anomalies in sea 

surface temperature 

(shading, °C) and 

sea level pressure 

(contours) associated 

with the warm phase 

for 1900–1992. Note 

that the main center 

of action for the PDO 

(left) is in the north Pacific, and for the ENSO 

(right) in the equatorial Pacific. Contour 

interval is 1 millibar, 

with additional contours drawn for +0.25 

and 0.5 mb. Positive 

contours are dashed; 

negative contours are 

solid. Source: Climate 

Impacts Group, University of Washington.

Ocean-atmosphere interactions: PDO and ENSO

Ocean-atmosphere interactions
36

Climate and Ecosystem Change in the U.S. Rocky Mountains and Upper Columbia Basin A Framework for Climate Analysis and Reporting for Greater Yellowstone and Rocky Mountain Networks

position is much more variable, and research 

suggests that it will be characterized by a 

high degree of spatial heterogeneity in relation to other variables that control treeline 

position, e.g., aspect, soils, and micro-topography (Lloyd and Graumlich 1997; Graumlich et al. 2005; Bunn et al. 2005, 2007). 

Changing climate conditions are also influencing disturbance processes that regulate 

ecosystem dynamics. Warming temperatures, 

earlier snowmelt, and increased evapotranspiration are increasing moisture stress on 

forest species and making them more susceptible to insect attack. An increase in the 

extent, intensity, and synchronicity of mountain pine beetle attacks in the western United 

States and Canada has been linked to forests 

stressed by drought, which makes trees less 

able to resist infestations (Nordhaus 2009; 

Hicke and Jenkins 2008; 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 by 

facilitating reproduction and dispersal (Carroll et al. 2004; Black et al. 2010). In some 

cases, past forest management (e.g., factors 

related to structural characteristics of host 

stands) may also facilitate beetle infestation 

(Nordhaus 2009; Black et al. 2010). The rate 

of fire disturbance is also increasing across 

the West, particularly in the northern Rockies (fig. 22, Westerling 2008).

The extent of the western United States 

burned in wildfires each year is strongly 

linked to interannual climate variability (Littell et al. 2008, 2009a; Morgan et al. 2008; 

Higuera et al. 2010). Changes in surface 

hydrology associated with reduced snowpack, earlier spring runoff and peak flows, 

diminished summer flows, and a lengthening 

fire season have all been linked to increased 

frequency of large fires, with the most evident impacts at mid-elevation forests in the 

northern Rockies since the mid-1980s (Westerling et al. 2006, fig. 22).

7.2 Northern U.S. Rocky Mountains

7.2.1 Temperature

Over the course of the 20th century, the 

instrumental record in the northern Rockies showed a significant increase in average 

seasonal, annual, minimum, and maximum 

temperatures (figs. 23, 24; Loehman and 

Anderson 2009; Pederson et al. 2010, submitted). Regional average annual temperatures increased 1–2°C (2–4°F) from 1900 to 

2000 (Pederson et al. 2010). Seasonal and 

annual minimum temperatures are generally increasing much faster than maximum 

temperatures (Pederson et al. 2010). In 

particular, summer and winter seasonal 

average minimum temperatures are increasing significantly faster than the season’s 

average maximum temperatures, causing 

a pronounced reduction in the seasonal 

diurnal temperature range (Pederson et al. 

2010). The magnitude of minimum temperature increases also 

appears seasonally variable: in 

areas with mid-elevation snow 

telemetry (SNOTEL) sites, Pederson et al. (submitted) estimated 

minimum temperature increases 

since 1983 of 3.8 ± 1.72°C (6.8 

± 3.10°F) in winter, 2.5 ± 1.23°C 

(4.5 ± 2.21°F) in spring, and 3.5 

± 0.73°C (6.3 ± 1.31°F) annually (fig. 24). The magnitude of 

changes varies locally, but there 

are few exceptions to this general 

warming trend.

Temperature trends within the 

northern Rockies generally track 

Northern Hemisphere trends 

Figure 22. Annual number of 

forest fires >1000 

acres (total column 

height) in the northern Rockies (black 

area) and other 

western states. 

Source: Westerling 

2008; reprinted 

with permission.
37

across temporal scales 

(fig. 23). This similarity between regional 

and continental trends 

suggests that large-scale 

climate forcings such as 

greenhouse gases, sea 

surface temperature patterns, volcanic activity, 

and solar variability also 

drive regional temperatures (Pederson et al. 

2010).

7.2.2 Precipitation

Throughout the West, high interannual, annual, and decadal variability in 

precipitation exceeds any century-long 

trends (Ray et al. 2008). General patterns 

throughout the latter part of the 20th century indicate that areas within the northern Rockies experienced modest but statistically insignificant decreases in annual 

precipitation (Mote et al. 2005; Knowles 

et al. 2006). Although few statistically 

significant trends are evident in regional 

20th century precipitation, rising temperatures throughout the West have led to an 

increasing proportion of precipitation 

falling as rain rather than snow (Knowles 

et al. 2006). Winter temperatures well 

below 0°C make the northern Rockies less 

sensitive than other western regions where 

small temperature increases in temperature 

are affecting the number of freezing days 

(Knowles et al. 2006).

7.2.3 Surface hydrology

Like most of the western United States, the 

snow water equivalent (SWE) of winter 

snowpack largely controls surface runoff and 

hydrology in the northern Rockies for the 

water year (e.g., Pierce et al. 2008; Barnett 

et al. 2008; Stewart et al. 2005; Pederson et 

al. submitted). Studies have demonstrated 

a statistically significant decrease in winter 

snowpack SWE across the region during the 

second half of the 20th century (Barnett et al. 

2008; Pederson et al. submitted).

7.2.4 Ocean-atmosphere interactions

The warm phase of the PDO was associated 

with reduced streamflow and snowpack in 

the northern Rockies during the 20th century, 

(Fagre et al. 2003), and the cool phase with 

increased streamflow and snowpack (Pederson et al. submitted). One ecological 

response to these shifts has been changes 

in the distribution of mountain hemlock 

(Tsuga mertensiana). At high elevations, 

where mountain hemlock growth is limited 

by snowpack-free days, a warm-phase PDO 

often results in decreased snowpack and 

increased mountain hemlock growth (Peterson and Peterson 2001). At low elevation 

sites where moisture is limiting, a warmphase PDO commonly leads to less moisture and consequently decreased 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 

in the northern Rockies. Winters with high 

Figure 23. Comparison of variability and trends in 

western Montana 

(blue-green line) 

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; reprinted 

with permission.

