CHAPTER 3

WATER AND CLIMATE CHANGE IN MONTANA

Wyatt F. Cross, John LaFave, Alex Leone, Whitney Lonsdale, Alisa Royem, Tom Patton, and Stephanie McGinnis 

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Water is the lifeblood of Montana. We depend on an adequate supply of clean water for nearly every aspect of our lives, including food production, hydroelectric power, domestic and industrial uses, and sustaining our treasured natural ecosystems. Water is also strongly influenced by climate, as changes in temperature and precipitation consistently alter patterns of water availability and quality throughout the state. It is thus critical that we understand the impacts of climate change on Montana’s water resources. This chapter synthesizes scientific information on how climate change is influencing the supply and distribution of water in Montana. The information presented here represents an essential first step—understanding what’s changing—within the longer-term, iterative process of adapting and improving our resilience to the challenges of an uncertain climate future.

 

	KEY MESSAGES Montana’s snowpack has declined over the observational record (i.e., since the 1930s) in mountains west and east of the Continental Divide; this decline has been most pronounced since the 1980s. [high agreement, medium evidence]11    Warming temperatures over the next century, especially during spring, are likely to reduce snowpack at mid and low elevations. [high agreement, robust evidence]   Historical observations show a shift toward earlier snowmelt and an earlier peak in spring runoff in the Mountain West (including Montana). Projections suggest these patterns are very likely to continue into the future as temperatures increase. [high agreement, robust evidence]    Earlier onset of snowmelt and spring runoff will reduce late-summer water availability in snowmelt-dominated watersheds. [high agreement, robust evidence]   Long-term (decadal and multi-decadal) variation in total annual streamflow is largely influenced by patterns of climate variability; the influence of climate warming on these patterns is uncertain. [high agreement, medium evidence]   Total annual streamflows are projected to increase slightly for most Montana rivers, but the magnitude of change across the state and agreement among models vary. [medium agreement, medium evidence]    Local responses of groundwater resources to climate change will depend on whether aquifers are directly sensitive to climate variability, are buffered from climate by water-use practices such as irrigation, or are used to meet water demands that exceed or replace surface water supplies. [high agreement, robust evidence]    Groundwater demand will likely increase as elevated temperatures and changing seasonal availability of traditional surface-water sources (e.g., dry stock water ponds or inability of canal systems to deliver water in a timely manner) force water users to seek alternatives. [high agreement, medium evidence]    Multi-year and decadal-scale droughts have been, and will continue to be, a natural feature of Montana’s climate [high agreement, robust evidence]; rising temperatures will likely exacerbate drought when and where it occurs. [high agreement, medium evidence]   Changes in snowpack and runoff timing will likely increase the frequency and duration of drought during late summer and early fall. [high agreement, medium evidence]   A warming climate will strongly influence Montana’s snowpack, streamflow dynamics, and groundwater resources, with far-reaching consequences for social and ecological systems. [high agreement, medium evidence]

 

 

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SNOWPACK

 

Key Message

Montana’s snowpack has declined over the observational record (i.e., since the 1930s) in mountains west and east of the Continental Divide; this decline has been most pronounced since the 1980s. Warming temperatures over the next century, especially during spring, are likely to reduce snowpack at mid and low elevations. [high agreement, robust evidence]

 

The influence of climate on snowpack is one of the major linkages between climate change and water supply. Snowpack in the mountains of Montana stores and provides water to downstream users and ecosystems in both the US and Canada. Water generated by Montana’s snowpack travels to the Pacific, Atlantic, and Arctic oceans. Indeed, western Montana is often called the Crown of the Continent because headwater streams originating there give rise to the major rivers that drain three of North America’s largest watersheds, those of the Columbia, Missouri-Mississippi, and Saskatchewan rivers. 

Precipitation that falls at higher elevations during the cold winter months accumulates as snow until spring when temperatures increase and snowmelt begins. In Montana’s mountainous areas, winter snowfall represents the majority (62-65%) of total annual precipitation (Serreze et al. 1999), while in the eastern plains, the contribution of snowfall to total precipitation is considerably less (WRCC undated). All of Montana’s major rivers that contain headwaters above 7000 ft (2100 m) elevation are considered snowmelt-dominated systems in which precipitation as snow is a primary driver of year-to-year variability in streamflow. This snowpack acts as a natural reservoir, slowly releasing water during the spring and early summer, sustaining approximately 2 million acres (0.8 million ha) of irrigated farmland in Montana (Pierce et al. 2008; Vano et al. 2010; USDA-NASS 2015). A sufficient supply of water (especially during the summer) is not only important for maintaining Montana’s agricultural industry, but it also underpins our natural ecosystems and the state’s rapidly growing tourism economy (Power and Power 2015, 2016). 

Most of Montana’s annual snowfall arrives from mid October through mid May (although snowfall has been observed in all 12 months in the mountains of Montana). Snowfall is strongly influenced by local and regional climate. Average annual snowfall varies considerably throughout the state, from roughly 20 inches (0.5 m) in the plains of northeastern Montana, to over 400 inches (10.1 m) in several mountain locations in the west (WRCC undated). (Note that annual snowfall totals are higher than annual precipitation totals—as in Figure 3.2—because of the different physical properties of frozen versus liquid water). 

 

Measuring snowpack

Reliable snowpack measurements are essential for estimating water supply and assessing the risk of drought or floods (MT DNRC undated). The US Department of Agriculture’s (USDA’s) Natural Resources Conservation Service (NRCS undated) measures Montana’s snowpack through two networks:

• Over 90 automated SNOwpack TELemetry (SNOTEL) sites.—First established in Montana during the early-1970s, SNOTEL sites gather high-resolution data year-round, and remotely transmit snowpack and climate information every hour.
• Roughly 100 Snow Course survey locations.—First established in Montana in the 1920s-1960s, Snow Course data consist of hand-collected snowpack measurements. These measurements, typically gathered near the first of each winter month, provide our longest direct records of regional snowpack. 

Scientists usually report snowpack as snow water equivalent (SWE). SWE represents the amount of liquid water contained within a column of snow or, more precisely, the height of water that would remain in a standardized area if the snowpack melted. When examining multi-year trends in snowpack, scientists and managers often use the April 1 SWE measurement to represent peak snowpack and total accumulated cold-season precipitation. Although April 1 SWE values can underestimate actual peak snowpack in the Northern Rockies (Bohr and Aguado 2011), this metric functions as a reasonable approximation for maximum snowpack at the watershed scale (Serreze et al. 1999; Pederson et al. 2011b) and as an indicator of potential spring streamflow in Montana. 

April 1 is considered an optimal date for examining trends because it is the most continuously collected date in the observational record; some sites have been recorded continuously for over 80 yr (Mote et al. 2005). 

 

Montana’s diverse geography and topography influence patterns of snowpack accumulation and snowmelt

Geography.—In Montana, the Continental Divide exerts a marked influence on climate patterns and resulting snowpack:

• Areas west of the Continental Divide typically exhibit milder winters, cooler summers, and a longer growing season due to the influence of warm Pacific air masses (see Climate chapter). Figure 3-2 shows that average annual precipitation is highest west of the Continental Divide (MT DNRC 2015). As a result, total water yields and water yield relative to watershed area are greatest in climate division 1 (Figure 3-5) (MT DNRC 2015).
• Areas east of the Continental Divide experience more extreme seasonal temperature fluctuations and a shorter growing season due to greater influence by drier continental air masses (see Climate chapter) (WRCC undated). 

