The waters of the world’s oceans are subject to a variety of forces that create regional and local variations in sea level. Winds and currents move water laterally in the ocean, creating anomalous spatial patterns of sea level that can persist for a decade or longer. The high winds and low atmospheric pressures associated with El Niños and other climate patterns can significantly elevate sea level along the west coast of the United States for intervals of several months, as well as generate damaging high waves and storm surges. Melting of glaciers and ice sheets adds new water to the oceans and the associated gravitational and deformational effects distribute it nonuniformly, raising sea level in some areas and lowering it in other areas. Geologic processes (e.g., tectonics, compaction) and human activities (e.g., withdrawal of groundwater) also raise or lower the coastal land surface, increasing variability in relative (or local) sea-level rise.
This chapter evaluates the current contributions of ocean circulation, short-term climate patterns and storms, modern land ice change, and vertical land motion to sea-level rise in California, Oregon, and Washington. The discussion draws largely from published studies on the variability of sea level in this region, although the Intergovernmental Panel on Climate Change (IPCC) Fourth Assessment Report also summarizes research results on ocean circulation and short-period climate changes in the northeast Pacific Ocean. This chapter concludes with the results of the committee’s analysis of tide gage records along the west coast of the United States.
Satellite altimetry data provide unambiguous evidence of significant regional differences in sea-level change in the oceans (Bindoff et al., 2007; Milne et al., 2009; Appendix B). Spatial variability in the North Pacific Ocean is associated with climate patterns—primarily the El Niño-Southern Oscillation (ENSO) but also the Pacific Decadal Oscillation (PDO; Box 4.1)—which affect ocean surface heating, surface air pressure, and wind patterns, and thus change ocean circulation (e.g., Mantua and Hare, 2002; Bond et al., 2003; Cummins and Freeland, 2007). Changes in ocean circulation change sea levels on seasonal to multidecadal timescales by redistributing mass and altering temperature and salinity in the upper ocean.
Estimates from the IPCC Fourth Assessment Report
Satellite altimetry records assessed by the IPCC showed that sea level fell about 0–6 mm yr-1 from 1993 to 2003 along the U.S. west coast and rose by 6 mm yr-1 to ~12 mm yr-1 in the tropical western Pacific Ocean (Bindoff et al., 2007). Temperature data from the upper 700 m of the ocean showed a similar sea-level pattern for the same period, indicating that regional sea level is influenced by changes in the thermal structure of the upper ocean, which are associated with changes in ocean circulation and surface heating. The IPCC (2007) suggested that the largest fraction of this short-term variation was caused by ENSO. Over longer periods, however, the thermosteric sea-level pattern along the U.S. west coast was different, showing a rise
Pacific Ocean Climate Patterns
ENSO. The El Niño-Southern Oscillation is a quasi-periodic climate pattern that occurs across the tropical Pacific Ocean about every 2 to 7 years. It is characterized by variations in the sea-surface temperature of the tropical eastern Pacific Ocean. In the warm El Niño phase, warm ocean temperatures in the tropical eastern Pacific are accompanied by high air surface pressures in the tropical western Pacific (Figure). In the cool La Niña phase, the pattern is reversed. The reversal in surface air pressure between the eastern and western tropical Pacific is known as the Southern Oscillation.
FIGURE (Top) Sea-surface temperature anomalies (shading) and sea-level pressure (contours) associated with the warm phase of ENSO (i.e., El Niño) for the 1900–1992 period. Positive contours are dashed and negative contours are solid. (Bottom) Multivariate ENSO index for 1950–2009. The index is based on variables observed over the tropical Pacific, including sea-level pressure, surface wind, sea surface temperature, surface air temperature, and cloudiness. Positive (red) index values indicate El Niño events and negative (blue) values indicate La Niña events. SOURCE: Figure and details on how the index is computed are given in <http://www.esrl.noaa.gov/psd/enso/mei/>.
PDO. The Pacific Decadal Oscillation is often described as a long-lived (i.e., decadal) El Niño-like pattern of Pacific climate variability. Like ENSO, the PDO has warm and cool phases, as defined by patterns of ocean temperatures in the northeast and tropical Pacific Ocean (Figure).
FIGURE The Pacific Decadal Oscillation. (Top) Typical winter patterns of sea surface temperature (colors), sea-level pressure (contours), and surface wind stress (arrows) during positive (warm) and negative (cool) phases of PDO. Temperature anomalies are in degrees Celsius. (Bottom) History of the PDO index (the principal component of monthly sea surface temperature anomalies in the North Pacific Ocean poleward of 20°N) from 1900 to 2010. SOURCE: Figure obtained with permission granted by Nate Mantua at the University of Washington’s Joint Institute for the Study of Atmosphere and Ocean.
FIGURE 4.1 Trend of thermosteric sea level (mm yr-1) for 1993–2009 (left) and 1961–2008 (right), based on an updated version of data from Ishii and Kimoto (2009). SOURCE: Courtesy of Masayoshi Ishii, Japan Meteorological Research Institute.
in sea level of about 0–0.8 mm yr-1 from 1955 to 2003, rather than a fall (Bindoff et al., 2007). This difference suggests that the spatial pattern of sea level varies on decadal and longer timescales.
Changes in wind-driven ocean circulation can play an important role in determining patterns of sea-level change in the northeast Pacific Ocean on seasonal to decadal and longer timescales (e.g., Timmermann et al., 2010; Bromirski et al., 2011; Merrifield, 2011; Sturges and Douglas, 2011). Recent studies show a decrease in the rate of sea-level rise along the west coast of the United States since 1993, which is consistent with IPCC (2007) findings, but no statistically significant trends appear in tide gage records (Bromirski et al., 2011), satellite altimetry data, or in situ temperature observations since 1980. For example, thermosteric sea-level calculations show falling sea level off the U.S. west coast from 1993 to 2009 (Figure 4.1, left) and rising sea level from 1961 to 2008 (Figure 4.1, right). Bromirski et al. (2011) suggested that the flat sea-level trend since 1980 and the decrease since 1993 are associated with PDO phase changes.
Seasonal and Interannual Variability
Among all the climate modes, ENSO is the dominant cause of sea-level variability in the northeast Pacific Ocean on interannual timescales (e.g., Zervas, 2009; Bromirski et al., 2011). Sea level rises off the west coast of the United States during El Niño events and falls during La Niña events. El Niños differ in magnitude and large-scale form (Barnard et al., 2011) but commonly produce an active winter storm season in the northeast Pacific. The associated winds and ocean circulation changes may elevate sea level by 10–30 cm for several months along the west coast (Chelton and Davis, 1982; Flick, 1998; Bromirski et al., 2003; Allan and Komar, 2006; Komar et al., 2011). In fact, the highest sea levels recorded along the west coast were usually associated with El Niño events (e.g., Figure 4.2). For example, on January 27, 1983, during one of the largest El Niños in half a century, seven tide gages along the west coast (San Diego, Los Angeles, Monterey, Crescent City, Charleston, Astoria, and Seattle) recorded their highest water levels.1 Peak sea level was 24 cm above predicted in San Diego (104 years of record), 31 cm above predicted in Los Angeles (87 years of record), and 76 cm above predicted in Seattle (112 years of record).
Large El Niño and La Niña events also can be seen in satellite altimetry data. The top panels of Figure 4.3 show the sea-level rise observed during the El Niño of 1997–1998 and the sea-level fall observed during the 1999 La Niña. The ENSO signal is strongly seasonal and reaches a peak amplitude in the Northern Hemisphere winter. Figure 4.3c shows the ocean seasonal cycle, which is occasionally magnified by ENSO.
FIGURE 4.2 San Francisco tide gage record showing relative sea-level increases during major El Niño events. SOURCE: Tide gage data from the Permanent Service for Mean Sea Level.
FIGURE 4.3 (a) Sea-level anomaly (SLA), the difference between mean sea level for 1993–2009 and sea level during the December 1997 El Niño. (b) Same as (a) but for a La Niña event in February 1999. Color scale on right is in cm. (c) Time series of monthly SLA offshore San Diego, San Francisco, and Seattle. The two black arrows correspond to the dates shown in the upper figures. SOURCE: AVISO satellite altimetry data from <http://www.aviso.oceanobs.com/>.
Decadal and Longer Variability
The low-frequency (decadal and longer) variability in sea level off the U.S. west coast often corresponds to forcing by regional and basin-scale winds associated with climate patterns such as the PDO and the North Pacific Gyre Mode (e.g., Lagerloef, 1995; Fu and Qiu, 2002; Jevrejeva et al., 2006; Cummins and Freeland, 2007; Miller and Douglas, 2007; Di Lorenzo et al., 2008, 2010; Bromirski et al., 2011; Sturges and Douglas, 2011; Merrifield, 2011). For example, ocean modeling by Bromirski et al. (2011) found that surface heating alone produced falling sea level—the opposite to that observed—whereas forcing by winds explained the rise in sea level along the U.S. west coast since 1950. They suggest that the lack of a significant trend in sea level observed in tide gages since 1980 reflects forcing by winds associated with phase changes of the PDO. Sea level rose when the PDO changed from negative (cool) to positive (warm) around 1976–1977, and it fell when the PDO changed from positive to negative at the end of the 1990s (see lower figure in Box 4.1). The PDO has largely been in a positive phase since 1977, although negative phases have occurred almost a half-a-dozen times since the 1990s.