Figure 24. Average 

winter (Dec–Feb; 

top), spring (Mar–

May; middle), and 

annual (bottom) 

minimum temperatures. Data from 

SNOTEL (water year 

Oct–Sep) and valley MET (Jan–Dec) 

stations. SNOTEL 

station Tmin (minimum temperature) 

records have been 

fit using a nonlinear quadratic 

equation due to 

characteristics of 

these time series. 

All trends shown 

are significant 

(p ≤ 0.05). Note the 

y-axis temperature 

scale changes for 

each panel. Source: 

Pederson et al. 

submitted.

20th century climate change and the instrumental record
38

Climate and Ecosystem Change in the U.S. Rocky Mountains and Upper Columbia Basin A Framework for Climate Analysis and Reporting for Greater Yellowstone and Rocky Mountain Networks

snowpack tend to be associated with the cool 

phase of the PDO, a weakened Aleutian Low, 

and low pressure centered poleward of 45°N 

across western North America (fig. 25). During years of high snowpack, mid-latitude cyclones tend to track from the Gulf of Alaska 

southeast through the Pacific Northwest and 

into the northern Rockies. 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. 

The ENSO is an important driver of snowpack and streamflow at interannual 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 strongly influence 

the timing of streamflow in the northern 

Rockies (fig.25). Springtime geopotential 

heights over western North America influence the amount and more importantly the 

timing of snowmelt and streamflow across 

the northern Rockies. Specifically, high pressure anomalies centered over western North 

America are associated with higher spring 

temperatures and consequently an increasing number of snow-free days and earlier 

arrival of snow melt-out and peak streamflow. Atmospheric circulation changes in 

March and April can, in turn, initiate surface 

feedbacks that contribute to surface temperature and hydrograph anomalies (fig. 25). 

Hence, warming temperatures in the northern Rockies lead to earlier snowmelt and 

runoff and associated decreasing snowpack 

and streamflow, but these patterns can be 

partially attributed to seasonally-dependant 

teleconnections and atmospheric circulation 

patterns, as well as to surface-albedo feedbacks that interact with broad-scale controls 

on snowpack and runoff (Pederson et al. 

submitted).

7.3 Central U.S. Rocky Mountains 

and the Greater Yellowstone Area

7.3.1 Temperature

Temperatures for the CR-GYA increased 

1–2°C (2–4°F) during the last century, with 

the greatest increases occurring in the 

second half of the 20th century (Vose et al. 

2005; Bonfils et al. 2008; Mote 2006, 2003). 

This rate of increase is slightly higher than 

in the Southwest and slightly lower than in 

the northern Rockies, following a pattern of 

more pronounced temperature increases at 

higher latitudes (Cayan et al.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 (Mote 2006; McCabe and Clark 2005; 

Stewart et al. 2004; Hidalgo et al. 2008).

7.3.2 Precipitation

CR-GYA precipitation records show highly 

variable patterns across gradients in elevation, latitude, and longitude. No long-term 

trends are evident over the last century; 

reconstructions of hydrology from tree-ring 

records indicate interannual, decadal, and 

multidecadal variation (fig. 26; Watson et 

al. 2009; Graumlich et al. 2003). A greater 

proportion of precipitation is likely falling as 

rain rather than snow in this region but the 

impacts are less pronounced than in other 

parts of the western United States (Knowles 

et al. 2006). In many parts of the CR-GYA, 

Figure 25. Idealized relationship 

between northern 

Rockies snowpack 

and streamflow 

anomalies with 

associated Pacific 

SSTs, atmospheric 

circulation, and 

surface feedbacks. 

Source: Pederson et 

al. submitted.
39

the 1930s and 1950s were significantly drier than average and the 

1940s wetter, although sub-regional 

variation is high, likely because the 

region is located between atmospheric circulation patterns, as 

discussed further below (Watson 

et al. 2009; Gray et al. 2007, 2004, 

2003; Graumlich 2003). 

7.3.3 Ocean-atmosphere 

interactions

The influence of ocean-atmosphere 

interactions on decadal, multidecadal, and interannual variation in 

climatic conditions is more spatially 

variable in the CR-GYA than in the other regions because it falls in a transition area between northwestern and southwestern U.S. 

circulation patterns (Gray et al. 2007, 2004; 

Graumlich et al. 2003), where variations in 

ocean-atmosphere interactions, topography, 

latitude, and longitude often result in opposite trends in climatic conditions at sites 

within the same region (Gray et al. 2004).

High-elevation snow basins within the 

western GYA typically respond to large-scale 

climate forcing in a manner similar to that of 

the Pacific Northwest, where the cool-phase 

PDO results in cool, wetter than average 

winters and the warm-phase PDO brings 

warmer and drier than average winters (Gray 

et al. 2007, 2004; Graumlich et al. 2003; 

Dettinger et al. 1998). Similar to the Pacific 

Northwest, these areas experience increased 

precipitation during La Niña events and 

decreased precipitation during El Niño 

events, and the ENSO seems to be linked to 

the magnitude of PDO anomalies, especially 

during the winter (Gray et al. 2007). Alternatively, lower elevation sites and eastern 

portions of the GYA respond more like the 

Southwest or show a variable response to 

ENSO that depends heavily on the strength 

of event and interactions with other climate 

drivers (Gray et al. 2004; McCabe et al. 

2007). Years with strong El Niño SSTs have 

increased winter precipitation and La Niña 

events bring drier conditions. This difference 

between high and low elevation precipitation 

regimes is common throughout the central 

Rockies, complicating predictions of future 

precipitation in the region (McCabe et al. 

2007).

7.4 Southern U.S. Rocky Mountains

7.4.1 Temperature

Temperatures have increased 0.5–1°C 

(0.9–1.8°F) throughout the southern Rockies 

during the last 30 years. The north central 

mountains of Colorado warmed the most 

(~1.35 ̊C, 2.43°F) and high elevations may 

be warming more quickly than lower elevations in some regions (Pepin and Lundquist 

2008; Diaz and Eischeid 2007). Warming is 

evident at almost all locations, but 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 shows a 

slight cooling trend during the 20th century; 

no trend is evident in this area for the second 

half of the century (Ray et al. 2008).

7.4.2 Precipitation

Precipitation records for the southern 

Rockies for the last century indicate highly 

variable annual amounts and no long-term 

trends (Ray et al. 2008; Dettinger 2005). 

Like elsewhere in the interior West, a greater 

proportion of precipitation is falling as rain 

rather than snow than in the past, but these 

changes are less pronounced than in the 

northern 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 

Figure 26. (a) Observed annual (previous July through 

current year June) 

precipitation for 

Wyoming Climate 

Division 1 (gray 

line) compared 

to precipitation 

estimates based on 

the stepwise regression model (black 

line). (b) The full 

stepwise version 

of reconstructed 

annual precipitation (black line) 

for 1173 to 1998. 

