Oceanography of the Pacific Northwest Coastal Ocean and Estuaries
with Application to Coastal Ecosystems
Barbara M. Hickey1 and Neil S. Banas
School of Oceanography Box 355351
University of Washington, Seattle, WA 98195-7940
Tel.: 206 543 4737
e-mail: [email protected]
Submitted to Estuaries
May 2002
1Corresponding author.
Hickey and Banas 2
Abstract
This paper reviews and synthesizes recent results on both the coastal zone of the U.S. Pacific Northwest
(PNW) and several of its estuaries, as well as presenting new data from the PNCERS program on links between the
inner shelf and the estuaries, and smaller-scale estuarine processes. In general ocean processes are large-scale on this
coast: this is true of both seasonal variations and event-scale upwelling-downwelling fluctuations, which are highly
energetic. Upwelling supplies most of the nutrients available for production, although the intensity of upwelling
increases southward while primary production is higher in the north, off the Washington coast. This discrepancy is
attributable to mesoscale features: variations in shelf width and shape, submarine canyons, and the Columbia River
plume. These and other mesoscale features (banks, the Juan de Fuca eddy) are important as well in transport and
retention of planktonic larvae and harmful algae blooms.
The coastal-plain estuaries, with the exception of the Columbia River, are relatively small, with large tidal
forcing and highly seasonal direct river inputs that are low-to-negligible during the growing season. As a result
primary production in the estuaries is controlled principally not by river-driven stratification but by coastal upwelling
and bulk exchange with the ocean. Both baroclinic mechanisms (the gravitational circulation) and barotropic ones
(lateral stirring by tide and wind) contribute to this bulk exchange, though tidal circulations appear to dominate
during the low-riverflow growing season on ~monthly scales. Because estuarine hydrography and ecology are so
dominated by ocean signals, the coast estuaries, like the coastal ocean, are largely synchronous on seasonal and event
time scales, though intrusions of the Columbia River plume can cause strong asymmetries between Washington and
Oregon estuaries during spring downwelling conditions. Property coherence increases between spring and summer as
wind forcing becomes more spatially coherent along the coast. Estuarine habitat is structured not only by large scale
forcing but also by fine scale processes in the extensive intertidal zone, such as differential solar heating or
differential advection by tidal currents.
Hickey and Banas 3
Introduction
Recent results on the physical oceanography of the U.S. Pacific Northwest (PNW) coastal region are
integrated in this paper to provide a framework for understanding ecosystem variability. The coastal region important
to the regional ecosystem includes both the nearshore zone and the coastal estuaries. Many species utilize both these
regions at different life stages. For example, Dungeness crab frequently utilize coastal estuaries for the first year of
their life, re-entering the ocean to become part of the fishery as juveniles (Gunderson et al. 1990). Salmon, on the
other hand, utilize the estuary at both the beginning and end of their life cycles and the coastal zone as adults.
As we will demonstrate, ocean variability in nearshore regions of the U.S. West Coast and, in particular, its
coastal estuaries, is distinctly different from that in estuaries and nearshore regions of the U.S. East Coast. The West
Coast is embedded within an Eastern Boundary Current System; the East Coast is embedded within a Western
Boundary System. Thus, whereas the West Coast is dominated by upwelling, the East Coast is not; whereas
upwelling provides plentiful nutrients to the West Coast and its estuaries, on the East Coast nutrients are more
commonly supplied by river outflow.
Ocean variability along the West Coast is generally very large scale (> 500 km), a result of large-scale
atmospheric systems (Halliwell and Allen 1987). Nevertheless, we will show that significant alongshore gradients
occur in coastal productivity in the PNW. Moreover, we will demonstrate that mesoscale features such as banks and
submarine canyons play important and even critical roles in ecosystem function.
In many ways, at least during the summer growing season, coastal estuaries in the PNW may be considered
as extensions of the coastal ocean: as we will discuss, flushing rates are on the order of a few days and property
variability is controlled by changes at the ocean end of the estuary rather than by riverflow at its head. Thus, like the
ocean processes, both the several-day and seasonal fluctuations that occur over the growing season occur nearly
simultaneously across the PNW coast estuaries. However, actual values of water properties such as temperature,
salinity and velocity will differ depending on the specific estuary configuration.