Topography.—Mountains west of the Continental Divide are generally situated at lower elevations than those east of the Continental Divide, yet the western mountains still receive more snowfall on average each year (Figure 3-2). SNOTEL stations record the highest snowfall totals west of the Continental Divide in the Kootenai, Flathead, and Clark Fork basins. Several of these stations are located at relatively low elevations (5000-6500 ft [1500-1980 m]), but receive over 40 inches (1 m) of SWE each year. By comparison, the highest annual snowpack totals east of the Continental Divide (20-35 inches [0.5-0.9 m] of SWE) are generally located at elevations over 8000 ft (2400 m) (NRCS 2016). 

For most mountainous areas in Montana, SWE typically peaks in April or early May, but this can vary depending on elevation, aspect (e.g., north versus south facing), and relative position west or east of the Continental Divide. Low-elevation SNOTEL sites west of the Continental Divide in the Kootenai Basin (approximately 4200 ft [1280 m]) typically record maximum snowpack at the end of March and snow is absent by early May. In contrast, high-elevation sites (over approximately 8500 ft [2590 m]) in the headwaters of the Yellowstone and Missouri basins exhibit peak SWE values in mid May, and some north-facing slopes can retain snow through the end of June (NRCS 2016). 

Snowpack accumulated at high elevations tends to be more stable and persist longer than at low elevations, largely as a result of colder temperatures at high elevations. Snowpack at higher elevations is also less prone to melt during short warm spells in the early spring that can degrade snowpack at lower elevations. 

 

Long-term variation in snowpack and the importance of ocean-atmosphere linkages

As discussed in the Climate chapter, large-scale atmospheric patterns associated with changes in sea-surface temperatures are largely responsible for variation in Montana’s weather and climate (Cayan et al. 1998; Abatzoglou 2011; Pederson et al. 2011a; Pederson et al. 2013a). Phase shifts in the Pacific Decadal Oscillation can be readily detected in the long-term records of annual snowfall. On shorter time scales, and layered on top of Pacific Decadal Oscillation variation, the Pacific North American pattern and the El Niño-Southern Oscillation cycles (see Climate chapter) can also affect variation in snowpack. During El Niño episodes, Montana tends to experience warmer-than-average temperatures and below-average precipitation, especially during the winter and spring. These anomalies decrease snowpack and result in early snowmelt (Climate Prediction Center 2016). In contrast, La Niña episodes typically result in below-average temperatures, above-average precipitation, and above-average snowpack. Exceptions to these patterns certainly exist.

 

Observed regional trends in snowpack

Regional trends in April 1 SWE demonstrate that average annual snowpack has declined in large portions of the American West over the period of reliable measurement (1930s to present; Figure 3-7) (Mote 2003; Hamlet et al. 2005; Mote 2006; Casola et al. 2009; Mote and Sharp 2016). Some regions, such as low-elevation sites in the northern Rocky Mountains (including Montana) and the Cascades, have experienced more drastic reductions than other sites, such as high-elevation loctions in the Sierras and central Rocky Mountains. 

 

Figure 3-7. Trends in April snowpack in the western US, 1955-2016. Red bubbles indicate areas with declining snowpack; blue bubbles indicate areas with increasing snowpack. The diameter of the bubbles is proportional to the percentage change between 1955 and 2016. Figure from Mote and Sharp (2016).

 

However, it is important to place these recent observations in the context of much longer-term (multi-century) changes in climate. Climate reconstructions based on tree-ring measurements provide a robust tool for producing quantitative comparisons of past and present climate (Fritts 2012). One such recent reconstruction showed that declines in snowpack since the 1950s are unusually severe and synchronous across the West when viewed in the context of the past 1000 yr (Figure 3-8) (Pederson et al. 2011b). Separate studies have suggested that these recent declines in snowpack can be directly attributed to elevated greenhouse gas emissions and associated warming (Barnett et al. 2008; Pierce et al. 2008).

 

Figure 3-8. Snow water equivalent (SWE) reconstruction for the Northern Rockies based on tree-ring measurements (figure from Pederson et al. 2013a). Z-scores standardize the data to represent the number of standard deviations above or below the long-term average.

 

In the Rocky Mountains, spring (February-March) warming since the 1980s has been largely responsible for recent snowpack declines at mid- and low-elevation sites (Pederson et al. 2013b). Most studies agree that general declines in snowpack across the West have resulted from warming spring temperatures (Mote 2003; Hamlet et al. 2005; Mote et al. 2005; Abatzoglou 2011; Kapnick and Hall 2012; Pederson et al. 2013a; Lute et al. 2015); however, declines in winter precipitation may also be important (Clow 2010). If spring temperatures continue to warm as projected (see Climate chapter), snowpack is likely to decline even further. 

 

Observed trends in Montana’s snowpack

We conducted an updated assessment of Montana’s April 1 SWE to assess variability and trends over the full record—from the late 1930s to present—of NRCS Snow Course observations.¹⁵ We used Snow Course rather than SNOTEL data because of its longer period of record. Although studies have shown a strong correlation between Snow Course and SNOTEL April 1 SWE measurements (Serreze et al. 1999; Bohr and Aguado 2001; Dressler et al. 2006; Pederson et al. 2011a), all snowpack measurements have limitations, including the potential for human measurement error, land-use change over the period of record, poor representation for watersheds with highly complex terrain, and misrepresentation of high-elevation sites that lack measurement stations (Gillan et al. 2010; Silverman and Maneta 2016). Nonetheless, Snow Course trends are more dependable and useful for describing long-term changes to snowpack across Montana’s large mountainous watersheds. 

Long-term cycles in April 1 SWE are evident among all Snow Course groupings (Figure 3-9). These cycles can be attributed in part to decadal-scale climate patterns, including the Pacific Decadal Oscillation. For example, the Pacific Decadal Oscillation was in a negative phase between approximately 1960-1980, leading to above average snowpack across the state. In contrast, from the mid 1980s through approximately 2000, a positive phase led to relatively low snowpack years (Figure 3-9).

Large interannual variability in snowpack can be nested within Pacific Decadal Oscillation (and Pacific North American) driven patterns (e.g., see the high snow years of 1996 and 1997 that occurred during a 25-year period of below average snowpack).

 

Figure 3-9. Normalized April 1 SWE based on Snow Course measurements west and east of the Continental Divide. The upper panel in each column shows data summarized from all Snow Course stations west or east of the Continental Divide. The middle and lower panels show patterns of SWE at high or lower elevations. Black lines represent simple downward trends and are not meant for statistical inference.

 

 

Snow Course groupings on both sides of the Continental Divide show long-term downward trends in April 1 SWE (Figure 3-9, Table 3-2). This observation is consistent with other studies that have described shrinking snowpack volumes in Montana and elsewhere in the western US (Mote et al. 2005; Pederson et al. 2013a; Mote and Sharp 2016). In general, April 1 SWE in Montana has declined roughly 20% over the last 80 yr, and this decline is most pronounced at lower elevation sites (Table 3-2). 