The horizontal line 

(solid gray) near 

400 mm represents 

the series mean, 

and the vertical 

line (dotted gray) 

at 1258 divides the 

well-replicated portion of the record 

from reconstructed 

values in the earlier, less replicated 

years. Source: Gray 

et al. 2007; reprinted with permission.

20th century climate change and the instrumental record
40

Climate and Ecosystem Change in the U.S. Rocky Mountains and Upper Columbia Basin A Framework for Climate Analysis and Reporting for Greater Yellowstone and Rocky Mountain Networks

land-surface interactions (fig. 27; Stewart 

et al. 2005).

7.4.3 Surface 

hydrology

Similar to trends 

evident throughout 

the interior West, 

more precipitation 

is falling as rain than 

snow in the southern Rockies, spring 

snowpack is decreasing, especially at elevations below 2500 

meters, 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). Summer flows are 

typically lower and annual flows show high 

variability but no significant trends in most 

locations (Ray et al. 2008).

7.4.4 Ocean-atmosphere interactions

Like the CR-GYA, the Colorado River Basin 

spans a transition area where the influence of 

Pacific Northwest and southwestern circulation patterns show opposite trends (Gray et 

al. 2007; Clark et al. 2001). During El Niño 

years, northern parts of this region experience drier than average conditions while the 

southern portions experience wetter than 

average conditions. The opposite conditions occur during La Niña years, 

and anomalies tend to be more 

pronounced in spring in southern 

portions. 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 PDO and 

warm-phase AMO (McCabe et 

al. 2004).

7.5 Upper Columbia Basin

7.5.1 Temperature

For most of the Upper Columbia 

Basin, average annual temperatures increased 0.7–0.8°C from 

1920 to 2003, and the warmest 

decade was the 1990s (fig. 28; Littell et al. 

2009b). Average temperatures have increased 

as much as 2°C (4°F) in parts of the region, 

and increases have been more pronounced 

at higher elevations (Mote 2006, 2003). During the mid-20th century, average and daily 

minimum temperatures increased more in 

the winter and spring than in other seasons 

and more than maximum temperatures. 

During the second half of the 20th century, 

minimum and maximum temperatures 

increased at similar rates (Watson et al. 

2009; Dettinger et al. 1994; Karl et al. 1993: 

Lettenmaier et al. 1994; Littell et al. 2009b).

7.5.2 Precipitation

Upper Columbia Basin precipitation trends 

are less clear than temperature trends, 

and observations indicate high decadal 

variability. Precipitation increased 14% for 

Figure 27. Observed time series 

(1895–2007) of 

annually averaged precipitation 

departures areaaveraged over the 

Upper Colorado 

drainage basin (top) 

and annual Colorado River natural 

flow departures at 

Lees Ferry in million 

acre-feet (bottom). 

The precipitation 

data are based 

on 4-km gridded 

PRISM data. Colorado River natural 

flow data from the 

Bureau of Reclamation. Source: Ray et 

al. 2008; reprinted 

with permission.

Figure 28. Trends 

in average annual 

Pacific Northwest 

temperature, 

1920–2000. Increases (decreases) 

are indicated with 

red (blue) dots, and 

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.
41

the entire northwestern United States, 

(1930–1995), and increases ranged between 13% and 38% within the region 

(fig. 29; Mote 2003), but these trends 

are often not statistically significant, 

depending on the area and time 

interval measured. Similar to much of 

the interior West, variability in winter 

precipitation has increased since 1973 

(Hamlet and Lettenmaier 2007).

7.5.3 Surface hydrology

Spring snowpack and SWE declined 

throughout the Upper Columbia Basin 

in the second half of the 20th century. 

The decline was most pronounced at 

low and mid-elevations, and declines 

of more than 40% were recorded 

for some parts of the region (fig. 30; 

Hamlet et al. 2005; Mote 2006, 2003). Declines in snowpack and SWE are associated 

with increased temperatures and declines in 

precipitation during the same period (Mote 

et al. 2005; Mote 2003). The timing of peak 

runoff shifted 2–3 weeks earlier for much of 

the region during the second half of the 20th

century (Stewart et al. 2004), and the greatest shifts occurred in the mountain plateaus 

of Washington, Oregon, and western Idaho 

(Hamlet et al. 2007). Because Upper Columbia Basin ecosystems rely on the release of 

moisture from snowpack, these shifts are 

significantly impacting plant species, causing some to bloom and leaf out earlier in the 

spring (Mote et al. 2005; Cayan 

et al. 2001; Schwartz and Reiter 

2000).

7.5.4 Ocean-atmosphere 

interactions

Variations in Upper Columbia Basin climatic conditions are related 

to ocean-atmosphere and landsurface interactions, namely the 

ENSO and PDO phenomena. In 

their warm phases, both the ENSO 

and PDO increase the chance for 

a warmer winter and spring in 

the Upper Columbia Basin and 

decrease the chance that winter 

precipitation will reach historical 

averages. The opposite tendencies are true during a cool-phase 

ENSO and PDO: they increase the odds 

that Upper Columbia Basin winters will be 

cooler and wetter than average (Clark et al. 

2001). While strong El Niño years are typically warmer than average, SWE anomalies 

are often less pronounced then, and winter 

precipitations are commonly close to historical averages (Clark et al. 2001). Clark et al. 

(2001) suggested that El Niño circulation 

anomalies are centered more in the interior 

West than are La Niña circulation anomalies 

and are most evident in mid-winter.

Figure 29. Trends 

in average annual precipitation, 

1920–2000. Increases (decreases) 

are indicated with 

blue (red) dots, 

and the size of the 

dot corresponds 

to the magnitude 

of change. Source: 

Climate Impacts 

Group, University 

of Washington.

Figure 30. Trends 

in April 1 SWE, 

1950–2000. Increases (decreases) 

are indicated with 

blue (red) dots, and 

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.

20th century climate change and the instrumental record

43

8.1.1 Small changes can have large 

impacts

Changes in the distribution of minimum 

temperatures and frost-free days illustrate 

how small changes in temperature (1–2oC, 

2–4°F) may result in large changes to surface 

hydrology (Barnett et al. 2004, 2005) as they 

contribute to earlier melt-off and diminished 

spring snowpack (Pederson et al. 2010, 2009; 

Mote 2006; Barnett et al. 2008; Stewart et al. 