ENSO may also play a significant role in decadal and longer sea-level variability (Newman et al., 2003). Indeed, ENSO and the PDO are not independent. ENSO can influence the PDO (Newman et al., 2003; Schneider and Cornuelle, 2005), and the PDO can modulate tropical Pacific circulation and ENSO (e.g., Vimont et al., 2009; Alexander et al., 2010).
The spatial variability of sea level in the Pacific Ocean is driven primarily by ENSO, which affects sea level on seasonal to decadal timescales, and is also associated with phase changes in the PDO, which affects sea level on decadal and longer timescales. Satellite altimetry, tide gage, and ocean temperature measurements all indicate a long-term increase in sea level off the U.S. west coast, with large amplitude seasonal to multidecadal variability. The measurements show no statistically significant sea-level trend since 1980, consistent with the PDO phase changes.
Any climate-induced increase in storm frequency and magnitude will induce short-term changes in sea level. This issue is critical to coastal planners because storm surges and wind-driven waves are responsible for most of the flooding and erosion damage along the west coast of the United States (Armstrong and Flick, 1989; Domurat and Shak, 1989; Allan and Komar, 2006). The most severe coastal impacts tend to occur when a storm surge coincides with high tides and/or during periods of anomalously high sea level, such as those caused by El Niños. For example, the simultaneous occurrence of anomalously high sea level, high waves in late January and early March, and high astronomical tides caused significant damage along the California coast during the El Niño winter of 1983 (Figure 4.4). The amplitude of local sea-level rise from storm and wave events can greatly exceed the projected amplitude of global and regional sea-level rise, even beyond 2100, so understanding their additive effects is crucial for coastal planning. This section describes the contributions of these factors to short-term sea-level rise and the extent to which they may be changing with climate change (Task 2b).
Contributions of Tides, Storms, and El Niños to Local Sea Level
High tides along the U.S. west coast occur twice daily, often of uneven amplitude, caused predominately by the gravitational attraction of the Moon and the Sun on the Earth. The Earth-Moon-Sun orbital geometry also results in heightened high tides twice monthly (spring tides, near the times of the full and new moon) and every 4.4 years and 18.6 years (Zetler and Flick, 1985). The largest tidal amplitudes of the year along the coasts of California, Oregon, and Washington usually occur in winter and in summer (Zetler and Flick 1985). Tides in the highest winter and summer months are often more than 20 cm higher than tides in the spring and fall months.2 The peaks in the 4.4-year and 18.6-year cycles produce monthly high tides that are about 15 cm and 8 cm, respectively, higher than they are in the intervening years (Flick, 2000). Flick et al.
FIGURE 4.4 (a) Hourly sea-level pressure (SLP; mb), (b) sea-level anomaly (cm) above tide-predicted levels, (c) predicted and (d) observed sea level (cm) relative to a mean sea-level datum, and (e) significant wave height (Hs, the average height of the highest one-third of waves [m]) from a buoy sensor near San Francisco during the El Niño winter of 1983. SOURCE: Adapted from Flick (1998).
(2003) reported increases in the range from high to low astronomical tide over multiple decades at some, but not all, U.S. west coast tide gages.
Storm surges are created when high winds, the Coriolis force, and low barometric pressures from coastal storms force sea water onto the shore. During the most severe winter storms, surface atmospheric pressure along the west coast drops by 20 mb or more from long-term average levels, typically with greater pressure drops in Washington and Oregon than in California. The drop in atmospheric surface pressure raises sea level by approximately 1 cm for every 1 mb decrease in atmospheric pressure. The resulting increase in sea level is usually regional, according to the regional scale of winter cyclones, and typically lasts only a few days at most (Flick, 1998). Woodworth and
Blackman (2004) investigated high-water levels from tide gages around the world since 1975 and found that the magnitude of sea-level extremes has risen in many locations, including some parts of the U.S. west coast, and that these extremes closely followed increases in the median sea level.
Strong ocean winds also produce surface gravity or wind waves. The most extreme such waves are of two types: sustained intervals of large waves (measured by the significant wave height, the average height of the largest one-third of the waves) and rogue waves, which have individual crests that are much larger than the significant wave height. Sustained intervals of large waves occur during strong storms. These storm waves can propagate over a long distance to the shoreline. Rogue waves are produced by interactions among waves and perhaps currents, and they have the greatest impact when they arise during a sustained interval of large waves. By definition, they are expected but relatively uncommon events (Baschek and Imai, 2011).
El Niños can significantly elevate sea level along the west coast during winter months (see “Changes in Ocean Circulation” above), especially along the California coast because the North Pacific storm track is displaced toward the equator during El Niño events (Seager et al., 2010). The wind and pressure patterns that elevate sea level above climatological normals along the west coast also may occur in winters when El Niño is not present. Winters with high sea-level anomalies have usually had a few large North Pacific storms with strong westerly, southwesterly, or northwesterly winds offshore, which generate storm surges and high waves along the coast of California and sometimes the coasts of Oregon and Washington.
The path and propagation speed of storms controls the wind direction and barometric pressure, which, in turn, affects the generation of wind waves and high water (e.g., O’Reilly and Guza, 1991). The highest winds, and hence waves, along the west coast of the United States nearly always occur during strong winter extra-tropical cyclones (Wang and Swail, 2001; Bromirski et al., 2003; Caires et al., 2004; Ruggiero et al., 2010; Barnard et al., 2011; Seymour, 2011). Tropical cyclones rarely travel as far north as California, although two cases have been recorded historically (Hurd, 1939; Chenoweth and Landsea, 2004). Significant wave heights recorded by offshore coastal buoys during extra-tropical events can exceed 10 m (Figure 4.5; Ruggiero et al., 2010; Seymour, 2011), although they are usually smaller as they approach the shoreline. Significant wave heights at the shoreline vary considerably depending on incident wave direction and nearshore bathymetry.
Wave swells generated by storms propagate long distances (e.g., from the central North Pacific to the U.S. west coast) over several days. Swells generated far from the west coast tend to peak at relatively long periods (12 seconds or more), whereas more locally generated wave swells tend to peak at periods of 10 seconds or less. The largest swells are generated by winter cyclones that produce high winds with a long fetch (the total distance that wind blows over the sea surface during the storm) directed toward the west coast. A broad, deep low-pressure system over the North Pacific favors these conditions (Figure 4.6; Bromirski et al., 2005). Synoptic timescale patterns like this tend to occur during El Niño winters, but not exclusively (Seymour et al., 1984; Bromirski et al., 2005; Allan and Komar, 2006). Larger than normal waves have occurred during El Niño winters along the California coast and some parts of the Oregon and Washington coasts (Bromirski et al., 2005; Allan and Komar, 2006). La Niñas have been shown to produce smaller than normal winter wave heights at some California locations, but not everywhere along the west coast (Allan and Komar, 2006). Overall, the occurrence of large storms and high waves is clustered in time, with particular years and groups of years having many large storms, and other years having few or no large storms.
Peaks in wind waves are generally much higher than sea-level anomalies (Seymour et al., 1984; Seymour, 1998; Storlazzi and Griggs, 1998; Ruggiero et al., 2010). High breakers induce a change in mean water level at the beach (set-up), which can be about 20 percent of the breaking wave height (Dean and Dalrymple, 1991). High wave events sometimes, but not always, coincide with high sea levels (Cayan et al., 2008; Ruggiero et al., 2010).
Changes in Storminess and Extreme Wave Heights
Evidence of changes in storminess (wind intensity) in the North Pacific Ocean is mixed. Bromirski et al. (2003) examined nontidal sea-level fluctuations from 1858 to 2000 in the San Francisco tide gage record
FIGURE 4.5 (a) Number of storm events per month off Oregon and Washington between 1976 and 2007, when the significant wave height (SWH) exceeded a threshold of 8.1 m at two deep-water wave buoys. (b) Days when the threshold of 8.1 m was exceeded (dots), annual maxima (circles), and the five largest storms per year (asterisks) for 1976–2007, illustrating the seasonality of the extreme wave climate. The 100-year significant wave height is shown by the solid horizontal line and its associated uncertainty is the dashed horizontal lines. SOURCE: Ruggiero et al. (2010).