Table 3-2. Linear trends in snowpack for particular elevations east and west of the Divide, calculated from data in Figure 3-9.¹⁶

		West			East
	All sites	>6000 ft (>1830 m)	<6000 ft (<1830 m)	All sites	>7000 ft (>2100 m)	<7000 ft (<2100 m)
Decline in inches (cm) of SWE/decade	-0.48 (-1.2)	-0.36 (-0.91)	-0.45 (-1.1)	-0.33 (-0.84)	-0.29 (-0.74)	-0.35 (-0.89)
Decline (percent over the record)	-19%	-12%	-23%	-20%	-14%	-27%

 

 

In general, April 1 snow water equivalent in Montana has declined roughly 20% over the last 80 yr, and this decline is most pronounced at lower elevation sites. 

 

Montana’s snowpack is particularly sensitive to warming

Both empirical studies and model projections demonstrate that snowpack in the Northern Rockies and inland Pacific Northwest is more vulnerable to warming than some other regions in the West. For example, Mote and Sharp (2016) showed that western Montana and the Pacific Northwest have experienced the most drastic declines in snowpack volume in the West over the past 80 yr. 

Unlike Rocky Mountain regions to the south, Montana stores a significant amount of snowpack at mid and low elevations (below 8000 ft [2400 m]), particularly within the Flathead, Kootenai, and lower Clark Fork basins of northwestern Montana. In regions such as these, the projected increase in temperatures will result in reduced winter snowpack and a higher-elevation snowline (Regonda et al. 2005; Klos et al. 2014). 

Low elevations west of the Continental Divide are exposed to relatively warm Pacific air masses. These regions have experienced an increase in precipitation falling as rain instead of snow since the 1950s (Knowles et al. 2006), a trend that is expected to continue under future climate conditions (Barnett et al. 2005). 

 

Snowpack projections for Montana

Here we present projections of April 1 SWE values for three of our focal snowmelt-dominated basins in Montana (Figure 3-10). Projections consist of two future scenarios of greenhouse gas emissions (Representative Concentration Pathways [RCPs], RCP4.5 and RCP8.5; see Climate chapter), for two periods in the future: 2040-2069 and 2070-2099. Model results were compared to a historical baseline period of 1970-2000. These projections highlight the general direction of projected changes and differences among watersheds across the state. Values near the dotted 0% line would represent model results that project no future change in snowpack relative to 1970-2000 data. While our results demonstrate relatively strong agreement among most of the climate models, they should only be used to project the direction of change and not specific future snowpack volumes.

 

Figure 3-10. APRIL 1 SWE projections for three snowmelt-dominated basins in Montana under two scenarios (RCP4.5 and RCP8.5) and two time periods (2040-2069 and 2070-2099). Data are presented as the projected percent change in April 1 SWE between the baseline period 1970-2000 and two future time periods (2040-2069: upper panel; 2070-2099: lower panel). Box and whiskers plots show variation in projections among the different models. These types of plots appear in other graphs below that depict model projections.  The line in the middle of the boxplot represents the median value of all model projections. The bottom and top of the box represent the 25th and 75th percentiles (or first and third quartiles), respectively, of model projections. The upper whisker (line extending from the box) extends from the box to the largest model value no further than 1.5*IQR from the box (where IQR is the inter-quartile range, or distance between the first and third quartiles). The lower whisker extends from the box to the smallest model projection that is no further than 1.5*IQR of the hinge. Few model projections fall beyond the end of the whiskers (i.e., outliers), and these are not shown in the figures.  For explanation of specific confidence levels, refer to Future Projections in Water Chapter.

 

Our projections show that

• snowpack volumes for the Montana basins studied will very likely decline in the future; and 
• the largest projected changes in snowpack appear to be located west of the Continental Divide, and are the same areas that have experienced the largest declines in April 1 SWE over the past 80 years (see, for example, the Clark Fork River in Figure 3-9). 

Small headwater basins west of the Continental Divide show this vulnerability because they occupy relatively low elevations that are likely to experience additional days with temperatures above the freezing point. In contrast, many small headwater basins east of the Continental Divide are at higher elevations (often 8000-10,000 ft [2400-3000 m]) and are thus less likely to experience temperatures above freezing during the winter-spring transition (see Climate chapter). 

 

 

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TOTAL ANNUAL STREAMFLOW

 

Key Messages

Long-term (decadal and multi-decadal) variation in total annual streamflow is largely influenced by quasi-cyclic changes in sea-surface temperatures and resulting climate conditions; the influence of climate warming on these patterns is uncertain. [high agreement, medium evidence]

Total annual streamflows are projected to increase slightly for most Montana rivers, but the magnitude of change and agreement among models vary across the state. [medium agreement, medium evidence] 

 

To this point, our streamflow discussion has focused on how climate influences the timing and distribution of flow throughout the year. However, it is also important to consider patterns of annual streamflow (or annual water supply), which is the total amount of runoff generated by a given watershed throughout an entire year. 

Annual streamflow derives from a variety sources including rainfall, snowmelt runoff, groundwater discharge, and glacial runoff. Annual streamflow is critical because it defines the potential volume of water available each year to influence groundwater, fill reservoirs and lakes, and support consumptive and non-consumptive uses of water. Climate-induced changes in annual streamflow have the potential to impact hydroelectric power generation, agricultural production, wildlife habitat, recreation, and other beneficial uses of Montana’s water resources.

Long-term records demonstrate that annual streamflow varies widely over time due to changes in both natural and human-related factors. Here again, we focus largely on atmospheric processes that influence annual streamflow (i.e., temperature and precipitation), which are themselves modified by both natural and human-related factors, such as greenhouse gas emissions. 

Interannual variation in precipitation tends to have the greatest influence on year-to-year variation in annual streamflow volumes (Karl and Riebsame 1989; McCabe and Wolock 2011). Years of high snowpack accumulation or high spring and summer rains tend to produce high annual streamflow volumes and hence greater potential water supply. While temperature effects on annual streamflow are much weaker and less consistent than the effects of precipitation, the relative importance of temperature is likely to increase as the climate warms (Tesemma et al. 2015; Woodhouse et al. 2016). 

 

Observed trends in total annual streamflow

On decadal time scales, annual streamflow variation and precipitation are driven by large-scale patterns of climate variability, such as the Pacific Decadal Oscillation (see teleconnections description in Climate chapter) (Pederson et al. 2011a; Seager and Hoerling 2014). Periods of high or low precipitation associated with these patterns generally translate to periods of high or low annual streamflow, respectively (Karl and Riebsame 1989; McCabe and Wolock 2011). Obvious periods of lower-than-average streamflow in most of the focal rivers in Montana include the drought years of the Dust Bowl (late 1920s to early 1940s), the 1950s, the late 1980s to early 1990s, and the early 2000s (Figure 3-13). 

Specific years of above- and below-average streamflow differ slightly among river basins due to Montana’s geographic diversity and the varying influence of large atmospheric circulation patterns east and west of the Continental Divide. Interestingly, large semi-cyclic patterns in total annual streamflow are detectable in most Montana rivers, suggesting that parallel changes in water use over time have not been large enough to mask these climate-driven trends. Such patterns in annual streamflow, however, are often hidden for rivers that are highly regulated by dams or large irrigation withdrawals (e.g., the Marias River below Tiber Reservoir) (MT DNRC 2014c).