2004; McCabe and Clark 2005), increases in 

the proportion of winter precipitation as rain 

rather than snow (Knowles et al. 2006; Bales 

et al. 2006), decreased snow season length at 

most elevations (Bales et al. 2006), and lower 

summer flows (Barnett et al. 2008). Evidence 

from a number of studies suggests that even 

small temperature increases can have dramatic impacts on water availability for much 

of the western United States. Along with 

changes in snowpack and earlier spring runoff, the predicted temperature increases will 

likely contribute to increased drought severity, duration, and frequency (fig. 31; Hoerling 

and Eischeid 2007; Barnett and Pierce 2009). 

Increased winter precipitation predicted for 

central and northern regions of the study 

area will likely be inadequate to offset the 

increased evaporation and plant water use 

driven by rising temperatures (Gray and 

McCabe 2010; Hoerling and Eischeid 2007; 

Seager et al. 2007).

8.1.2 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 the water available in these 

areas depends heavily on the mountain 

snowpack dynamics (Gray and Anderson 

2010). While 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 

Chapter 8: What can we learn from 20th century observations?

Figure 31. Modeled changes in annual mean precipitation minus evaporation (P–E) over the 

Southwest (125°W to 95°W and 25°N to 40°N, land areas only), averaged over ensemble 

members for each of the 19 climate models participating in the Fourth Assessment Report 

(AR4) of the IPCC (Solomon et al. 2007). The historical period used known and estimated 

climate forcings; 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 to 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 ca. 0.3 mm/day. Source: Seager et 

al. 2007; reprinted with permission.

1.1
44

Climate and Ecosystem Change in the U.S. Rocky Mountains and Upper Columbia Basin A Framework for Climate Analysis and Reporting for Greater Yellowstone and Rocky Mountain Networks

The Intergovernmental Panel on Climate 

Change (IPCC, Solomon et al. 2007) included studies to determine whether a detected 

climate change is significantly different from 

natural variations of the climate system. Attribution studies seek to establish the principal causes for observed climate phenomena, including trends in temperature and 

extreme climate events, and whether they 

are related to human activities. In order to 

attribute a detected change, scientists must 

demonstrate that the change is consistent 

with an identified anthropogenic cause and 

inconsistent with any alternative, physically 

plausible explanation that excludes anthropogenic causes (Houghton et al. 2001). If 

attribution is established, the IPCC may 

assign a likelihood for the probability that 

the identified cause resulted in the observed 

conditions.

Attribution studies use empirical analyses 

of past climate relationships and evaluate 

cause-and-effect relationships with climate 

models. Model simulations of past climate 

are compared with the observed record using statistical analysis, including estimates of 

natural variability and trends from climate 

models, historical observations, or paleoclimate reconstructions. “Fingerprint” methods seek the unique signature of climate 

change by simultaneously looking at changes 

in many variables. In studies conducted to 

determine the cause of the observed warming of temperatures in western and northern 

North America over the last half-century, 

annually averaged North American surface temperatures from 1950 to 2007 were 

computed using the IPCC (CMIP3) models 

forced with the observed record of greenhouse gases, volcanic aerosols, and solar 

forcing during 1950 to 1999 and subsequently (2000–2007) with the A1B scenario 

of greenhouse gas emissions. (The A1B 

scenario = rapid economic growth, global 

population that peaks in mid-century and 

declines thereafter, and the rapid introduction of new and more efficient technologies). Comparison of these modeled 

temperatures (fig. 1, top panel) with observations (fig. 19) suggests that anthropogenic 

greenhouse gas emissions have contributed 

about 1°C (1.8°F) of the observed warming in the last 30 years. Similarities between 

the modeled climate and the observed 

trends provide the best available evidence 

for external (rather than natural variability) 

forcing of surface temperature change by 

anthropogenic greenhouse gases because 

the bulk of the warming occurs after about 

1970 in both time series and the externally 

forced warming of about 1°C (1.8°F) since 

1950 is close to the observed rate.

A series of recent studies sought to detect 

and attribute climate change in the western 

United States (Bonfils et al. 2008; Pierce et 

al. 2008; Hidalgo et al. 2009; Das et al. 2009) 

using the same downscaled projections 

and PRISM data as the westwide projections shown below (figs. 35, 36). Bonfils and 

colleagues conducted a detailed analysis of 

Figure 1. Annual average North American 

surface temperature 

(1950–2007) from 22 

IPCC model simulations with greenhouse 

gas, aerosol, solar, and 

volcanic forcing from 

1950 to 1999, and the 

A1B emissions scenario 

from 2000 to 2007 (top 

panel). Annual anomaly 

values of surface temperature averaged 

over the whole of 

North America (compared to 1971–2000 

average, lower panel). 

The smoothed curve 

highlights multi-year 

variations. Source: Ray 

et al. 2008, adapted 

from CCSP 1.3, figure 

3.3; reprinted with 

permission. 

How do we know if observed changes are related to 

human-caused climate change? 
45

Contents

whereby a major 

redistribution of 

vegetative communities ensues (Gray 

et al. 2006; Jackson 

et al. 2009b). 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. 32).

Small increases in temperature (e.g., 1–2oC, 

2–4°F) will result in greater evaporative 

losses from lakes, streams, wetlands and 

terrestrial ecosystems, and it is likely that this 

enhanced evaporation will lead to significant 

ecosystem and water management impacts 

(Arnell 1999; 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 a 15–25% decrease 

in average Yellowstone 

River flows from a 1.5–3°C 

(2.7–5.4°F) 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 can lead to an 

increased frequency of hot weather relative 

to historical conditions and extreme precipitation events (Karl et al. 2009; Solomon et 

al. 2007; Groisman et al. 2005; Kunkel et al. 

2003; Madsen and Figdor 2007; fig. 33). Multiple assessments point to a potential shift in 

precipitation such that storms will 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.

Figure 32. Relationships between 

climate change, 

coping range, vulnerability thresholds, and adaptation. Idealized 

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 AR4 WGI 

2007; reprinted 

with permission.

models thought to best simulate the climate 

of the western United States, and using 

these models, found that natural variability 

is insufficient to explain the increase in daily 

minimum and maximum temperatures, a 

sharp decline in frost days, a rise in degree 

days above 0°C (32°F), and a decline in 

snowpack at low and mid-elevations. They 

ruled out solar variability and volcanic forcing as a cause. They found that the anthropogenic signal is detectable by the mid-1980s 

in a signal-noise ratio of minimum temperature. Other attribution papers focus on 

streamflow (Hidalgo et al. 2009), snowpack 

(Pierce et al. 2008), and the structure and 

detectability of hydrological variables (Das et 

al. 2009). These studies have estimated that 

up to about half of the trends in temperature 

and associated hydrologic variables can be 

attributed to anthropogenic causes (Barnett 

et al. 2008; Pierce et al. 2008).