FIGURE 4.6 Atmospheric circulation during periods of high waves along the central California coast exhibits broad-scale low pressure over the North Pacific. This map shows anomalies of 700 hPa height in meters during the 15 winter months (November through March) from 1981 to 2003 when wave energy offshore San Francisco was greatest. The region of anomalously low 700 hPa indicates a low-pressure trough and increased storminess in the central and eastern North Pacific. Significant negative and positive anomalies are blue and red, respectively. SOURCE: Adapted from Bromirski et al. (2005).
and found significant decadal variability. Although the record showed an increase in storminess from 1950 to 2000, the storm intensity in recent decades did not significantly exceed that in the decades prior to 1950 (Bromirski et al., 2003). On the other hand, the IPCC Fourth Assessment Report cited several studies that reported increases in the strength of the winter westerly wind circulation across the North Pacific during the past few decades (Trenberth et al., 2007).
Lowe et al. (2010) described climate change effects on storm intensity as inconclusive, with no consensus among different model simulations on local changes in storm frequency. A simulation of San Francisco sea-level anomalies forced by 21st century climate change simulations (Cayan et al., 2008) found considerable interannual and decadal variability, driven partly by storm characteristics, superimposed on an assumed long-term rise in mean sea level. Several climate models discussed in the IPCC Fourth Assessment Report project that the mid-latitude storm tracks in both the southern and northern hemispheres will migrate poleward over the 21st century (Meehl et al., 2007). A subsequent projection by Salathé (2006) also showed a northward shift in the North Pacific winter storm track over the next several decades. The storm tracks and Pacific wind fields in some global climate model projections suggest that future wave heights might diminish somewhat over the open ocean and along the coast from southern and central California to Oregon (Salathé, 2006; Cayan et al., 2009).
If frequency or intensity of storminess changes as a result of climate change, the frequency of high sea-level extremes also would likely change. Even if the storminess regime does not change, sea-level rise will increase the exposure of the coast to storm-driven surge and high waves, magnifying their impact on the coast.
Analyses of marine weather reports discussed in the IPCC Fourth Assessment Report showed an increase in significant wave height of 8–10 cm per decade over the central and eastern North Pacific from 1950 to 2002 (Trenberth et al., 2007). Gulev and Grigorieva (2006) attributed these increases to longer period, longer distance sources of swell as well as to more locally generated wind waves. The tendency for an increase in wave energy over the eastern North Pacific is also indicated by wave hindcasts (Graham and Diaz, 2001), buoy observations (e.g., Allan and Komar, 2006), some wave buoy records (Ruggiero et al., 2010), and satellite altimeter observations (Young et al., 2011a).
A study of North Pacific wind variability on 2- to 10-day timescales from the National Centers for Environmental Prediction (NCEP) Reanalysis (Kalnay et al., 1996) indicated that wind speed trends are variable, owing to the occurrence of relatively infrequent large events. From the 1950s through the 1990s, wave model reanalyses over the North Pacific (Graham and Diaz, 2001; Caires et al., 2004) indicate a trend toward increasing wave height. From a series of buoy observations beginning in the late 1970s, Storlazzi and Wingfield (2005), Allan and Komar (2006), Ruggiero et al. (2010), and Seymour (2011) found that the largest waves along the coast from California to Washington state were larger in the period after 1990 than in the period before (Figure 4.7). This change was associated with a deepening of the winter low pressure system over the North Pacific Basin and partly to the incidence of some relatively strong El Niño years since 1995.
Increases in wind speed and wave heights in the northeastern Pacific Ocean have been reported recently,
FIGURE 4.7 Increases in the annual maximum wave height (green; m), average of the five largest wave events per year (blue), winter average height (red), and annual average height (black) from northeast Pacific wave buoy sensors. Open circles represent years with too much missing data (i.e., winter months missing more than 60 percent of data). SOURCE: Ruggiero et al. (2010), after Allan and Komar (2006).
but the interpretation of these changes is controversial. Analyses of global ocean winds from ship observations (Tokinaga and Xie, 2011), satellite microwave sensors (Wentz et al., 2007), and satellite altimeters (Young et al., 2011a) indicate that wind speeds have risen over the global oceans, although the trends found by Young et al. (2011a) are greater than those derived from Tokinaga and Xie (2011) and Wentz et al. (2007) by approximately a factor of two (Wentz and Ricciardulli, 2011; Young et al., 2011b). The Young et al. (2011a) analysis also found that wind speeds within the highest 1 percent of events have risen over much of the extra-tropical oceans over the past two decades, including an increase of about 1 percent per year in the northeast Pacific, and that this increase is accompanied by increases in the extreme wave heights. The latter occurs in particular in the northeast Pacific Ocean, which is consistent with increasing extreme wave heights (by as much as 2 m over the record period) during big storms recorded in near coastal deep-water buoy records from northern California to Washington (Allan and Komar, 2006; Menéndez et al., 2008; Ruggiero et al., 2010). However, further analysis by Gemmrich et al. (2011) suggests that much of this change is spurious, caused by changes in buoy hardware and data processing. All of these estimates were made from records that are only a few decades long, and thus partly reflect changes in wind forcing associated with natural climate variability such as the Pacific Decadal Oscillation and other interannual-interdecadal fluctuations. However, the global extra-tropical pattern of extreme wave increase found by Young et al. (2011a) is atypically widespread for most decadal natural variability, and thus might indicate a longer trend. As yet there is no good explanation for why such a trend would occur.
Periods of anomalously high sea levels and wave heights along the west coast of the United States exhibit considerable variability on synoptic, interannual, and decadal timescales, in association with ENSO and other climate patterns. Some evidence suggests that wave heights have increased along the west coast from northern California to Washington during the past few decades. However, it is likely that much of this increase is associated with interannual- to decadal-scale natural variability of the Pacific atmosphere-ocean system. Some global climate models predict that the North Pacific storm track will shift northward as global climate warms during the next several decades, which would generate extreme wave heights and storm surges along the Oregon and Washington coasts. However, a northward shift in the North Pacific storm track has not yet been confirmed.
All climate models project ample winter storm activity in the North Pacific in future decades, suggesting that periods of anomalously high sea level and high waves will continue to occur along the west coast. Storm-generated bursts of high sea levels and waves are expected to vary from year to year and decade to decade. Over the next few decades, these anomalies will likely eclipse the secular rise in sea level (few to several mm per year). Short-period fluctuations of sea level may sometimes exceed 20 cm, and storm-driven wave heights of 1 m or even higher amplitudes than are seen in the historical record could easily occur. These variations will have greatest impact when they occur on days with high tides.
As glaciers and ice sheets melt and lose mass and the melt water is transferred from the continents to the ocean, the solid earth deforms and the gravitational field of the planet is perturbed. The addition of new water to the ocean basins and the associated gravitational and deformational effects create regional patterns of sea level change. Both modern melting and deglaciation of the ancient ice sheets affect sea-level change along the west coast of the United States. Melting of the ancient ice sheets caused the solid earth to rebound (glacial isostatic adjustment), resulting in significant vertical land motions in the vicinity of the California, Oregon, and Washington coasts. In contrast, modern melting affects land motions at the ice masses, which are far from the U.S. west coast, but the gravitational effect influences the height of the sea surface in the northeast Pacific Ocean. This section describes the effects of modern land ice melt on sea-level rise off the coasts of California, Oregon, and Washington. The effects of ancient ice melt are discussed in the following section (see “Glacial Isostatic Adjustment” below).
Modern melting of land ice affects sea level along the west coast of the United States in two ways. First, the large mass of glaciers and ice sheets generates an additional gravitational pull that draws ocean water closer, raising relative sea level near the ice masses. As the ice melts, the amount of ice mass on land declines, decreasing its gravitational pull on the ocean water. The loss of mass also results in uplift of the land mass under the ice. The combination of these effects causes relative sea level to fall in the vicinity of the ice mass. The fall extends, at decreasing rates, in the region within a few thousand km of the melting ice. Second, ice melt enters the ocean, raising global mean sea level. Because of gravitational and deformational effects, however, the distribution of new ice melt is nonuniform over the globe. Relative sea level falls near the shrinking ice mass and rises everywhere else. This effect is shown schematically in Figure 4.8. The combined effect of new water mass entering the ocean and altered gravitational attraction results in a spatial pattern of sea-level rise that is unique for each ice sheet or glacier (Mitrovica et al., 2001; Tamisiea et al., 2003). As a consequence, these sea-surface geometries have come to be known as sea-level fingerprints.
Only a few studies have attempted to map the sea-level fingerprints of melting land ice along the west coast of the United States (e.g., Tamisiea et al., 2003, 2005). Figure 4.9A shows the sea-level fingerprints of the three largest sources of land ice that are most likely to have significant effects on west coast sea level: Alaska, Greenland, and Antarctica. The figure shows that melting of Alaska glaciers creates a strong north-south gradient in relative sea-level change along the west coast. The gradient from uniform melting of the Greenland Ice Sheet is much smaller (Figure 4.9B). Uniform melting of either the Antarctic Ice Sheet or the West Antarctic Ice Sheet leads to a uniform change in relative sea level along the entire west coast (Figure 4.9C).