A key question is whether annual streamflows have changed over time in Montana and, if so, why. Pederson et al. (2011b) reported no recent change in annual streamflow for a number of rivers in the northern Rocky Mountains, including in Montana, despite significant reductions in snowpack. The authors attributed this finding to recent increases in spring precipitation that may have offset reduced snowpack. 

Luce and Holden (2009) reported declines in annual streamflow during the driest years (i.e., lowest 25^th flow percentile) for a set of Pacific Northwest rivers, including some rivers in Montana west of the Continental Divide. Other work has suggested that streamflow declines in the Pacific Northwest have resulted from reduced mountain precipitation rather than warming (Luce et al. 2013). 

 

Figure 3-13. Long-term patterns of total annual streamflow in our focal rivers. Each panel shows the annual discharge (gray line) expressed as cubic feet per second or CFS (metric unit is m3/s). The blue and red lines show the percentage deviation above (blue) or below (red) the long-term average for each year. The dark black line represents the 5-year moving average. The red shading represents the most significant periods of hydrologic drought for each focal river.

 

Norton et al. (2014) also reported recent (1960-2010) declines in annual streamflow for 30 rivers in the Upper Missouri watershed. It is unclear whether these declines are attributed to changes in climate or other factors such as changing patterns of land or water use (e.g., conversion of agricultural lands to subdivisions, or changing irrigation methods and practices). 

 

Factors that influence total annual streamflow

At local scales and over shorter periods, annual streamflow responds to seasonal changes in climate variables (e.g., temperature, precipitation) and related processes such as evapotranspiration. The relative influence of these factors on annual streamflow differ across the state’s watersheds. Thus, identifying the most important factors that influence annual streamflow in each basin can help us understand how changing climate may influence future water supplies.

Analysis for the Upper Missouri River Headwaters Study quantified climate factors impacting annual streamflow variability for a number of important watersheds in Montana (USGS undated; analysis contributed by Connie Woodhouse, University of Arizona). The work, which covered 1936-2010, considered average monthly temperatures and total monthly precipitation for the water year (prior October to September) as possible predictors of annual streamflow. Importantly, this particular analysis was conducted with streamflow data corrected for water use and human modification (i.e., naturalized flows). Thus, these results provide a window into the direct climate factors that impact streamflow. A summary of results explaining interannual variability of streamflow follows.

• For the Missouri River at Toston, fall-to-early-winter (October-January) precipitation accounts for nearly half (45%) of the variability in total annual streamflow (Figure 3-14).¹⁸ Years of high winter precipitation lead to years of high annual streamflow. Snowmelt season (March-July) temperatures explain an additional 14% of the interannual variation in flow, likely because warmer temperatures during this time can lead to greater evapotranspiration and reduced annual streamflows. 
• In contrast, for the Marias River basin near Chester,¹⁹ spring-to-early-summer temperatures, not fall-to-winter precipitation, account for the largest amount (40%) of annual streamflow variation (Figure 3-14), although reasons for this observation are not entirely clear. Prior November-January precipitation is the second most important factor (17% of the variation), with spring-early summer (May-June) precipitation being third (5%).
• For the Powder River near Locate, May-June precipitation accounts for close to half (43%) of the annual variability in streamflow, probably because southeastern Montana receives the majority of its annual precipitation in the spring and early summer. January precipitation accounts for 15% of annual variation, showing that winter precipitation from the Powder River’s alpine headwaters in Wyoming is also important.
• For the Yellowstone River at Corwin Springs,²⁰ at the northern edge of Yellowstone National Park, fall and early winter precipitation (Oct-Jan) account for 70% of the annual variability. Runoff season temperatures have a minimal effect on annual flow, perhaps because much of the watershed consists of high-elevation terrain and snowpack that is less affected by variation in spring temperatures. 

The Missouri River analysis suggests that large snow years that are associated with cold air temperatures during runoff lead to the greatest annual water supplies. The Marias River analysis suggests that rising spring temperatures could exacerbate low flows in the Marias during years of below-average precipitation. All analyses include significant other variation that remains unexplained, suggesting that a) the models do not include all factors influencing annual streamflow, and b) there is some observational and statistical error. 

 

Figure 3-14. Climate factors associated with naturalized streamflow in four Montana river basins. The size of pie pieces correspond to how strong the particular climate factor influences total annual streamflow. Some of these factors lead to greater flow (positive), while others lead to reduced annual flow (negative). See text for further explanation.

 

 

Annual streamflow projections

Model projections for annual streamflow (Figure 3-15) show little agreement among models regarding the direction and magnitude of change in our two focal rivers west of the Continental Divide (Middle Fork of the Flathead and Clark Fork at St Regis). In contrast, projections show moderately-high to high agreement that total annual streamflow will increase east of the Continental Divide (e.g., Missouri River, Yellowstone River, Musselshell River), especially under the RCP8.5 emission scenario (see Climate chapter).

 

Figure 3-15. Total annual streamflow projections for the focal rivers under RCP4.5 and RCP8.5 for 2040-2069. Data are presented as the projected percent change in runoff between 2040-2069 and the baseline period of 1970-2000.  (Boxplots are explained in the caption of Figure 3-10.) For explanation of specific confidence levels, refer to Future Projections in Water Chapter.

 

 

 

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DROUGHT

 

Key Messages

Multi-year and decadal-scale droughts have been, and will continue to be, a natural feature of Montana’s climate [high agreement, robust evidence]; rising temperatures will likely exacerbate drought when and where it occurs. [high agreement, medium evidence]

Changes in snowpack and runoff timing will likely increase the frequency and duration of drought during late summer and early fall. [high agreement, medium evidence] 

 

Drought is a recurrent climate feature that occurs throughout the world, although it varies substantially from region to region. Drought can have broad and potentially devastating economic and environmental impacts (Wilhite 2000); thus, it is a topic of ongoing, statewide concern. 

Through time, Montana’s people, agriculture, and industry, like its ecosystems, have evolved with drought. Today, many entities across the state address drought, including private and non-profit organizations, state and federal agencies, and landowners, as well as unique watershed partnerships. These groups incorporate drought preparedness and management goals into Montana water policy (see Building Drought Resilience sidebar). 

Drought is a complex phenomenon, driven by both climate and human-related factors. Although a clear definition of drought is elusive, most definitions fall into four interrelated categories: 

• meteorological drought, defined as a deficit in precipitation and above average evapotranspiration that lead to increased aridity;
• hydrological drought, characterized by reduced water levels in streams, lakes, and aquifers following prolonged periods of meteorological drought;
• ecological drought, defined as a prolonged period over which an ecosystem’s demand for water exceeds the supply (the resulting water deficit, or shortage, creates multiple stresses within and across ecosystems);²¹ and
• agricultural drought, commonly understood as a deficit in soil moisture and water supply that lead to decreased productivity (in this assessment, we will treat this form of drought as an important component of ecological drought).

Here, we focus on hydrological drought, in keeping with the emphasis on water availability and streamflow. 

Drought is also discussed in terms of its duration. This section will address both persistent drought, which we define as multiple years of below-average streamflow (within which individual seasons of above-average flow may occur); and seasonal drought, defined as below-average streamflow lasting months. In Montana, seasonal drought is most common and of greatest concern during the warm growing season in summer and early fall. We therefore refer to this as warm-season drought.