Figure 33. Schematic for a normal 

temperature distribution showing the 

effect on extreme 

temperatures when 

the mean temperature increases. 

Source: IPCC AR4 

WGI 2007; reprinted with permission.

What can we learn from 20th century observations?

47

Contents

Many of the trends in climate evident in the 

last century are expected to continue in the 

future. Projections shown in this report are 

based on the global climate model (GCM) 

projections done for the IPCC (Solomon et 

al. 2007), a coordinated large set of climate 

model runs known as the Coupled Model 

Intercomparison Project, Phase 3 (CMIP3), 

performed at modeling centers worldwide 

using 22 global climate models. Output of 

most of these models is at large resolution, 

often a 200-kilometer grid. Although it was 

common in past climate impact studies to 

present the results of only one or two global 

climate models, research now suggests that 

the average of multiple models provides a 

better approximation, and use of ensembles 

is made possible by increasing computing 

capacity and technical abilities for analyzing 

multiple model simulations (Salathé et al. 

2010). The CMIP3 models and knowledge 

from these comparisons are the current 

state of the art in climate modeling and 

assessments.

9.1.1 GCM projections for North 

America

These global models project broad-scale 

increases in temperature in North America 

through the mid-21st century. Projected 

changes compared to a recent baseline 

(1950–1999 average) through mid-century 

(2040–2060 average) are shown in figure 

34. For much of the interior western United 

States, the multi-model average projects an 

annual mean warming of about 2°C (4°F, in 

orange) by 2050. Individual global models 

also show a broad-scale pattern of warming, 

though of different magnitudes across models. The range of individual GCM projections (10th and 90th percentiles of the model 

projections) is from about +2.5F to +5.5°F 

(1.4–3.1°C). GCM projections show summers warming by about +5°F (range: 3–7°F, 

2–4°C) and winters by about +3°F (range: 

2–5°F, 1–3°C) (fig. 34, top row). The multimodel average and many individual global 

models show less warming within several 

hundred kilometers of the Pacific coast. This 

Chapter 9: What can we expect in the future?

Figure 34. Temperature and precipitation changes 

in North America 

projected for 2050 

(2040–60 average) 

by an ensemble of 

22 climate models 

used in the IPCC 

AR4. Changes are 

shown relative to 

the 1950–1999 average. The top row 

is the multi-model 

average temperature change for the 

annual mean (left), 

winter (center), and 

summer (right). 

The second row 

shows the percentage change 

in total precipitation (data source: 

CMIP3 multi-model 

dataset, PCMDI). 

The bottom row 

shows multi-model 

agreement. Source: 

Ray et al. 2008; 

reprinted with permission.
48

Climate and Ecosystem Change in the U.S. Rocky Mountains and Upper Columbia Basin A Framework for Climate Analysis and Reporting for Greater Yellowstone and Rocky Mountain Networks

feature may be a result of the inability of 

the global models to simulate the effects of 

the coastal mountain ranges and hence the 

moderating coastal influence penetrates too 

far inland. Regional climate modeling studies 

corroborate this (Salathé et al. 2010), showing 

large values of summertime warming much 

closer to the coast than for the global models.

Although temperatures are expected to 

increase across the landscape, natural 

short-term variation (i.e., years to decades) 

is still expected. One example would be the 

unusually cool Northern Hemisphere winter 

of 2009–10, which is likely linked to a natural 

climate variation called the North Atlantic 

Oscillation (NAO; Hurrell 1995), a largescale circulation feature that can alternately 

block or allow cold air from the Arctic to 

enter the mid-latitudes of North America, 

Europe, and Asia. In 2009–10, the NAO was 

positioned to spawn cold winters as well as 

strong storms and heavy snowfalls. However, when considering the global average, 

the December 2009–February 2010 period 

still ranks as the 13th warmest in the last 131 

years (http://www.ncdc.noaa.gov/sotc/). 

Rates of warming in the Northern Hemisphere have slowed somewhat in the last 

decade, even though the 9 of the 10 warmest 

years on record occurred during this period. 

This phenomenon has been largely attributed to a temporary decrease in the water 

vapor in the lower stratosphere which acts 

as a greenhouse gas (Solomon et al. 2010). 

Because of this natural short-term variability, 

some scientists prefer to report the averages 

of projections for 20 to 30 years, and this is 

done for the projections presented here.

Changes in the amount and spatial distribution of precipitation are still poorly understood and thus difficult to project (Solomon 

et al. 2007), especially against a background 

of substantial year-to-year variability. This 

is largely because precipitation is controlled 

by complex interactions between global and 

hemispheric-scale circulation features and 

ocean-and-land-surface atmospheric interactions that occur 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 storm tracks (Barnett et al. 2008; Bonfils et al. 2008). 

For total annual precipitation, the dominant 

pattern in North America projects a wetter climate in the northern tier and a drier 

climate in the southwestern United States 

(fig. 34, middle row), with small (≤10%) but 

important changes in annual precipitation in 

much of the four climate regions, although 

individual models (not shown) exhibit a 

range of projected changes. The importance of even modest changes relates to the 

timing of increased or decreased precipitation. Summer precipitation is expected to 

decrease for much of the western United 

States, causing increasingly dry warm-season 

conditions, whereas increased precipitation in the winter could result in dramatically higher volume streamflows in parts of 

the Upper Columbia Basin and northern 

U.S. Rocky Mountains. While models are 

in better agreement for projected increased 

winter precipitation for parts of the northern 

Rockies and the Upper Columbia Basin and 

decreases in the Southwest, overall uncertainty remains high (Solomon et al. 2007). 

There is only weak agreement among the 

models as to whether annual precipitation 

will increase or decrease (fig. 34, bottom 

row), but there is an indication of a seasonal

decrease in summer precipitation for parts 

of the four climate regions, and an increase 

in winter precipitation (and more agreement among the models for the latter). In 

the central Rockies and GYA, model results 

vary widely, and it is unclear how conditions 

might change in coming decades. In addition, all of these areas feature pronounced 

natural variability at multiyear to multidecadal scales, and this natural variation may 

mask or enhance the regional expression of 

any broad-scale precipitation trends. 