To estimate the effect of fingerprinting from these three ice masses on relative sea level, it is necessary only to multiply the global sea-level equivalent of the mass loss from each source by the appropriate scale factor (colored contours) indicated in the figure and then add the contributions from all three sources. Scale factors greater than 0 indicate that the sea-level fingerprint increases relative sea-level rise at that location, and scale factors greater than 1 indicate that the rise is higher than the global sea-level equivalent value. Scale factors less than 0 mean that the effect of mass loss from a source causes the relative sea level to fall. Scale factors for other ice sources (e.g., European Alps, northeastern Canadian Arctic, Patagonia) are not available at the resolution shown in Figure 4.9, but these sources are likely too small and/or too distant to affect the gradient in sea-level change along the U.S. west coast.
The scale factors and ice loss rates used to calculate the adjusted rates of relative sea-level rise are given in Table 4.1. Modeling or estimating individual regional land ice losses is beyond the scope of this study, so
FIGURE 4.8 Schematic view of the changing sea level caused by a shrinking land ice mass. Relative sea level at time t1 exceeds the mean sea level near the ice mass and is less than the mean at some distance beyond the mass. As the land ice mass decreases (time t2), the local gravitational attraction decreases and the land in the vicinity of the ice rises, causing the relative sea level to fall, even though the mean sea level increases. SOURCE: Adapted from Tamisiea et al. (2003).
FIGURE 4.9 Sea-level responses in the northeast Pacific to ice loss from three major ice masses. The responses are shown as scale factors, which are the local sea-level equivalent divided by the global mean sea-level equivalent. (A) Response to melt from the Alaskan glacier system, as modeled in Tamisiea et al. (2003). (B) Response to uniform melting over the entire grounded portion of the Greenland Ice Sheet. (C) Response to melting across the entire Antarctic Ice Sheet (left) or the West Antarctic Ice Sheet (right). All of the calculations underlying this figure treat the Earth as elastic; that is, the timescale of response is assumed to be sufficiently rapid that viscous effects can be neglected. SOURCE: Courtesy of Jerry Mitrovica and Natayla Gomex, Harvard, based on calculations described in Mitrovica et al. (2011).
the committee used ice loss rates averaged from data reported in Appendix C. To simplify the analysis, scale factors were picked from Figure 4.9 for three representative locations along the U.S. west coast: the north coast (approximately Neah Bay, Washington), the central coast (approximately Eureka, California), and the south coast (approximately Santa Barbara, California).
In the absence of a sea-level fingerprint effect, the expected sea-level rise along the U.S. west coast from ice loss in Alaska, Greenland, and Antarctica would be 0.79 mm yr-1, the sum of the ice mass loss rates in Table 4.1. The overall effect of the fingerprint is to lower sea-level rise along the entire west coast. Although melting of Alaska glaciers contributes less water to the oceans than melting of the Greenland Ice Sheet, the Alaska glaciers are closer to the U.S. west coast and have a greater effect on relative sea level in the region. The adjusted rates of relative sea-level rise for the three sources (found by multiplying the loss rate by the fingerprint scale factors) are 0.46 mm yr-1 for the north coast, 0.60 mm yr-1 for the central coast, and 0.68 mm yr-1 for the south coast. This adjusted rate of sea-level rise is 42 percent lower than the rate for melting of the three ice sources (0.79 mm yr-1) for the north coast, 24 percent lower for the central coast, and 14 percent lower for the south coast.
This simple calculation provides only an approximate estimate of the magnitude and sign of relative sea-level change due to gravitational and deformational effects of modern land ice melting. Uncertainties in the rate of ice loss and, to a lesser extent, the neglect of fingerprints of other sources of land ice can lead to significant uncertainties in the adjusted rates of relative sea-level rise. In particular, the steep gradient caused by Alaska’s proximity to the study region, combined with the high uncertainty in the rate of ice loss from Alaska compared to the ice sheets, yield a wide range of possible adjustments to relative sea-level rise (see Appendix C). When the uncertainty in loss rates from the three sources is considered, the adjusted rate of relative sea-level rise due to melting of these ice masses ranges from 0.1–0.9 mm yr-1 for the north coast, 0.1–1.1 mm yr-1 for the central coast, and 0.1–1.3 mm yr-1 for the south coast.
The large mass of glaciers and ice sheets creates a gravitational pull that draws ocean water closer. As the ice melts, the gravitational pull decreases, ice melt enters the ocean, and the land and ocean basins deform as a result of this loss of land ice mass. These gravitational and deformational effects produce a spatial pattern of regional sea-level change commonly referred to as a sea-level fingerprint. The land ice masses that most affect sea level along the California, Oregon, and Washington coasts are in Alaska, which is nearby, and Greenland and Antarctica, which are large. Melting in Alaska and, to a lesser extent, Greenland, causes relative sea level to fall at decreasing rates from northern Washington to southern California. Melting in Antarctica causes a uniform sea-level rise along the entire west coast of the United States. The net result is a reduction in the contribution of Alaska, Greenland, and Antarctica ice melt to relative sea-level rise off Washington, Oregon, and California. The magnitude of this reduction decreases from about 42 percent along the north coast (Neah Bay) to 24 percent along the central coast (Eureka) to 14 percent along the south coast (Santa Barbara) for 1992–2008.
TABLE 4.1 Ice Loss Rates, Fingerprint Scale Factors, and Adjusted Rates of Relative Sea-Level Rise for Three West Coast Locations
|North Coast||Central Coast||South Coast|
|Ice Source||Ice Mass Loss Rate(mm yr-1 SLE)a||Scale Factor||Adjusted Sea-Level Rise (mm yr-1)||Scale Factor||Adjusted Sea-Level Rise (mm yr-1)||Scale Factor||Adjusted Sea-Level Rise (mm yr-1)|
a Based on the average of published rates for 1992-2009 for Greenland and Antarctica and 1992-2008 for Alaska, as described in Appendix C.
b Average of east and west Antarctic values.
Vertical land movements that affect relative sea level may be caused by geologic processes (e.g., glacial isostatic adjustment, tectonics, compaction) or anthropogenic activities (e.g., groundwater or oil extraction). Each of these processes can in principle be modeled or observed (Box 4.2), although data coverage is sparse and uncertainties are large. The estimated rates of vertical land motions resulting from these processes are on the order of several mm yr-1—about the same as the rate of global sea-level rise—with magnitudes that vary over spatial scales ranging from one to thousands of km.
Glacial Isostatic Adjustment
The solid earth and oceans continue to respond to the decay of ice sheets since the last deglaciation through glacial isostatic adjustment (GIA). The loss of ice mass produces uplift in regions under the former ice masses, including northern Washington, and subsidence in areas at the ice margin and beyond, including the rest of Washington, Oregon, and California (Box 1.2). In addition, the transfer of melt water to the oceans and the consequent subsidence of the ocean basins in response to the increased water load produce a change in the absolute sea level (or geoid, an equipotential surface of the Earth’s gravity field that coincides with the mean sea surface). Both processes are manifested in geological records of relative sea-level rise, geodetic observations, and GIA models.
GIA models commonly focus on predictions of sea level because many of the time series used as constraints are from paleo sea-level data (e.g., Engelhart et al., 2011). The sea-level predictions in GIA literature (e.g., Peltier, 2004) are typically a measure of relative sea-level change, according to the following equation:
Relative sea-level change =
Change in absolute sea level -
Change in height of the solid earth surface,
where the changes in absolute sea level and the height of the solid earth surface are measured relative to a common datum (e.g., the Earth’s center of mass). In GIA models, the solution is obtained using the sea-level equation (Farrell and Clark, 1976; Peltier, 1976;
Geodetic Observations of Vertical Land Motion
Ground-based and space-based geodetic techniques are used to observe vertical land motion at sub-cm vertical precision. Leveling measures the vertical component of land motions from ground stations spaced hundreds to thousands of meters apart. Height differences between points are measured by setting a level on a tripod and orienting it so that the line of sight is horizontal. For short distances between benchmarks (e.g., 1 km, similar to the spacing used for tide gage leveling), a vertical accuracy of about 1 mm can be achieved (see Appendix D). For longer lines (e.g., 10 km), such as are used for tectonic studies, expected accuracies are about 2 mm.
Satellite-based systems, including the Global Navigation Satellite System (GNSS) and Interferometric Synthetic Aperture Radar (InSAR), have been available for selected regions since the 1990s. The GNSS comprises constellations of navigation satellite systems, including 24 Global Positioning System (GPS) satellites, circling the Earth with accurately determined orbits and broadcasting their precise locations and velocities. The global network of GNSS stations, along with other space geodetic techniques (e.g., satellite laser ranging, very long baseline interferometry, Doppler Orbitography and Radiopositioning Integrated by Satellite [DORIS]), provide the fundamental reference frame that makes accurate positioning and time transfer possible. The radio signals sent by the satellites are received at fixed ground stations or low-orbiting satellites. Because errors in the vertical component are typically twice as large as errors in the horizontal components, only continuous GPS (CGPS) stations are used routinely to measure vertical land motion. The National Science Foundation’s Plate Boundary Observatory significantly increased the number of CGPS stations in the western United States. Stations along the west coast are spaced ~25–50 km apart. If time series are long (> 5 years) and the location of the station is accurately known, vertical resolution can reach ~1–2 mm yr-1.