 

	Building Drought Resilience in Montana In the past, Montana often addressed drought and flooding as temporary emergencies, with reactionary responses to an immediate crisis. Over the past decade, western water planners have learned that the best time to prepare for the impacts of drought or flooding is before those events occur. Thus, they have developed plans advocating a proactive hazard management approach.  Today, Montana’s drought planning efforts take this approach, seeking to apply foresight, commitment, technology, and cooperation to diminish the impacts of drought. For example, water managers and users now employ improved short-term drought forecasting methods to better plan for and mitigate drought impacts. Even so, our forecasting abilities must further improve for Montanans to better prepare for short-term variation in weather patterns and expected long-term impacts associated with climate change.   Canoeing the Jefferson River. Photograph courtesy of Scott Bischke.

National Drought Resilience Partnership Montana Demonstration Project Through the Montana Demonstration Project partners, Montana is forging new ground in bringing together agencies, resource managers, and communities to plan for drought impacts and build drought resilience within watershed communities. Teaming with the National Drought Resilience Partnership—a collaborative of federal and state agencies, watershed stakeholders, and non-governmental organizations—the Montana group is working to leverage and deliver technical, human, and financial resources to help address drought in the arid West.  The team selected the Missouri Headwaters Basin in southwest Montana for a national demonstration project. This basin experiences frequent drought, plays an important role in landscape connectivity in the Northern Rockies, and faces rapidly changing population and land use. The Montana Demonstration Project partners are working collaboratively to engage and train community-based drought coordinators to lead planning, mitigation, and project implementation in each of the eight watersheds in the basin. The individual watershed planning efforts will provide the basis for a scaled-up, integrated Headwaters Basin plan.  This unique partnership is successfully demonstrating a) the value of enhanced coordination, and b) how to effectively leverage federal, state, and private resources to build community and ecosystem resilience to prepare and adapt to a changing climate. Text contributed by Ann Schwend, MT DNRC.

 

Persistent drought

During the past century, Montana experienced significant periods of persistent drought in 1917-1919, the late 1920s to early 1940s (the Dust Bowl droughts), the 1950s, the late 1980s to early 1990s, and the early 2000s (Figure 3-13). Most consider the Dust Bowl drought to be the worst multi-year drought in the observational record in Montana. 

While these major droughtss affected all of Montana—and indeed much of the West—the severity, duration, and timing of each drought varied across the state (Figure 3-13), including in their influence on the larger river basins (MT DNRC 2015). In the 1950s, for example, rivers east of the Continental Divide show multiple years of below-average flow, while the Clark Fork at Saint Regis experienced above-average flows during the entire decade. 

Studies of tree-ring-based reconstructions of drought, snowpack, and streamflow offer important insights about the long-term history of drought, as well as the natural variability of climate over the last millennium (Jackson et al. 2009). Reconstructions of snowpack, streamflow, and drought indices show that wet and dry periods, persisting between 10-25 yr, occurred throughout the western US (Cook et al. 2004; Pederson et al. 2006; Pederson et al. 2011b). Further, tree-ring reconstructions of drought (using the Palmer Drought Severity Index) show that a) the frequency and duration of droughts in the West were greater prior to 1200 AD than during the 20^th century, and b) the past 300 yr have been wet relative to the long-term average (Cook et al. 2004; Cook et al. 2010). 

 

Regional and local factors that influence persistent drought 

A complex interplay of climate, hydrologic and ecosystem processes, and human impacts influences drought. For this assessment, we focus on the first two. However, it is notable that humans significantly impact streamflow and water supply, and, hence, patterns of drought. Those impacts must be included in future efforts to assess drought risk, manage water use and supply, and build resilience to climate change. 

Natural variability in precipitation and temperature will continue to characterize Montana’s climate in the future, resulting in droughts of varying duration and intensity. Within the context of this natural variability, human-driven changes in temperature and precipitation will affect future patterns of drought in Montana. For this assessment, we focus on important factors, described below, that affect the natural variability of persistent drought in Montana, as well as potential shifts in drought occurrence as a result of a changing climate.

• Precipitation.—Interannual variability in precipitation is widely accepted as the primary climate factor driving drought. While annual precipitation is expected to increase in many parts of Montana, precipitation projections are less certain than changes in temperature, making accurate assessment of future drought risk based on those projections difficult (Cook et al. 2014). Additionally, the total volume of annual precipitation is only a single factor that helps to predict drought; the frequency, intensity, character, and seasonality of precipitation are equally important (Sheffield and Wood 2008). For example, shifts from snow to rain in headwater areas and potential decreases in summer precipitation could have negative consequences for water supply in the seasons of highest water demand (see snowpack, snowmelt and runoff, and seasonal drought sections). 
   
• Temperature.—Temperature variability can also affect drought, although its influence is much smaller than that of precipitation (Dai 2011; Livneh and Hoerling 2016). Historically, temperature appears to be a secondary response to drought, rather than an initial driver. In the prolonged absence of precipitation, soils dry out and the fraction of energy that once went into evaporation heats the land surface and forces temperatures higher (Lukas et al. 2014). In severe circumstances, a positive feedback then occurs, with high temperatures further exacerbating the drought. Several recent studies suggest that, while precipitation remains the primary driver of drought, the influence of high temperatures on drought is increasing, as shown for recent droughts in California and the Great Plains (Hoerling et al. 2014; Wang et al. 2014; Livneh and Hoerling 2016).
   
• Evapotranspiration and drought.—Rising temperatures cause increased rates of evaporation and plant transpiration, which together are referred to as evapotranspiration. This increased evapotranspiration will be one of the most significant influences on drought resulting from rising temperatures. In the absence of increased precipitation, higher rates of evapotranspiration can move substantial amounts of water back to the atmosphere (Figure 3-1), leading to reductions in streamflow, soil moisture, and groundwater recharge. Recent studies suggest a) a global trend toward drying of land surfaces since the 1980s (Sheffield and Wood 2008; Dai 2011; Dai 2013), and b) an increase in water deficits in the Northern Hemisphere since 2000 (McCabe and Wolock 2015), both resulting from rising temperatures and elevated levels of evapotranspiration. However, regional changes in evapotranspiration are less certain than global trends (Cook et al. 2014). Additionally, quantifying the effects of climate change on evapotranspiration—and subsequently to the water balance—is complex; so much so that future projections of drought risk vary significantly (Zwiers et al. 2013; Sheffield et al. 2012) depending on assumptions made about how evapotranspiration will respond to climate change. Among other factors, complexity results from uncertainty in how plants will respond to elevated greenhouse gases and changes in water availability, as shown below.
  • Evapotranspiration is expected to increase with warming and, yet, plants can respond to elevated atmospheric CO₂ (that occurs in parallel with warming) by using water more efficiently leading to less water loss through evapotranspiration (Tesemma et al. 2015; Swann et al. 2016).
  • Evapotranspiration is limited by water supply and, thus, long-term or seasonal increases in aridity will constrain potential increases in evapotranspiration caused by rising temperatures (Huntington and Niswonger 2012; Trenberth et al. 2014).

Surface water stored in reservoirs provides important warm-season water supplies in much of Montana, particularly in central and eastern areas that receive the lowest levels of average annual precipitation. A significant percentage of this stored water is currently lost to evaporation. For example, 7% of the water budget for the lower Missouri River Basin evaporates annually from Fort Peck Reservoir (MT DNRC 2014a). Additionally, in the many arid parts of Montana, runoff efficiency—the proportion of precipitation converted to streamflow—is already low (e.g., 4% of precipitation in the Musselshell leaves the basin as streamflow [MT DNRC 2014a]). Higher rates of reservoir evaporation due to rising temperatures could exacerbate both problems, resulting in reduced water supply and decreased ability for reservoirs to buffer summer periods of low streamflow.