The CMIP3 projections, which document a 

broad spatial scale of warming, are at large 

resolution (e.g., 200-kilometer [124-mi] grid) 

that is of limited use for assessing impacts 

more locally. Hence, we present results from 

downscaled modeling efforts more relevant 

for regional or local planning: westwide at 
49

Contents

4 kilometers, temperatures at specific sites 

driven by three emissions scenarios (B2, 

A1B, and A2) for three future periods, and 

projections of mid-century conditions for 

three climate and hydroclimate variables 

based on the A1B scenario (rapid economic 

growth, global population that peaks in 

mid-century and declines thereafter, and the 

rapid introduction of new and more efficient 

technologies).

9.1.2 Westwide climate: Statistically 

downscaled projections

For much of the West, GCMs project about 

a 2°C (4°F) rise in temperatures for 2050 

(the orange shading in fig. 34, top row), with 

somewhat less warming near the Pacific 

coast. As part of a project for the U.S. Fish 

and Wildlife Service, NOAA used statistically downscaled projections to illustrate 

what the projected rise in temperatures 

would mean for the western regional climate 

compared to the existing north-south and 

elevational gradients of climate in the West 

(Ray et al. 2010). Downscaled temperature 

data from the CMIP3 22 model average 

projection for the A1B emissions scenario 

(from IPCC AR4) were added to the PRISM 

climatology (~4 km, 2 mi) for the June–

August season. This downscaling method 

makes minimal, physically based corrections 

to the global simulation while preserving 

much of the statistics of interannual variability in the climate model (described by 

Salathé 2005). This method is similar to the 

so-called “delta method,” in which the temperature changes (the “deltas”) from GCMs 

are spatially interpolated and added to a 

high-resolution climatology. 

The maps depict average daily temperature 

for the northern and central U.S. Rocky 

Mountains and Greater Yellowstone Area 

(fig. 35) and the southern Rockies (fig. 36) 

for the 1950–1999 climatology and projections for 20-year averages around 2025, 

2050, 2090. These graphics illustrate that at 

large spatial scales, by 2050 the projected 

changes in summer climate can be visualized 

as a shift of temperature zones northward 

and upward in elevation (3rd panel in each 

figure). This shift of temperature zones continues through the end of the 21st century 

Figure 35. Summer 

observed average 

temperatures and 

statistically downscaled projections 

for the northern 

and central U.S. 

Rockies and Greater 

Yellowstone Area 

(left) and Southern 

Rockies (right). 

Observed average June–August 

temperature for 

1950–1999 (top 

panel). Projections 

were calculated by 

adding the multimodel average temperature changes to 

the 4-km PRISM climatology. Observed 

climatological 

averages are from 

PRISM (DiLuzio et 

al. 2008), projected 

changes from 

the IPCC (CMIP3) 

22-model average 

for the A1B emissions scenario. 

Source: Ray et al. 

2010; used with 

permission.

What can we expect in the future?
50

Climate and Ecosystem Change in the U.S. Rocky Mountains and Upper Columbia Basin A Framework for Climate Analysis and Reporting for Greater Yellowstone and Rocky Mountain Networks

(lower panel in each figure). These maps do 

not illustrate the year-to-year or day-to-day 

variability that will also occur. Furthermore, 

there are a number of unknowns about how 

climate effects may reduce or amplify the 

large-scale pattern of widespread warming 

that is projected over the western United 

States. It is unclear how the details will play 

out at any given location.

9.1.3 Climate projections downscaled 

to specific alpine sites

As part of the USFWS project, NOAA 

also generated temperature projections 

statistically downscaled to 22 mountain 

ranges in the western United States (Ray et 

al. 2010). Graphics for four sites not illustrated in that report are presented below 

(Glacier National Park and the Gallatin 

Mountains, Montana; Niwot Ridge, Colorado; and Clearwater Mountains, Idaho). This 

analysis illustrates implications of modelprojected changes for the seasonal cycle, the 

relationship of projected climate change to 

historical climate variability, the spread of 

the individual model projections, and the 

evolving nature of the ensemble of projections throughout the century. This analysis 

used a modified version of the statistically 

Figure 36. June–August 20-year temperature projectionscentered on 2025 (left panels), 2050 (right panels) for 

Glacier National Park (top panels, elevation 1866m) and the Gallatin Mountains, Montana (elevation 2778m, bottom 

panels) for a 4-km grid cell (approximately 30 x 40 mile). Each graphic compares observed monthly average temperatures to projections for the period. The observed monthly averages (solid black) and 10th and 90th percentiles values 

(dashed black lines) are based on observations during 1950–1999. Projected monthly climatologies (thin red lines) are 

from the multi-model ensemble for the 20-year period centered on 2050. The average of the projections is shown as 

a heavy red line. Data are derived from Maurer 2007. Note that the magnitude of projected temperature change is 

comparable to or greater than variations in the historical record. Source: Ray et al. 2010, used with permission.
51

Contents

downscaled CMIP3 Climate Projections 

created by the Department of the Interior 

Bureau of Reclamation and the University 

of Santa Clara. The statistical downscaling 

technique is known as “bias corrected spatial 

disaggregation” (BCSD) and was originally 

developed for hydrologic impact studies (Wood et al. 2004; Maurer 2002). This 

dataset downscales the projections to a 1/8o 

(12-km, 7-mi) grid (see details in Ray et al. 

2008), which we adapted to the 4-km (2-mi) 

PRISM climatology (DiLuzio et al. 2008) to 

be smaller in scale for ecological applications in mountainous regions. 

The resulting estimates adjusted to PRISM 

are among the best inferences available for 

temperature at this scale, albeit representative of a 4-km (2-mi) average and based on 

interpolation from station observations 

that may be distant from the grid box (See 

Ray et al. 2008 p 30). The results are shown 

in Figures 36–37. Projected temperatures 

from the BCSD/PRISM downscaling for 

the A1B emissions scenario are shown in 

red (thin lines) with the average projection 

(heavy line). (A1B scenario = rapid economic growth, global population that peaks 

in mid-century and declines thereafter, and 

Figure 37. June–August 20-year temperature projections centered on 2025 (left panels), 2050 (right panels) for 

Niwot Ridge, Colorado (top panels, elevation 3267m) and Clearwater, Idaho (bottom panels, elevation 2467m) for a 

4-km grid cell (approximately 30 x 40 mile). Each plot shows observed monthly average temperatures compared to 

projections for that period. The observed monthly averages (solid black) and 10th and 90th percentiles values (dashed 

black lines) are based on observations during 1950–1999. Projected monthly climatologies (thin red lines) are from 

the multi-model ensemble for the 20-year period centered on 2050. The average of the projections is shown as a 

heavy red line. Data are derived from Maurer 2007. Note that the magnitude of projected temperature change is 

comparable to or greater than variations in the historical record. Source: Ray et al. 2010, used with permission.