InSAR uses phase differences between radar images from repeat satellite passes to infer changes in the round-trip travel time of the radar signals between the earth surface and the satellite. These changes can be used to generate interferograms to infer line-of-sight surface deformation. The high-resolution image swath size is 60–100 km, and the spatial resolution of the measurement tends to be on the order of 40–100 m (pixel size). The vertical resolution of the measured surface change is less than 1 cm.
Clark et al., 1978), which assumes that the volume of water in the earth system must be conserved. Note that the GIA literature that uses the sea-level equation frequently interchanges the terms “absolute sea level,” “sea surface height,” and “geoid,” which creates a problem when comparing and discussing predicted GIA contributions to altimetry and GRACE observations (see the discussion in Tamisiea, 2011). “Height change,” “radial displacement,” and “vertical motion” also are used interchangeably.
The committee compared the GIA predictions of relative sea-level change at 21 tide gage locations in California, Oregon, and Washington from an ensemble of 16 models (Figure 4.10). The time period of these models is ± 250 years relative to the present day. The models differ significantly from one another, depending on the earth rheology parameters and deglaciation model used (Table 4.2). Most GIA models employ a self-gravitating, spherically symmetric Earth model, with Maxwell rheology. Some use laterally varying viscosity and mantle thicknesses (e.g., Wu, 2006).
The new ICE-5G reconstruction of the surface topography and land-ice distribution at the last glacial maximum differs significantly from its ICE-4G precursor at all Northern Hemisphere locations that were glaciated (see Peltier, 2002a; Tarasov and Peltier, 2002). ICE-5G and ICE-6G (Peltier, 2010) contain a similar mass for the Laurentide Ice Sheet and cover the same surface area of the North American continent. They differ in the relative thickness of the ice sheet, which in the case of ICE-6G has been adjusted to eliminate the misfits between the vertical motion predictions of the model and the GPS observation analyses by Argus and Peltier (2010). Changes include a thickening of the ice cover over Labrador, Yellowknife, and the northwestern border between British Columbia and Alberta, as well as a thinning of the ice cover to the south of Hudson Bay.
All the GIA models shown in Figure 4.10 predict a similar pattern of variability in relative sea-level change along the Pacific coast, rising from 32° latitude to a maximum around 47° latitude, and then declining
FIGURE 4.10 Ensemble of 16 GIA models showing the predictions of relative sea-level rise (expressed as change in absolute sea level minus change in height of the solid earth surface) at the latitudes of 21 tide gages off the California, Oregon, and Washington coasts. The time period of these models is ± 250 years relative to the present day. SOURCES: ICE4GVM2 (Peltier, 1998) and ICE-5GVM2 (Peltier, 2004) models and their variations are from <http://www.sbl.statkart.no/projects/pgs/authors>. Other GIA models (Wang and Wu, 2006; Paulson et al., 2007; van der Wal et al., 2009; Sasgen et al., 2012; H. Wang, personal communication) were provided by the respective authors. Predicted values from ICE5G and ICE6G models and their variations were computed for this study by Richard Peltier, University of Toronto.
TABLE 4.2 Earth Rheology Parameters Used in Selected GIA Models
|Viscosity (×10-21Pa s)|
|G1A Model||Ice History Model||Lithosphere Thickness (km)||Upper Mantle||Lower Mantle|
|ICE4GVM2 (Peltier, 2002a)||ICE-4G||90||04~1.5||1.3~3.9|
|ICESGVMa (Peltier, 2004)||ICE-5G||90||0.4~1.5||1.3~3.9|
|ICE5GVM4a (Peltier, 2004)||ICE-5G||90||0.4~0.9||0.9~3.9|
|Paulson-Zhong-Wahra (Paulson et al., 2007)||ICE-5G||98||09||3.6|
|Sasgen-Klemann-Martineca (Sasgen et al, 2012)||HUY, NAWIb||100||0.52||5.9|
|van der Wala (van der Wal et al., 2009)||ICE-5G||98||0.9||3.6|
|Wang Wu ICE4G (Wing and Wu, 2006)||ICE-4G||115||0.6||LM1 = 3, LM2 = 6, β = 0.4c|
|Wing Wu ICE5G (H. Wang, personal communication)||ICE-5G||115||0.6||LM1 = 3, LM2 = 6, β = 0.4c|
a Models that considered rotational feedback.
b HUY is the Antarctica ice model (Huybrechts, 2002), scaled to 12 m of sea-level rise since the last glacial maximum. NAWI is the Northern Hemisphere ice model (Huybrechts, 2002).
c Laterally varying mantle viscosity. LM1 is a shallow lower mantle, and LM2 is a deep lower mantle. Lateral variation is inferred from lateral shear wave velocity anomalies given in the seismic tomographic model S20A with a scaling factor β.
sharply. The strong latitudinal gradient in Washington illustrates the importance of glacial isostatic adjustment in regions under or at the margins of the extinct Laurentide Ice Sheet. In Cascadia, uplift is expected at the far north locations, which had been covered by the ice sheet, and subsidence is expected at the other locations, which are along the former margins of the ice sheet. In Oregon and California, the variance among models is almost as large as any apparent trend. The mean relative sea-level rise from the GIA model ensemble at each tide gage location is given in Table 4.3.
It should be noted that some studies suggest that the global earth rheology parameters (e.g., mantle viscosity) used to study the GIA process may not be suitable for subduction zones such as Cascadia. For example, James et al. (2000) used a regional, rather than global, deglaciation history to analyze GIA in southern Vancouver Island. Local paleo sea-level data show rapid uplift 12,000 years before present, which best fits a mantle with much lower mantle viscosity (~1019 Pa s) than is used in the GIA models shown in Table 4.3. James et al. (2000) extrapolated these results, concluding that vertical land motion from glacial isostatic adjustment along the Cascadia Subduction Zone is negligible compared to the influence of tectonics.
The U.S. west coast is characterized by two tectonically distinct regions: (1) the Cascadia Subduction Zone, where lithospheric plates are colliding north of Cape Mendocino, California, and (2) the San Andreas Fault Zone, where the plates are sliding past one another south of Cape Mendocino (Figure 1.8). Vertical land motions in both regions are caused by strain buildup along faults and release during an earthquake. Vertical land motions associated with the Cascadia Subduction Zone (e.g., Hyndman and Wang, 1993) are generally larger than those associated with the San Andreas Fault Zone (Argus and Gordon, 2001). Surficial crustal motions along the San Andreas Fault Zone are primarily horizontal, although convergence in some areas can produce locally significant rates of vertical deformation (e.g., Argus et al., 1999; Argus and Gordon, 2001). South of the San Francisco Bay area, the principal fault trace extends inland as much as 50–100 km, further reducing its effect on coastal vertical land motion.
The history of crustal strain accumulation and release above subduction zone faults over hundreds of years is described by the earthquake deformation cycle (Nelson et al., 1996; Satake and Atwater, 2007). During an earthquake (known as the coseismic period), vertical land motion can change almost instantly by more than a meter (see “Rare Extreme Events” in Chapter 5). Between earthquakes (known as the interseismic period), rates of vertical land motion can be on the order of mm yr-1 and thus can have a significant impact on the relative sea level. Vertical land motions for the Cascadia Subduction Zone and San Andreas Fault Zone are described below.
TABLE 4.3 GIA Predicted Relative Sea-Level Rise for ± 250 Years Relative to the Present Day Using an Ensemble of 16 GIA Models at 21 West Coast Tide Gage Locations
|GIA Predicted Relative Sea-Level Rise (mm yr-1)|
|Cherry Point, WA||48.87||-122.75||-0.16||0.44|
|Friday Harbor, WA||48.55||-123.00||0.14||0.46|
|Neah Bay, WA||48.37||-124.62||0.58||0.64|
|Port Townsend, WA||48.12||-122.75||0.40||0.48|
|Toke Point, WA||46.72||-123.97||1.03||0.53|
|South Beach, OR||44.63||-124.05||1.00||0.34|
|Charleston II, OR||43.35||-124.32||0.86||0.32|
|Port Orford, OR||42.73||-124.50||0.81||0.32|
|Crescent City, CA||41.75||-124.20||0.67||0.31|
|N. Spit, Humboldt Bay, CA||40.77||-124.22||0.63||0.32|
|Point Reyes, CA||38.00||-122.98||0.53||0.30|
|San Francisco, CA||37.80||-122.47||0.47||0.29|
|Port San Luis, CA||35.17||-120.75||0.45||0.27|
|Santa Monica, CA||34.02||-118.50||0.34||0.25|
|Los Angeles, CA||33.72||-118.27||0.36||0.25|
|La Jolla, CA||32.87||-117.25||0.34||0.25|
|San Diego, CA||32.72||-117.17||0.35||0.25|
NOTE: Relative sea-level change is the change in absolute sea level minus the change in height of the solid earth surface. Relative sea-level rise has a negative sign compared to uplift of the earth surface due to GIA.