 

Drought and the dominant role of sea-surface temperatures

As discussed in the Climate chapter, large-scale atmospheric circulation patterns connected to changes in sea-surface temperatures strongly influence natural variations in precipitation and temperature (e.g., Cayan et al. 1999; Mantua and Hare 2002). Shifts in sea-surface temperatures in both the Pacific and Atlantic oceans can produce conditions that lead to periods of drought (McCabe et al. 2004, Seager and Hoerling 2014). 

A deeper understanding of these circulation patterns is required to predict persistent drought in Montana and the West accurately (Cook et al. 2007; Trenberth et al. 2014). The relationship between changes in sea-surface temperature and drought is complicated by many factors, including a) the large number of meteorological or other environmental phenomena involved; b) the widely varying timescales and large distances those phenomena act over; and c) the fact that those phenomena can amplify or dampen each other’s effect on weather and climate (Schubert et al. 2016). Indeed, our current understanding of how sea-surface temperatures respond to climate change is relatively weak (see Climate chapter), severely limiting our ability to forecast persistent drought (Dai 2011; Seager and Hoerling 2014; Trenberth et al. 2014). 

 

Likelihood of persistent drought

Given the known occurrence of long-term drought in Montana over the observed historical and paleo-climate records, there is very high likelihood that persistent drought will continue to be part of Montana’s future climate, regardless of the effects of climate change. 

There is relatively little consensus about how climate change will affect the incidence of persistent drought at global and regional scales, in large part due to the uncertainties discussed above. In addition, exclusion of human-related impacts such as irrigation, land use, and water diversion from most current climate models makes reliable projection of drought even less certain (Sheffield and Wood 2008). Across the western US, there is considerable variation in projected future drought risk, both regionally and among climate models. 

Projections for the northern Rocky Mountains and northern Great Plains, including precipitation only, suggest that long-term droughts will not increase in frequency (Strzepek et al. 2010). However, projections for these regions that incorporate other changes in climate (such as temperature and evapotranspiration) predict increasing drought frequency in the latter half of the 21^st century (Strzepek et al. 2010; Dai 2011; Cook et al. 2014), suggesting an increasing influence of temperature on drought. Amidst debates over changes in drought frequency, there is widespread agreement that rising temperatures will exacerbate drought when and where it occurs, leading to more rapid onset of drought and increased intensity (Strzepek et al. 2010; Peterson et al. 2013; Lukas et al. 2014; Trenberth et al. 2014). Such effects may already be occurring in some areas (Lukas et al. 2014).

In addition, strong evidence exists that climate change is likely to impact the occurrence and severity of warm-season drought (Cook et al. 2004; Sheffield and Wood 2008; Pederson et al. 2011a; Dai 2013; Trenberth et al. 2014).

 

Warm-season drought

Drought during the warm season is a common phenomenon in arid and snowmelt-dominated regions in the West, including much of Montana. In these areas, the majority of precipitation arrives as snow in winter and melts in spring to produce high streamflow that generally diminishes through summer. Even for areas in eastern Montana that receive most of their annual precipitation in spring and summer, water in larger watersheds is predominantly derived from mountain snowpack (MT DNRC 2014b), and thus streamflow there follows a similar seasonal pattern (Figure 3-6). 

Warm-season drought can occur during years of persistent drought as well as years of average precipitation if, for example, high spring temperatures rapidly reduce snowpack. Changes in late-summer flows are likely to be more critical to people than changes in annual flows because the demand for water is highest in summer. Thus, understanding current trends and potential changes in warm- season drought is essential for building water resource resilience in Montana. 

 

Observed trends in warm-season drought

Evidence for declining summer flows across much of the West comes from large-scale studies in the Sierra Nevada; the Columbia, Colorado (Das et al. 2011), and Upper Missouri basins (Norton et al. 2014); and many small watersheds in western Montana, Idaho, Alberta, and British Columbia (Rood et al. 2008; Leppi et al. 2012). Widespread declines in August streamflow and increased frequency of low flows have been reported in both pristine and regulated watersheds of western Montana, Idaho, and Wyoming, with the most pronounced trends in pristine sites (Leppi et al 2012). 

Most studies link declining summer flows with increased winter and spring temperatures, reduced snow accumulation (see snowpack section), and earlier snowmelt and spring runoff (see snowmelt and runoff section) (Rood et al. 2008; Kim and Jain 2010; Leppi et al. 2012). The problems of declining summer flows are compounded in watersheds with significant water use.

 

Factors associated with low summer flows in Montana 

Here, we investigate factors associated with summer low flows in our focal rivers by examining correlations between historical (1929-2015) climate and August streamflow (e.g., the relationship between winter or spring precipitation and August flow; Figure 3-22). While this investigation cannot be used to predict August streamflow in the future, it can help reveal patterns of seasonal temperature and precipitation that tend to produce higher or lower summer flows. It should be acknowledged, however, that increased water use during the summer makes it harder to explicitly separate the effects of climate from water use.

Factors that determine August flows vary across the state.—In our focal watersheds, we find that for rivers fed primarily by mid- to high-elevation snowmelt (e.g., the Yellowstone at Billings, the Clark Fork at Saint Regis, and the Missouri at Toston), August flows have a strong positive relationship with winter (November-March) precipitation and a strong negative relationship with spring (April-June) temperatures. Thus, we can expect that years with low winter precipitation and high spring temperatures will lead to low August flows. In addition, for the Yellowstone at Billings, when recent years (1980-2015) are compared to the entire period of record (1929-2015), the relationship between high spring temperatures and low August flows (not pictured in Figure 3-22) is strengthened, a finding consistent with literature suggesting that recent warming is exacerbating low summer flows (Hay et al. 2011; Leppi et al. 2012; Huntington and Niswonger 2012).

Although these three rivers receive the vast majority of their annual precipitation in winter and spring (70-78%), summer precipitation can also have an important influence on August flows (Figure 3-22). For the Marias River near Shelby and the Musselshell River at Mosby, August flows are lowest during years of high summer temperatures (Figure 3-22) and low summer precipitation. For the Marias River near Shelby, high spring temperatures also negatively influence August flows, likely because of accelerated snowmelt. 

 

Figure 3-22. Relative influence of temperature and precipitation on August flows for the focal rivers of this assessment. In general, warmer temperatures have a negative influence on August streamflow, while precipitation has a positive influence on flows. Differences exist among seasons and rivers.

 

For the two focal rivers in eastern Montana (Powder River near Locate and Poplar River near Poplar), August flows are most dependent upon summer precipitation (Figure 3-22), likely because eastern Montana receives a higher percentage of precipitation in spring and summer relative to the other parts of the state.

Climate projections and warm-season drought.—Future changes in climate are likely to increase the incidence of warm-season drought (Cook et al. 2004; Sheffield and Wood 2008; Pederson et al. 2011a). Based on the relationships between climate and August flows described above, projected warming in winter and spring will likely lead to lower summer flows and/or low flows of longer duration. High warm-season temperatures show a negative relationship with August flows in several watersheds, an effect that is likely to be magnified with rising summer temperatures and the projected increase in number of days over 90°F (32°C) (see Climate chapter).