What can we expect in the future?
52

Climate and Ecosystem Change in the U.S. Rocky Mountains and Upper Columbia Basin A Framework for Climate Analysis and Reporting for Greater Yellowstone and Rocky Mountain Networks

the rapid introduction of new and more 

efficient technologies.) For comparison 

purposes, the 1950–1999 PRISM climatology 

of the monthly average temperature (solid 

black line) and the 10th and 90th percentiles 

(dashed black lines) are also shown. These 

percentiles represent the five warmest and 

coolest months from 1950 to 1999. 

At all four sites (as well as the other 16 not 

shown here), the temperature increases are 

largest in summer. The July temperatures 

from almost all the model projections at 

these sites lie at or above the 90th percentile 

of the present climate. Most of the projections suggest that typical summer temperatures will equal or exceed the extreme warm 

summers of the last half of the 20th century. 

The projected temperature changes are 

somewhat smaller in winter and the year-toyear variations are larger. While the proportion of warm winter months is projected to 

increase, most years, even by 2050, will not 

be above the 90th percentile of the present 

climatology. Winter warming will be manifest 

in the relative absence of months colder than 

the current average and in the cumulative 

effects of consecutive warm winters, with 

an increase in the number of extreme warm 

winter months. 

The spread of the light red lines in figures 

36–37 indicate the range of the individual 

models and the uncertainty in the projection of 20-year average climates, even for a 

given emissions scenario. High-end projections are approximately 1°C warmer than 

the multi-model average, and would indicate 

increased risk at a number of sites. However, the uncertainty implied by this spread 

may be larger than the true uncertainty due 

to differences in climate sensitivity among 

the models studied. The spread about the 

average projection is a result of two factors: 

differences in model climate sensitivity (the 

response of a particular model to climate 

forcing) and model-simulated multidecadal variability. That is why many scientists 

prefer to emphasize the multi-model average projection. Because the BCSD/ PRISM 

downscaling method is based on the CMIP3 

projections, the multi-model average projections shown in these figures are consistent 

with the large-scale patterns of warming in 

the GCM temperature change maps (see fig. 

34, top panel). 

The overall pattern that emerges is for hotter 

summers and somewhat warmer winters. 

The 2050 summer projection is consistently 

about 3°C (5°F) higher than in recent climatology, which is the westwide projected 

increase. The low model projection (the 

10th percentile of the distribution) is in most 

cases higher than the 90th percentile of the 

recent climatology, suggesting that the coolest summers of the mid-21st century will be 

warmer than the warmest summers of the 

recent past. Precipitation projections are 

not provided, but a recent similar downscaling effort for Colorado found that, unlike 

temperature projections, potential changes 

in precipitation are smaller than the yearto-year and decade-to-decade variations 

observed in the historical record (Ray et al. 

2008).

9.1.4 Model projections of future 

climatic and hydrologic conditions

As part of a larger U.S. Forest Service and 

U.S. Fish and Wildlife Service project evaluating future conditions, the Climate Impacts 

Group (CIG, University of Washington) is 

modeling future climatic and hydrologic 

conditions (e.g., SWE, soil moisture, and 

evapotranspiration) for much of the western 

United States (Littell et al. in press). Preliminary results for a few key variables from this 

study are shown in figure 38.

9.1.5 Downscaled model methodology

The CIG project applied a range of climate 

change projections from the WCRP CMIP3 

multi-model dataset used for IPCC AR4 

(Solomon et al. 2007) to hydrologic model 

simulations and evaluated the impact of 

climate change on the hydrology of the 

region (after Elsner et al., in review). These 

models were drawn from a common set of 

simulations of 21st century climate archived 

from 21 GCMs (Mote and Salathé 2010), 

using greenhouse gas emissions scenarios as 

summarized in the IPCC’s Special Report on 

Emissions Scenarios (SRES) (Nakićenović 

and Swart 2000). CMIP3 simulations were 

archived predominantly for three SRES 

emissions scenarios (A1B, B1, and A2) for 
53

Figure 38. Estimate 

of (a) mean SWE, 

(b) mean soil mois-

ture (June–August), 

and (c) mean poten-

tial evapotranspira-

tion (June–August) 

for 1916–2006 and 

2030–2059. Data 

derived from an en-

semble of 8 GCMs 

that perform the 

best over the four 

climate regions. The 

temperature and 

precipitation data 

from the GCMs is 

used to drive the 

VIC hydrologic 

model, and this 

gives the SWE 

values. Great Basin 

and lower Colorado 

are modeled at 1/8o 

(~12 km), the rest 

at 1/16o (~6 km). 

Source: University of 

Washington Climate 

Impacts Group 

(USFS, and USFWS, 

in prep.); used with 

permission.

(a)

(b)

(c)

What can we expect in the future?
54

Climate and Ecosystem Change in the U.S. Rocky Mountains and Upper Columbia Basin A Framework for Climate Analysis and Reporting for Greater Yellowstone and Rocky Mountain Networks

most of the 21 GCMs, with A2 following the 

highest trajectory (most warming) for future 

CO2

 emissions at the end of the 21st century. 

This work focuses on A1B (moderate warming) because it was simulated by the most 

GCMs, and our study focuses on mid-21st

century change, at which point none of the 

scenarios is consistently the highest and for 

which a larger source of uncertainty is the 

variability in GCM models. We chose to 

use the eight GCMs for this study that have 

the best fit to the observed seasonal cycle 

of climate as well as the lowest bias for the 

observed precipitation and temperature records in all three modeled basins: Columbia, 

Missouri, and Colorado.

The spatial resolution of GCM output is 

generally too coarse to be meaningful for 

hydrological studies. Therefore, we downscaled the GCM output to 1/16o (~6 km, 4 

mi) spatial resolution and applied a delta 

method approach to develop an ensemble 

based on the average of the eight scenarios 

(see e.g., Hamlet and Lettenmaier 1999; 

Snover et al. 2003). In the delta method, projected changes in precipitation and temperature as determined by GCM simulations are 

applied to the historical record at the resolution of hydrologic models. We performed 

hydrologic simulations using the historical 

record perturbed by these monthly changes 

in the Variable Infiltration Capacity (VIC) 

model (Liang et al. 1994; Nijssen et al. 1997) 

at 1/16o latitude by longitude spatial resolution over the entire region. The VIC model 

is a macroscale model intended for application to relatively large areas, typically from 

10,000 km2 (3,861 mi) to continental and 

even global scales. A key underlying model 

assumption is that sub-grid scale variability 

(in vegetation, topography, soil properties, 

etc.) can be modeled rather than represented 

explicitly.
55

10.1.1 Climate conditions

● Increasing temperatures (figs. 34–37).