In the following, the large-scale processes acting on the PNW coastal zone are first described. This is
followed by a discussion of nutrient variability (Section 2). With this setting the effects of important mesoscale
features such as submarine banks, canyons and river plumes are presented (Section 3). Following the description of
coastal processes, processes and variability within the coastal estuaries is described (Section 4), with particular focus
on the estuaries studied in the PNCERS program. The interaction of the coastal ocean with these estuaries and the
Hickey and Banas 4
similarity of water property variation in the several estuaries are demonstrated using time series and survey data
collected in the PNCERS program.
1. Large scale processes in the Pacific Northwest coastal ocean
a) Seasonal variability
The U.S. Pacific Northwest coastal zone is embedded within the California Current System (CCS), a system
of currents with strong interannual, seasonal and several-day (event) scale variability (Fig. 1) (Hickey 1998). The
California Current System includes the southward California Current, the wintertime northward Davidson Current,
the northward California Undercurrent, which flows over the continental slope beneath the southward upper layers,
as well as "nameless" shelf and slope currents with primarily shorter-than-seasonal time scales. The PNW includes
one major river plume (the Columbia), several smaller estuaries, and (primarily in the north) numerous submarine
canyons. The dominant scales and dynamics of the circulation over much of the CCS are set by several
characteristics of the physical environment; namely, 1) strong alongshore winds; 2) large alongshore scales for both
the winds and the bottom topography (Halliwell and Allen 1987); and 3) a relatively narrow and deep continental
shelf. Because of these characteristics, coastal-trapped waves (disturbances that travel northward along the shelf and
slope) are efficiently generated and propagate long distances along the continental margins of much of western North
America. Thus, much of the variability in the PNW is caused by processes occurring southward of the region (i.e.,
“remote forcing”). Because of the generally southward alongshore wind stress in spring and summer, coastal
upwelling is the dominant process controlling water property variability (see review in Smith 1995).
The California Current flows southward year-round offshore of the U.S. West Coast from the shelf break to
a distance of 1000 km from the coast (Hickey 1979, Hickey 1998) (Fig. 1). The current is strongest at the sea
surface, and generally extends over the upper 500 m of the water column. Seasonal mean speeds are ~10 cm s-1. The
California Undercurrent is a relatively narrow feature (~10-40 km) flowing northward over the continental slope of
the CCS at depths of about 100-400 m as a nearly continuous feature, transporting warmer, saltier Southern water
northward along the coast. The Undercurrent has a jet-like structure, with the core of the jet located just seaward of
and just below the shelf break and with peak speeds of ~30-50 cm s-1. The Undercurrent provides a possible
northward transport route for larvae, larval fish and even phytoplankton seed stock. Because of its proximity to the
Hickey and Banas 5
shelf break, the Undercurrent is the source of much of the water supplied to the shelf during coastal upwelling. The
onshore transport of this water during upwelling offers a mechanism for onshore transport of plankton entrained in
the Undercurrent.
A southward undercurrent (the “Washington Undercurrent”) occurs over the continental slope in the winter
season in the PNW (Werner and Hickey 1983). This undercurrent occurs at deeper depths than the northward
undercurrent (~300-500 m). The existence of this undercurrent, like that of the northward undercurrent, likely
depends on the co-occurrence of opposing wind stress and alongshore pressure gradient forces. The Davidson
Current flows northward in fall and winter north of Point Conception. This northward flow is generally broader
(~100 km in width) and sometimes stronger than the corresponding subsurface northward flow in other seasons (the
"Undercurrent") and extends seaward of the slope.
Currents and water properties of the CCS both over the shelf and in the region offshore of the shelf undergo
large seasonal fluctuations. The California Current and Undercurrent are strongest in summer to early fall and
weakest in winter. The Davidson Current is strongest in winter. Seasonal mean shelf currents are generally southward
in the upper water column from early spring to summer and northward the rest of the year. Over the shelf, the
seasonal duration of spring-summer southward flow usually increases with distance offshore and with proximity to
the sea surface (Strub et al. 1987b). A northward undercurrent is commonly observed on shelves during the summer
and early fall. Off the coast of Vancouver Island a northward flowing buoyancy driven current exists year-round
from the coast to at least mid shelf (the Vancouver Island Coastal Current) (Thomson 1981, Hickey et al. 1991). This
current opposes the southward shelf break jet current that connects to southward flow off the outer Washington shelf.