Any potential decline in summer precipitation is also projected with medium confidence for many areas in Montana (see Climate chapter). Small decreases in summer precipitation could exacerbate the occurrence and severity of warm-season drought because a) many smaller watersheds in eastern Montana are fed more by spring and summer precipitation than by winter snowpack (MT DNRC 2014b), and b) low August flows show a strong correlation with summer precipitation in all focal watersheds. 

Projected hydrographs for our focal watersheds demonstrate reductions in late-summer flows for all rivers except the Musselshell, although variation in the magnitude of change and projection confidence exists (see Figure 3-12). These projections, therefore, generally illustrate that a higher proportion of the annual flow will leave Montana watersheds earlier in the year, resulting in lower flows during the summer months. Lower flows are a concern for multiple reasons, as described below.

• Although Montana has experienced a long history of warm-season drought, projected changes in temperature and precipitation could have a substantial impact on the severity of warm-season drought in the future. 
• Short-term drought during the season of highest demand can a) test water supply infrastructure, and b) have severe consequences for human and natural systems (Luce and Holden 2009). 
• Given the projected increases in streamflow during winter and spring, maintaining streamflows during warm season months will likely necessitate reconsideration of water storage practices and reservoir management.
• Changing seasonality of water availability will likely put additional stress on the rigid and legally encumbered water rights system, making it difficult to access water at critical times (Udall 2013). 
• Changes in stream temperature due to lower flows and rising air temperature are likely to have catastrophic impacts on some aquatic species, with ripple effects on Montana’s important river-based recreation industry (see Warming Rivers and Streams sidebar). 

Warming Rivers and Streams Montana’s 40,000 miles (64,000 km) of perennial rivers and streams support world-famous trout fisheries. They also provide habitat for rare and temperature-sensitive species like bull trout (Salvelinus confluentus), cutthroat trout (Oncorhynchus clarkii), and pearlshell mussels (Margaritifera margaritifera).  Researchers have recently developed high-resolution stream climate maps (Isaak et al. 2016) based on extensive stream temperature data collected by several agencies across the state. The maps show that summer stream temperatures vary considerably throughout the state. Those temperatures generally reflect patterns in average air temperatures—usually being coldest in the high mountains and warmest at low elevations and in the eastern plains.  Changes in climate described in this assessment, especially declining summer flows (Rood et al. 2008; Leppi et al. 2012) and increasing air temperatures (Pederson et al. 2010), have caused temperatures to increase in the state’s rivers and streams at the rate of 0.18-0.36°F (0.1-0.2°C/decade) (Isaak et al. 2012). Stream warming rates are slower than air temperature warming rates due to the buffering effects of groundwater, but any temperature increase can be important for cold-blooded aquatic species. Two ramifications of increasing stream temperatures follow. Fish moving upstream.—Studies already show that distributions of brown trout (Salmo trutta) and bull trout have shifted upstream as fish seek to access cooler habitats (Eby et al. 2014; Al-Chokhachy et al. 2016). In addition, warm-water fishes (e.g., smallmouth bass [Micropterus dolomieu]) have been caught with increasing frequency in historically cold sections of some rivers, such as the Yellowstone River near Livingston. Cold headwater streams are poised to provide important climate refuge for species requiring cold waters (Isaak et al. 2015).  Possible changes in Montana fishing regulations.—In larger rivers at lower elevations, warming trends may result in more frequent fishing season closures and disease outbreaks, such as the mountain whitefish (Prosopium williamsoni) kill on the Yellowstone River in the fall of 2016 (MFWP 2016, Wright 2016). Some sections of rivers that currently support trout fisheries may transition gradually into bass fisheries. It is difficult to know precisely how much warmer Montana’s rivers and streams will become in this century, but across the state water temperatures will likely follow rising air temperatures.   Climate map showing average summer temperatures in Montana’s rivers and streams during historical baseline period of 1993–2011 (top panel) and a late-century scenario in which temperatures are warmer than historical conditions by 3.6°F (2.0°C; bottom panel). River segments colored red are usually too warm for popular cold-water species like trout. A more detailed version of this map, as well as the stream temperature data used in it, are available from the NorWeST website (USFS-RMRS undated). Sidebar text and figure contributed by Daniel J. Isaak, US Forest Service, Rocky Mountain Research Station.

 

Flooding.—While drought likely represents the greatest persistent water-resource concern in Montana, flooding has also occurred regularly throughout the state’s history, resulting in loss of life and substantial damage to property, infrastructure, and riparian ecosystems. Flash flooding events typically occur with little warning, are difficult to predict, and are caused by a variety of climate and human-related factors. The geographical extent of flooding is often more limited than that of drought; flood history in Montana therefore varies significantly by watershed and basin (Table 3-3).

In Montana, flood events can occur at any time of the year, but the causes of flooding vary among seasons. Most spring floods are caused by rapid snowmelt, particularly during rain-on-snow events in which rain infiltrates and degrades the existing snowpack. The most severe and destructive floods in the state’s observational record have resulted from rain-on-snow events (Table 3-3) (Paulson et al. 1991; MT DNRC 2015) that occurred after a period of relatively cold weather. During these events, frozen soils prevent the infiltration of surface water into soils, resulting in greatly elevated runoff (MT DNRC 2015). In northwestern Montana, rain-on-snow events are one of the most frequent causes of annual maximum streamflows (MacDonald and Hoffman 1995; Ferguson 2000). 

Natural variability in precipitation also plays a significant role in flooding in Montana, sometimes in combination with rain-on-snow events. For example, in the huge flood of June 1964, 13 inches (33 cm) of rain fell in 24 h near Augusta, Montana, nearing the average annual precipitation for the region in a single day. This record-breaking rainfall on a higher-than-average and late-melting snowpack along the Rocky Mountain Front caused the overtopping of the Gibson Dam on the Sun River, and the failure of Swift Dam on Birch Creek and Lower Two Medicine Dam on Two Medicine Creek. The flood caused extensive damage and resulted in 30 fatalities, all of which were on the Blackfeet Reservation.

Flood events can also occur in winter as a result of ice jams, which impede flow in the river channel and lead to floodplain inundation. These floods are most common east of the Continental Divide during persistent cold weather fronts. Interestingly, Montana has recorded more ice jam events than any other state in the continental US (US Army CRREL undated). Floods that occur during summers are generally caused by large convective rainstorms and are most common in the eastern plains, particularly in the lower Yellowstone River Basin. 

Human factors play a significant role in modifying flood regimes. Activities, such as urbanization, forest clearing, wetland drainage, and stream channelization, tend to amplify flooding, while water management practices, such as reservoir storage operation, can often prevent or moderate the peak flows that lead to large floods (Kunkel at al. 2003; Rood et al. 2016). Consideration of these factors will be critical for preventing and mitigating floods into the future.

Table 3-3. Montana flood history from 1908-2011 from the National Water Summary and recent observations (Paulson et al. 1991).