● Increased but highly variable precipitation 

for parts of the Upper Columbia Basin 

and the northern and central U.S. Rocky 

Mountains (figs. 34, 38).

● Highly variable annual precipitation for 

parts of the central and southern U.S. 

Rocky Mountains, with possible decreases; however, recent changes in most 

regions are not statistically significant 

(figs. 34, 38).

● Increased evapotranspiration for most of 

the western United States which is unlikely to be offset by increased precipitation 

(figs. 31, 38; Hoerling and Eischeid 2007; 

Seager et al. 2007).

10.1.2 Surface hydrology

● Larger proportion of winter precipitation 

falling as rain rather than snow (Knowles 

et al. 2006; Bales et al. 2006).

● Decreased snow season length at most 

elevations (Bales et al. 2006).

● Less spring snowpack (fig. 38; Pederson et 

al. submitted; Mote 2006, 2003; Mote et al. 

2005).

● Earlier snowmelt runoff (Stewart et al. 

2005, 2004; Hamlet et al. 2005; Clow 

2007).

● Increased frequency of droughts and low 

summer flows (Gray and Andersen 2010; 

Meko et al. 2007).

10.1.3 Extreme conditions: droughts, 

floods, heat waves

● More episodes of extreme temperatures 

(Parson 2001; Karl et al. 2009).

● Increased frequency of extreme precipitation events, rain-on-snow events, and 

consequent winter and spring floods in 

mountains (Madsen and Figdor 2007; 

Groisman et al. 2005; Kunkel et al. 2003).

● More frequent dry periods as a result of increased temperatures, evapotranspiration, 

and changes to surface hydrology (Gray 

and Andersen 2010; Meko et al. 2007).

10.1.4 Productivity and phenology

● Earlier blooming dates for many plant species (by as much as two weeks; Cayan et al. 

2001; Schwartz and Reiter 2000).

● Longer growing season and increased 

productivity where moisture/soil fertility 

and other factors are not limiting (Bales et 

al. 2006).

10.1.5 Disturbance

● Higher frequency of large fires, longer 

fire seasons, and increased area burned 

by wildfires in the western United States 

(Westerling et al. 2006; Morgan et al. 2008; 

Littell et al. 2008, 2009a; Spracklen et al. 

2009; Higuera et al. 2010).

● Greater drought stress will likely result in 

more insect infestations and disease affecting forests (Black et al. 2010; Nordhaus 

2009; Romme et al. 2006; Logan et al. 2003; 

Carroll et al. 2004; Breshears et al. 2005).

Chapter 10: Climate projections for the western United States
Chapter 11: 
57

Introduction Chapter 11: Planning for the future

Planning for future conditions that are highly 

uncertain presents a significant challenge 

for land managers. However, techniques 

developed for business, finance, and military 

applications offer a roadmap for planning 

in the face of large uncertainties. “Scenario 

planning” is one such approach, and it 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 possible future 

conditions, and as a means for examining 

how management strategies might address 

multiple drivers of change.

Jackson et al. (2009a) developed an example to illustrate this process in which 

alternative futures are arrayed along two 

axes: 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 studies of fire history can place bounds on potential fire size. 

This exercise yields four quadrants, each 

comprising a distinct combination of climatic and fire-regime change (fig. 39). Each 

quadrant provides a contrasting scenario 

or “storyline” for exploring potential impacts 

on species or ecosystems and examining the 

relative costs and benefits of various mitigation and adaptation measures.

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 evident 11,000 

years ago in 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 

Figure 39. 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 

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 defined quantitatively or based on qualitative assessments 

alone. Source: Jackson et al. 2009a; reprinted with permission.
58

Climate and Ecosystem Change in the U.S. Rocky Mountains and Upper Columbia Basin A Framework for Climate Analysis and Reporting for Greater Yellowstone and Rocky Mountain Networks

may reveal that management response may 

be similar for a wide range of possible outcomes. Managers can then move beyond the 

challenges presented by an uncertain future 

to identify management responses that 

address ecological responses to a range of 

climatic conditions.
59

The Protocol Framework

The influence of human activities on the 

climate system, primarily through increased 

greenhouse gases and aerosol concentrations, will be superimposed on natural 

drivers of climatic change. As records of the 

past and model projections for the future 

suggest, our certainty about the patterns of 

climate change varies widely. Temperatures 

will most likely increase over large spatiotemporal scales, but the patterns and direction of change in other climatic variables 

are less clear. Investigating past change also 

instructs us that, at smaller spatio-temporal 

scales, changes in future climatic conditions 

will likely vary greatly, especially in heterogeneous mountain environments characterized 

by steep biophysical gradients. Temperatures 

are predicted to increase in the ensuing 

decades, making much of the western United 

States vulnerable to increased frequency 

and duration of drought (Barnett et al. 2008; 

Seager et al. 2007; Diffenbaugh et al. 2005). 

Small increases in temperature (e.g., 1–2oC, 

2–4°F) will require large increases in precipitation to offset increased evapotranspiration, 

especially in settings where increased temperatures significantly alter available moisture and surface-energy feedbacks (Hoerling 

and Eischeid 2007). 

High levels of uncertainty about how 

ecosystems will respond to changing conditions should not prevent managers from 

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 adaptive resource management. 

For example, management practices that 

focus on maintaining diversity, increasing 

connectivity, providing buffer habitat around 

protected areas, and mediating human 

impacts within and around public lands will 

also facilitate the dynamic and heterogeneous redistribution of vegetation that will 

likely occur across the West.

Paleoenvironmental records suggest that 

ecosystems are responsive to climatic change 

over millennial to decadal time and spatial 

scales, but they also suggest that we should 

anticipate novel vegetation assemblages as 

species respond individualistically to climate change and communities reorganize 

(Whitlock et al. 2003; MacDonald et al. 2008; 

Williams et al. 2007; Jackson et al. 2009b). 

These changes also suggest that natural 

disturbances, non-native species, disease, 

and unforeseen synergistic interactions will 

be associated during periods of ecological 

disequilibrium. Management actions that 

work to accommodate dynamic and difficult to predict redistributions of vegetation 

by protecting a full suite of biophysical and 

environmental gradients, increasing connectivity, and minimizing human impacts in 

landscapes surrounding protected areas may 

best protect valuable resource attributes in 

the face of climate change.

Chapter 12: Summary conclusions

61

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