Cascadia Subduction Zone
Along much of the Oregon and Washington coasts, the earthquake cycle yields a characteristic pattern of vertical land movements (Figure 4.11). In the first stage of the cycle, slow interseismic strain accumulation over hundreds of years causes the upper plate to bend upward, leading to gradual uplift along the coasts above this part of the subduction zone. In areas closer to the plate boundary (usually the continental shelf) and further inland, the slow bending of the upper plate causes gradual subsidence. In the second stage of the cycle, the plate-boundary megathrust fault slips in a great earthquake, releasing hundreds of years of accumulated strain along many hundreds of kilometers of the plate boundary. During the earthquake, the former slow vertical deformation of the upper plate is reversed: coastal areas suddenly subside as much as 2 m and formerly subsiding areas landward and seaward are suddenly uplifted.
Current rates of interseismic vertical deformation can be estimated using dislocation models constrained by geodetic, thermal, and seismic data (e.g., Okada, 1985; Hyndman and Wang, 1993, 1995; Flück et al., 1997; Wang et al., 2003). To estimate interseismic deformation along the Washington and Oregon coasts, the committee used results from the CAS3D-2 model (He et al., 2003; Wang et al., 2003; Wang, 2007), a three-dimensional, viscoelastic, spherical earth, finite element model that assumes negligible present-day influence of GIA (following the work of James et al., 2000). The model has been further constrained by comparisons between geological estimates of coseismic subsidence of the 1700 earthquake and predictions from elastic dislocation models of slip on the Cascadia subduction zone (Leonard et al., 2004, 2010; Hawkes et al., 2011).
Table 4.4 shows the predicted rates of vertical land motion for the Cascadia Subduction Zone for 2010–2030 from the CAS3D-2 model assuming a continuation of the interseismic phase of the earthquake deformation cycle. The projections suggest that coastal sites, which are closest to the offshore subduction boundary, should be experiencing uplift, whereas more inland locations (Anacortes and Seattle) should
FIGURE 4.11 (Top) Deformation associated with a subduction-zone thrust fault on a coastline during an earthquake cycle. (Bottom) Idealized seismic cycle for a subduction zone, showing a long period of uplift, followed by small-scale subsidence and then a sudden drop in land elevation during a great earthquake. SOURCE: Modified from Horton and Sawai (2010).
TABLE 4.4 Vertical Land Motion Rates Predicted by the CAS3D-2 Model for 2010–2030
|Location||Latitude||Longitude||Rate of Vertical Land Motion (mm yr-1|
|Long Beach, WA||46.58||-123.83||1.87|
|Pacific City, OR||45.38||-123.94||1.69|
|Coos Bay, OR||43.36||-124.30||2.33|
SOURCE: Rates provided by Kelin Wang, Geological Survey of Canada, using the CAS3D-2 model (He et al., 2003; Wang, 2007). The model deformation history includes a coseismic rupture of the entire Cascadia subduction fault, representing the 1700 M 9 great earthquake, followed by locking of the fault, modeled using the conventional backslip approach (Savage, 1983). A mantle wedge viscosity of 10 Pa s was used, consistent with the results of postglacial rebound analyses at northern Cascadia and values adopted at other subduction zones.
be experiencing subsidence. Comparisons of the model projections with GPS data are discussed below (see “Current Rates of Vertical Land Motion Along the U.S. West Coast”). Model projections further forward in time are given in Chapter 5.
San Andreas Fault Zone
Unlike the Cascadia Subduction Zone, vertical land motions along the San Andreas Fault Zone cannot be characterized by a single tectonic model. The San Andreas Fault Zone comprises multiple sub-parallel
faults, each with limited extent and unique seismotectonic character. Although crustal displacement is primarily horizontal (Figure 4.12), local vertical motions result from rock uplift associated with restraining bends (e.g., Anderson, 1990) and active contractional processes associated with the Transverse ranges and the Ventura and Los Angeles basins (Namson and Davis, 1991; Donnellan et al., 1993; Yeats, 1993; Shaw and Suppe, 1994, 1996; Yeats and Huftile, 1995; Dong et al., 1998; Orme, 1998; Argus et al., 1999, 2005; Hager et al., 1999; Shaw and Shearer, 1999; Argus and Gordon, 2001; Bawden et al., 2001). A comprehensive analysis of tectonically induced vertical land motions for the San Andreas Fault Zone has not been done.
Compaction may rearrange the mineral matrix of sediment, reducing its volume (Kaye and Barghoorn, 1964; Allen, 2000; Brain et al., 2011). The amount of compaction depends on a number of factors, including the mechanical and chemical properties of the sediment (e.g., composition, porosity), the water content, and the loading history (Brain et al., 2011). For example, deposits with a high sand fraction undergo little compaction, whereas peat may compact as much as 90 percent by volume (Jelgersma, 1961).
Early studies of wetlands in North America (Kaye and Barghoorn, 1964) and Europe (Jelgersma, 1961) illustrated the importance of sediment compaction to relative sea-level rise. However, only a few studies have quantified compaction rates of coastal sediments. Törnqvist et al. (2008) analyzed wetland sediments from the Mississippi Delta and found compaction rates of 5 mm yr-1 on millennial timescales and more than 10 mm yr-1 in some areas on decadal to century timescales. These high rates of compaction were thought to contribute significantly to the high rates of relative sea-level rise (10 mm yr-1 over the past century) in the Mississippi Delta. Horton and Shennan (2009) found compaction rates of 0.4 ± 0.3 mm yr-1 during the past 4,000 years in eastern England, with higher values in large estuaries and considerable local variability depending on sediment types and drainage histories. Galloway
FIGURE 4.12 Faults (black lines) and GPS-defined horizontal velocities (red arrows) for sites in the western United States relative to stable North America. Circles are error ellipses at the 95 percent confidence level. SOURCE: Bennett et al. (1999).
et al. (2001) found that compaction of organic soils in the Sacramento Bay Delta (2–7 cm yr-1), combined with reclamation and agriculture, has resulted in islands sinking below sea level (see also “California Bay Delta Case Study” in Chapter 6).
Comprehensive studies of compaction rates for the types of geomorphic environments that dominate the U.S. west coast (see “Geographic Variation Along the U.S. West Coast” in Chapter 1) are not available. Most of these environments, particularly the peat- and mud-rich estuaries and tidal marshes, will subside as a result of compaction.
Groundwater and Petroleum-Related Drawdown and Recharge
Withdrawal of groundwater and petroleum can lower large areas of the land surface. Subsurface fluid extraction depressurizes underground reservoirs, altering the arrangement of in situ stresses within the reservoir and surrounding rock or sediment (Donaldson et al., 1995). The elastic compaction can be recovered if the fluid level rises again (e.g., Schmidt and Bürgmann, 2003), but the inelastic compaction becomes permanent, resulting in subsidence (Sun et al., 1999). Some of the best documented examples of subsidence due to groundwater withdrawal along the U.S. west coast are in California (Figure 4.13). Intense cultivation in the Santa Clara Valley during the first half of the 1900s caused the land surface to subside up to 4 m in San Jose and 0.6–2.4 m near the southern end of San Francisco Bay, putting 44 km2 below the high-tide level (Galloway et al., 2001). In the San Joaquin Valley, one of the world’s most productive agricultural regions, the land surface dropped 0.3–9 m over 75 years, mainly due to groundwater pumping and compaction. Since 1969, groundwater recharge and the supplemental use of surface water for irrigation has slowed land subsidence in both valleys.
In some cases, subsidence is partly offset by groundwater recharge. For example, long-term subsidence in the Santa Ana Basin (Los Angeles area) is ~12 mm yr-1, but groundwater recharge produces seasonal vertical oscillations of up to 55 mm (Bawden et al., 2001).
FIGURE 4.13 Areas in Washington, Oregon, and California where significant subsidence has been attributed to groundwater withdrawal (blue). The impact of groundwater withdrawal has been greater in California than in Oregon or Washington. SOURCE: Modified from Galloway et al. (2001).
Petroleum production requires the withdrawal of subsurface liquid hydrocarbons and also significant quantities of groundwater (Yuill et al., 2009). Ground surface subsidence related to petroleum withdrawal has been documented in a number of areas, including the California San Joaquin Valley, Las Vegas, New Orleans, and Houston. The best documented example is the Wilmington oil field in Long Beach, California, which subsided up to 9 m over 27 years (Mayuga and
Allen, 1969; Nagel, 2001). However, use of secondary recovery techniques, such as pumping seawater into the reservoirs to increase oil production, can stabilize compaction and halt subsidence. Large active oil fields along the coastal west U.S. coast are located mainly in the area between Santa Barbara and the Los Angeles Basin.