Date	Area affected	Recurrence interval (yrs)	Description	Cause
June 1908	Clark Fork Basin, Missouri headwaters	Unknown	Widespread and severe; worst flood until 1964; lives lost, 6	Excessive spring rains and snowmelt runoff
Sept 1923	Powder River basin	Unknown	Largest known discharge at Moorhead.	Intense rain
May-June 1948	Clark Fork, Flathead, Kootenai basins	25 to 50	Severe	Intense rain and rapid snowmelt runoff
April 1952	Milk River basin	25 to >100	Severe on Milk River main stem	Rapid snowmelt runoff
May-Jun 1953	Missouri headwaters	25 to100	Moderate to severe	Intense rain and rapid snowmelt runoff
June 1964	Missouri headwaters (Sun, Teton, Marias), Clark Fork basin	50 to >100	Worst on record; lives lost, 30 (all on the Blackfeet Reservation)	Intense rain and rapid snowmelt runoff
Jan 1974	Kootenai River basin	25 to >100	Severe on several Kootenai tributaries	Intense rain and rapid snowmelt runoff
May-July 1975	Missouri headwaters, Clark Fork basin	25 to 100	Severe in most areas affected in 1964	Intense rain and rapid snowmelt runoff
May 1978	Bighorn, Powder and Tongue basins	10 to >100	Severe on larger tributaries; lives lost, 1	Intense rain and rapid snowmelt runoff
May 1981	West-central Montana	10 to >100	Severe; centered on Helena	Intense rain and rapid snowmelt runoff
Sept 1986	Milk River basin	10 to >100	Severe on larger tributaries; lives lost, 1	Intense rain
June 1996	Statewide	50 to 100	Ice jam flooding (Feb) in the Clark Fork and Yellowstone basins and widespread spring flooding	Intense rain and rapid snowmelt runoff
June 1997	Statewide	50 to 100	Severe and widespread flooding impacting wide geographic area,	Record statewide snowpack, Intense rain and rapid snowmelt runoff
May-June 2011	Statewide	50 to >100	Largest flood of record for Missouri River near Wolf Point and Yellowstone River near Livingston. Flooding most severe on the Musselshell River.	Above average mountain snowpack, spring rainfall, and extensive and late melting prairie snowpack

 

Trends in flood-related precipitation.—Determining trends in flood events and their underlying causes is difficult due to the complex interplay of climate and human-related factors. Many studies have therefore examined flood-related precipitation events instead (Karl and Knight 1998; Kunkel 2003; McCabe et al. 2007). 

Studies have shown an increase in global and North American extreme precipitation events since 1970 (Karl and Knight 1998; Peterson et al. 2013; Rood et al. 2016). In the US, increases in extreme precipitation have been most substantial in the East, while trends in the West appear to be mixed and location specific (Salathé et al. 2014). 

McCabe et al. (2007) analyzed the frequency of rain-on-snow events at over 400 sites in the western US between 1949 and 2003 and found declining trends at lower elevations and increasing trends at higher elevations. Reductions in rain-on-snow events at lower elevations are attributed to declines in the extent of low-elevation snowpack caused by warming (McCabe et al. 2007). Increases at higher elevations are likely due to a high-elevation snowpack that has been largely unaltered by warming (Mote et al. 2005), combined with increased variability of late-fall-winter precipitation.

The above studies and others (e.g., Hamlet and Lettenmeier 2007) suggest that change in flood risk during the latter half of the 20^th century has been a function of both precipitation (increased variability) and temperature (warming in mid winter). Climate-driven changes in both of these variables will continue to affect flood risk in the future. 

Future flood risk.—Warming will continue to reduce mountain snowpack, and this could reduce flood risk related to rain-on-snow events by reducing the quantity of water available for release stored as snow (Cohen et al. 2015). Yet warming is also likely to increase the amount of winter and spring precipitation that falls as rain (particularly in rain-snow transition zones), which will accelerate snowmelt and could increase flood risk, depending on antecedent snowpack, soil moisture, and other conditions. As such, rising temperatures alone will influence flood risk, regardless of trends in precipitation (Salathé et al. 2014); yet the effects will likely be location- and event-specific and therefore difficult to predict (Cohen et al. 2015).

Future precipitation projections show a general increase in extreme events at a global scale (Min et al. 2011; Rood et al. 2016), and regional climate models also consistently predict increases in extreme precipitation in the northwestern US. In Montana, the frequency of wet events (days with more than 1 inch [2.5 cm] of rain) and variability in interannual precipitation are both projected to increase slightly by mid to late century (Figures 2-25 and 2-20). 

There is considerable uncertainty surrounding future flood risk in response to climate change, and some research suggests that extreme precipitation events can actually intensify more quickly than what is projected by general circulation models (Min et al. 2011; also see section on GCMs in Climate chapter). Additionally, flood risk depends on specific storm characteristics that are difficult to capture in most models (Salathé et al. 2014). Moreover, the particular effects of projected changes in temperature and precipitation on flood risk will depend on location, elevation, and antecedent weather conditions, as well as human practices that impact flooding.

 

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CONCLUSIONS

 

Future changes in climate will alter Montana’s hydrology. Although the specific magnitude of changes remains uncertain, two conclusions regarding the general character of changes can be made with high confidence:

• Rising temperatures will reduce accumulation of snowpack, shift historical patterns of streamflow in Montana, and likely result in additional stress on Montana’s water supply, particularly during the summer and early fall.
• Rising temperatures will exacerbate persistent drought periods that are a natural part of Montana’s climate.

The goal of this chapter has been to provide scientific data on the impacts of climate change on Montana’s water supply, which are crucial to the health of Montana’s agriculture, industry, municipalities, and human and natural ecosystems. Building resilience for the future will require: 

• cooperation between legislators, planners, scientists, managers and water users across the state;
• a water use system that is flexible and able to adapt to changes in timing of water supply; 
• a focus on other means for natural and artificial storage of water for use during times of high demand; and
• explicitly addressing the issue of water use and demand in conjunction with best data on climate and water supply 

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¹¹ A reminder that throughout the MCA we assess our confidence in the key messages by considering a) the level of agreement among experts with relevant knowledge, and b) the quality of the evidence. We use these two factors and the criteria described in the National Climate Assessment to assign the confidence ratings expressed in this chapter. See sidebar titled “Expressed Confidence in MCA Key Messages” in the Introduction chapter.

¹² 1 acre-foot is 325,851 gal (1233 m³), enough water to cover an acre of land 1 ft (0.3 m) deep.

¹³ For more detail on our focal rivers and watersheds, see Appendix 3-1 on the MCA website.

¹⁴ The characterization of climate division 4 focuses on the plains basins within the division.

¹⁵ For more detail on our snowpack analysis methods see Appendix 3-2 on the MCA website.

¹⁶ For more detail on our snowpack analysis methods see Appendix 3-2 on the MCA website.

¹⁷ Center of timing refers to the calendar date at which half the total annual volume of streamflow has passed a given streamflow
gauging station.

¹⁸ For an explanation of methods see Appendix 3-2 on the MCA website.

¹⁹ Marias at Chester is the closest downstream gage to Marias at Shelby, one of our focal gages. Gages in this particular analysis were selected by the Basin Study of the Missouri River watershed (see earlier sidebar). 

²⁰ Yellowstone at Corwin Spring is not one of our focal rivers, but it is included here to show additional variation in drivers of annual flow.

²¹ Ecological drought has also been defined recently by a Science for Nature and People Partnership working group. Their definition is a prolonged and widespread deficit in naturally available water supplies that creates multiple stresses across ecosystems (Science for Nature and People Partnership undated). 

 

 

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