Current Rates of Vertical Land Motion Along the U.S. West Coast
Observations of vertical land motion in coastal California, Oregon, and Washington are given in Table 4.5. The values in the table represent the total vertical land motion, which is often caused by a combination of processes. For example, in the Los Angeles Basin, subsidence due to hydrocarbon and groundwater withdrawal, together with faulting, raised or lowered the surface elevation by upwards of 10 mm yr-1 from 1992 to 2000, with seasonal oscillations as high as 55 mm yr-1 (Box 4.3).
The spatial distribution of published data on vertical land motions is not optimal for assessing sea-level rise along the west coast. Consequently, the committee characterized the spatial variability of vertical land motion using the Scripps Orbit and Permanent Array Center velocity model and continuous GPS (CGPS) velocity data taken within ~15 km of the coast. The CGPS data provide an accurate, self-consistent vertical land motion estimate with well-defined and conservative error estimates (see Appendix A). The vertical land motion rates are shown in Figure 4.14. Most of the coastal CGPS vertical land motion rates fall within ± 3 mm yr-1 (Figure 4.14b). The average rates with obvious outliers removed (Figure 4.14c, d) are similar to longer-term estimates from leveling data for Cascadia (Burgette et al., 2009) and the San Andreas region (Appendix D). Annual rates of vertical land motion are generally positive in Washington and Oregon and generally negative in California (Figure 4.14a). This spatial pattern suggests that the tectonic boundary at the Mendocino Triple Junction is a fundamental and, most likely temporally stationary, boundary for vertical land motion. Uplift in Washington and Oregon is consistent with the buildup of interseismic strain in the Cascadia Subduction Zone as described by the CAS3D-2 model (He et al., 2003; Wang, 2007), rather than the subsidence predicted by GIA models. Subsidence in California is consistent with glacial isostatic adjustment; most GIA models predict subsidence south of the Mendocino Triple Junction (gray band in Figure 4.14b; see also Sella et al., 2007; Mazzotti et al., 2008; Argus and Peltier, 2010). As noted above (Box 4.3, Table 4.5), however, large vertical land motion signals associated with local tectonics and/or subsurface fluid movements can locally overwhelm the regional tectonic signal. This effect appears to be most prevalent toward southern California, although the paucity of
TABLE 4.5 Current Rates and Causes of Vertical Land Motion Along the U.S. West Coast
|Source||Location||Method||Period (yr)||Rate of Vertical Land Motion (mm yr-1|
|Cascadia Subduction Zone|
|Mazzotti et al. (2008)||Cascadia Subduction Zone||GPS||1993–2003||1.1–3.5|
|Burgette et al. (2009)||Cascadia Subduction Zone||Leveling||1925–2006||-0.28–3.29|
|San Andreas Fault Zone|
|Cooke and Marshall (2006) and||Palos Verdes Fault||Geodesy and modeling||Holocene–Quaternary||-0.5–0.4
|Wills et al. (2006)||Santa Monica Fault
Angeles Basin interior faults
|Bürgmann et al. (2006)||San Andreas System||InSAR||1992–2000||-2.0–1.5|
|California Aquifers and Oil Fields|
|Bawden et al. (2001)||Santa Ana Aquifer, long term
Santa Ana Aquifer, seasonal
|Argus et al. (2005)||Santa Ana Aquifer, seasonal
Long Beach Oil Field
Huntington Beach Oil Field
Wilmington Oil Field
|InSAR and GPS||1992–1999||-62–35
Spatial Variability of Vertical Land Motion and Relative Sea-Level Change in Los Angeles
Vertical land motions in the Los Angeles Basin vary on small spatial scales because of subsidence from groundwater and hydrocarbon withdrawal and active thrust faulting (Bawden et al., 2001; Lanari et al, 2004; Argus et al., 2005). Brooks et al. (2007) used InSAR to create a vertical land motion map of the Los Angeles Basin. The figure shows the rapid spatial change in land elevation at sub-15 km scales in this area.
Brooks et al. (2007) also used land motion rates to adjust local tide gage records to produce a profile of relative sea-level change along the coast. Vertical land motion differs on the west and east side of the Palos Verdes Peninsula. To the west, relative sea level was nearly constant from 1992 to 2000, with most values less than zero. To the east, approaching the Long Beach/Wilmington oil field, relative sea-level rates varied from -1.7 to 1.3 mm yr-1 and by as much as ~3 mm yr-1 over distances as short as ~5 km. The Brooks et al. (2007) results show the danger of assuming that a tide gage is representative of relative sea level for a region undergoing uplift or subsidence. Interpretation of the Los Angeles Harbor tide gage alone would miss the spatial variability in sea level to the east and assume the wrong sign of relative sea-level change to the west.
FIGURE Land motion (line-of-sight, 23 degrees inclined from vertical) from 1992 to 2000 in the Los Angeles Basin determined from InSAR (colors coded in mm yr-1) and GPS (red circles), showing variability due to tectonics and hydrocarbon and groundwater fluctuations. Tide gages are shown as yellow squares. SOURCE: Brooks et al. (2007).
FIGURE 4.14 Continuous GPS vertical land motion (VLM) rates. (A) Map of the west coast of the United States showing major tectonic boundaries and locations of GPS stations color-coded for vertical land motion rates. Squares are stations within ~15 km of the coast. Circles are other stations, which are shown to demonstrate the overall spatial variability of vertical land motion in the western United States. MTJ = Mendocino Triple Junction. (B) Vertical land motion versus latitude for the coastal GPS stations (squares in panel A), compared with predictions of current uplift from the CAS3D-2 model (green diamonds; from Table 4.5) and the ensemble of GIA models (gray shading; from Table 4.3). GPS errors are 1 standard deviation. (C) Histogram and normal density function for the Cascadia coastal stations in panels A and B. (D) Histogram and normal density function for the San Andreas coastal stations in panels A and B. In both areas, obvious outliers have been removed. SOURCE: GPS data from the Scripps Orbit and Permanent Array Center, <http://sopac.ucsd.edu/processing/refinedModelDoc.html>.
data adequate to sense the km-scale variations in vertical land motion precludes complete characterization of these strong local signals along the entire west coast.
The west coast of the United States is undergoing active vertical deformation due to a combination of tectonics, sediment compaction, fluid withdrawal and recharge, and glacial isostatic adjustment. Assessing their relative contribution to the observed vertical land motion is complicated by a shortage of data and by the wide spatial and temporal variability of the various processes. Continuous GPS measurements over the past two decades, in concert with 20th century leveling studies, show that the coast north of Cape Mendocino is rising on the order of ~1.5–3.0 mm yr-1, likely as a result of building interseismic strain along the Cascadia Subduction Zone. In contrast, the California coast south of Cape Mendocino is subsiding at a mean rate of ~1 mm yr-1 or more, although GPS-measured vertical land motions vary widely (-3.7–0.6 mm yr-1). The boundary between uplift and subsidence takes place at the Mendocino Triple Junction, highlighting the importance of regional tectonics in relative sea-level rise. Subsidence south of Cape Mendocino is consistent with models of glacial isostatic adjustment. However, more detailed analysis of potential reference frame bias and sensitivity tests of GIA models have to be carried out to determine whether GIA is responsible for the regional subsidence. Local tectonics, sediment compaction, and fluid withdrawal and recharge can cause much higher rates of subsidence or uplift than the regional mean, especially in California, but at spatial scales too small (as little as 1 km) to have a significant impact on sea-level change in the region.
The sea level along the west coast of the United States reflects contributions from both the global sea level and the local and regional processes discussed above. Tide gage data can be used to estimate rates of relative sea-level change, but only a few such estimates have been made for the west coast of the United States. Douglas (1991) compared tide gage records for 1930–1980 and found large differences in rates of sea-level rise between coastal California and northernmost California, Oregon, and Washington, consistent with a major tectonic influence (Table 4.6). Only a few other tide-gage-based estimates of sea-level change along the U.S. west coast have been published (e.g., Tebaldi et al., 2012), and most are based on the Douglas (1991) data (e.g., Peltier, 2001) or consider records from only a few gages (e.g., Nakada and Inoue, 2005; Bromirski et al., 2011; Table 4.6).
The committee obtained records from 28 tide gages along the California, Oregon, and Washington coasts archived at the Permanent Service for Mean Sea Level. Of these, 12 are currently operating, contain no long gaps, and have been recording sea level for at least 60 years, and thus were considered suitable for determining long-term trends in sea-level rise. For each gage, the rate of relative sea-level rise was determined by fitting a straight line through the monthly mean data plotted as a function of time (see Appendix A for details). The committee’s estimated rates of relative sea-level change at the 12 tide gages are given in Table 4.6 and shown geographically in Figure 4.15. Most of the gages north of Cape Mendocino (Crescent City to Neah Bay) indicate that relative sea level is falling, which is consistent with uplift associated with the buildup of interseismic strain along the Cascadia Subduction Zone, whereas most of the gages south of Cape Mendocino show that relative sea level is rising, which is consistent with land subsidence. Some gages (e.g., Friday Harbor, Seattle) deviate from these regional sea-level trends, likely as a result of local tectonic, compaction, or fluid withdrawal or recharge effects. The average rate of relative sea-level rise is 0.03 ± 1.49 mm yr-1 north of Cape Mendocino and 1.38 ± 0.64 mm yr-1 south of Cape Mendocino for the past 6–10 decades.
The change in relative sea level is what coastal residents experience and state and local managers factor into planning. To compare west coast sea-level trends with the global sea-level trend, it is necessary to adjust the relative rates of sea-level rise for changes in atmospheric pressure and vertical land motions, both of which affect the local water level (see Appendix A). Figure 4.16 illustrates the effect of these corrections on the sea-level trend for 108 years of monthly tidal data for Seattle, Washington. The slope of the blue straight line gives the rate of relative sea-level rise, in
TABLE 4.6 Rates of Relative Sea-Level Rise Estimated from U.S. West Coast Tide Gages
|Source||Tide Gage||Period||Rate of Sea-Level Rise (mm yr-1)|
|Douglas (1991)||Friday Harbor, WA||1930–1980||0.6|
|Douglas (1991)||Neah Bay, WA||1930–1980||-1.6|
|Douglas (1991)||Seattle, WA||1930–1980||2.5|
|Bromirski et al. (2011)||1930–1980||2.47|
|Douglas (1991)||Astoria, OR||1930–1980||-0.4|
|Douglas (1991)||Crescent City, CA||1930–1980||-0.9|
|Douglas (1991)||San Francisco, CA||1930–1980||1.8|
|Bromirski et al. (2011)||1930–1980||1.91|
|This report||Alameda, CA||1939–2008||0.70|
|This report||Port San Luis, CA||1945–2008||0.68|
|This report||Santa Monica, CA||1933–2008||1.41|
|Douglas (1991)||Los Angeles, CA||1930–1980||0.2|
|Douglas (1991)||La Jolla, CA||1930–1980||1.8|
|Douglas (1991)||San Diego, CA||1930–1980||1.7|
|Bromirski et al. (2011)||1930–1980||1.80|
this case, 2.01 mm yr-1. The green line in Figure 4.16 shows that the atmospheric correction is small. The atmospheric adjusted sea-level rise (using data from the National Ocean and Atmospheric Administration’s Earth System Research Laboratory) is 2.10 mm yr-1, about 4 percent higher than the relative sea-level rise.
Ideally, vertical land motions would be corrected using GPS data collected at the tide gage. However, none of the tide gage stations analyzed in this report include GPS instruments. Consequently, the committee followed the practice of using data from the closest CGPS station, as long as it is within 15 km of the gage (e.g., Mazzotti et al., 2007; Wöppelmann et al., 2007). Data from all CGPS stations within a 15 km radius of the gage also were analyzed to assess the spatial variability of vertical land motions near the tide gages.
The CGPS data were obtained from the Scripps Orbit and Permanent Array Center. Although GPS records extend back only a few decades, the committee assumed that current motions are representative of motions over the entire history of the tide gage and that land motion does not vary between the gage and the GPS station. It is likely that rates of vertical land motion near at least some of the tide gages have varied over the past century because of earthquakes, groundwater extraction and recharge, or other processes, but the absence of detailed geologic histories for each gage precluded a more sophisticated approach. The vertical land motion correction to the sea-level record was often relatively large, changing rates in one case by almost 150 percent (see Table A.2 in Appendix A). For five of the gages analyzed, correcting for vertical land motion changed the sign of sea-level change.
The rate of sea-level rise at the tide gage, adjusted for vertical land motion and atmospheric pressure, is the slope of the red line in Figure 4.16, which is 2.3 mm yr-1 for Seattle, about 15 percent higher than the rate of relative sea-level rise. Along the coast, the mean adjusted rates of sea-level rise are 1.59 ± 0.80 mm yr-1 north of Cape Mendocino and 1.02 ± 1.73 mm yr-1 south of Cape Mendocino, both of which are lower than global mean sea-level rise.
FIGURE 4.15 Rates of relative sea-level change estimated from long tide gage records (63–108 years) analyzed in this report.
FIGURE 4.16 Monthly sea level for Seattle, Washington, from 1900 to 2008. Straight-line fits to the data show the relative sea-level rise (blue line), the sea-level rise adjusted for atmospheric pressure (green line), and the sea-level rise adjusted for vertical land motion and atmospheric pressure (red line).
Sea level at any given place and time depends on the global sea level and the net contribution of atmospheric, oceanographic, geologic, and anthropogenic processes operating in the area. Processes that raise relative sea level in the northeastern Pacific Ocean include warm phases of climate oscillations (El Niños, positive phase of the PDO) and land subsidence due to glacial isostatic adjustment, sediment compaction, and the withdrawal of groundwater or hydrocarbons. Processes that lower relative sea level include cool phases of climate oscillations (La Niñas, negative phase of the PDO), gravitational and deformational effects of modern melting of glaciated land masses, and land uplift due to tectonics or fluid recharge.
The highest sea levels recorded along the west coast are usually associated with El Niño events, which can elevate coastal sea level by 10–30 cm for several winter months. Cool climate phases have less influence on local sea level than warm climate phases. Changes between warm and cool climate phases, which occur on seasonal to multidecadal timescales, cause large-amplitude variations in the relative sea-level trend.
Modern melting of glaciers and ice sheets adds new water to the ocean basins and produces gravitational and deformational effects that create regional patterns of relative sea-level change. The glaciated land masses that most effect sea level along the west coast of the United States are Alaska, which is close, and Greenland and Antarctica, which are large. The gravitational and deformational effects reduce the contribution of melting of these three ice sources to relative sea-level rise for 1992–2008 by about 42 percent along the north coast (Neah Bay, Washington), 24 percent along the central coast (Eureka,
California), and 14 percent along the south coast (Santa Barbara, California).
Vertical land motions along the west coast of the United States are caused by a complex combination of tectonics, glacial isostatic adjustment, sediment compaction, and fluid withdrawal and recharge. The area straddles two tectonic regimes: (1) the Cascadia Subduction Zone, where the buildup of interseismic strain is causing coastal uplift north of Cape Mendocino, California; and (2) the San Andreas Fault Zone, where the lateral motion of the lithospheric plates produces relatively little vertical land motion south of Cape Mendocino. Glacial isostatic adjustment is producing uplift in northernmost Washington, which had been covered by the former Laurentide Ice Sheet, and subsidence in areas peripheral to the center of the former ice mass, including the rest of Washington, Oregon, and California. Land levels in some areas also are rising or sinking because of local tectonics, compaction of wetland sediments, and/or fluid withdrawal or recharge. Continuous GPS measurements over the past two decades and leveling studies over the past eight or nine decades shows that the coast north of Cape Mendocino is rising at rates of 1.5–3.0 mm yr-1 and the coast south of Cape Mendocino is subsiding at a mean rate of about 1 mm yr-1, although with considerable spatial variability (-3.7–0.6 mm yr-1).
Tide gage records along the west coast of the United States indicate that relative sea-level change is variable along the coast. Most gages north of Cape Mendocino show relative sea-level fall for the past 6–10 decades, consistent with coastal uplift along the Cascadia Subduction Zone. Most gages south of Cape Mendocino show relative sea-level rise, consistent with land subsidence. When adjusted for vertical land motions and for atmospheric pressure effects, the rates of relative sea-level rise along the U.S. west coast are lower than the rate of global mean sea-level rise.
Although rates of sea-level rise are relatively low along the west coast of the United States, the combination of sea-level rise and winter storms increases the potential for significant coastal damage. Historically, most coastal damage has occurred when storm surges and large waves coincided with high astronomical tides and El Niños—a combination that can raise short-term sea level above sea levels projected for 2100. All climate models project ample winter storm activity, but a clear consensus has not yet emerged on whether storm frequency or intensity will change in the northeast Pacific. Several climate models predict a northward shift in the North Pacific storm track over the 21st century, and some observational studies report that a northward shift has been detected. However, most observational records are not long enough to determine whether a shift has begun.
Several observational studies have reported that high waves have been getting higher and that winds have been getting stronger in the northeastern Pacific over the past few decades. The magnitude and cause of these changes are under investigation; at least part of the observed increase likely reflects natural climate variability. But even if storminess does not increase in the future, sea-level rise will magnify the adverse impact of storm surges and high waves on the coast of California, Oregon, and Washington.