Seasonal fluctuations are continuous with similar fluctuations in the Alaskan gyre and the majority of
seasonal change in the sea surface height and hence geostrophic currents have been shown to occur within a few tens
of kilometers of the coast (Strub and James 2002). Seasonal currents are largely driven by alongshore wind stress
(see review included in Batteen 1997). Satellite altimetry data illustrate that seasonal features gradually migrate
offshore and out into the main California Current, so that alternating seasonal bands of northward and southward
flow (superimposed on the long term mean California Current) are observed as far as several hundred kilometers
from the coast (Strub and James, 2002).
The transition of currents and water properties over the shelf and slope between winter and spring, the
"Spring Transition," is a sudden and dramatic event in the CCS (Strub et al. 1987a). Along much of the coast, during
the transition, sea level drops at least 10 cm, currents reverse from northward to southward within a period of several
Hickey and Banas 6
days and isopycnals slope upward toward the coast in response to coastal upwelling (Smith 1995). The transition is
driven by changes in the large scale wind field and these changes are a result of changes in the large scale
atmospheric pressure field over the CCS. A similar rapid transition between summer upwelling and fall downwelling
oceanic characteristics does not occur (Strub and James 1988).
b) Several-day time scales
In spite of strong seasonal variability in the PNW, the dominant variability occurs at several-day time scales
(Hickey, 1989). Thus, on the shelf, seasonal conditions as described above are often reversed for shorter periods of
time. Because of bottom friction, reversals occur more frequently nearshore (Brink et al. 1987). Fluctuations in
currents, water properties and sea level over the shelf at most locations are dominated by wind forcing, with typical
scales of 3-10 d. A schematic of the locally wind-driven ocean surface circulation in the PNW for winds toward the
south (“fair weather”) and winds toward the north (“poor weather”) is shown in Fig. 2. During periods of fair
weather the stress of the southward winds on the sea surface accelerates the coastal currents, producing offshore- and
alongshore-directed currents in the surface Ekman layer, and alongshore currents elsewhere in the water column
(geostrophically balanced to first order, Allen et al. 1995). Under these conditions, plumes of fresher water
originating at coastal estuaries tend to spread offshore and to the south (Garcia-Berdeal et al. 2002). Upwelling
occurs within a few kilometers of the coast (typically, within one Rossby radius, about 10 km). During periods of
poor weather circulation patterns reverse and freshwater plumes move back onshore (Hickey et al. 1998).
The action of the alongshore wind stress on the sea surface results in an alongshore, baroclinically and
frictionally sheared coastal jet in the direction of the wind stress (see model studies in Allen et al. 1995 and Allen
and Newberger 1996). The location of maximum speed in the coastal jet moves progressively farther offshore as
long as the wind stress continues to act. Typical cross-shelf velocity profiles for Washington and Oregon are shown
in Hickey (1989) and Huyer and Smith (1974). The speed maximum most typically occurs near mid shelf (Hickey,
1989). Velocity can decrease by a factor of more than two from top to bottom (or even reverse sign) and by a factor
of more than two from the inner shelf to the mid shelf. The cross shelf and vertical structure of the velocity field is
important when considering transport of larvae by the coastal current system (Rooper et al. 2002).
The cartoon of shelf circulation shown in Fig. 2 does not include the important effects of remote forcing.
Alongshore gradients in alongshore coastal wind stress are significant, with stronger winds (typically upwelling-
favorable) south of the PNW in the spring and summer (Hickey, 1979). Because the West Coast north of Point
Conception has no promontories sufficient to significantly disrupt the coastal wave guide, coastal-trapped waves are
Hickey and Banas 7
generated and these waves to first order add to the local wind-generated alongshore currents. Waves off central
Washington have been shown to originate primarily from northern California (Battisti and Hickey 1984). At any
given time and location, the ratio of remote and local forcing varies and their relative importance has significant
interannual variability due to the dependence on alongshore wind stress gradients (Battisti and Hickey 1984). In
summer, free waves are usually important in the PNW, particularly at more northern latitudes such as the British
Columbia coast (Hickey et al. 1991). In winter, local wind forcing dominates in the PNW, especially in regions
where winter storms are accompanied by strong northward winds whose strength increases in the direction of
propagating waves.
Fluctuations in cross-shelf velocity are not as well understood as those in alongshelf velocity. Although
model results show onshore and offshore flow in the surface and bottom boundary layers after sufficient spin-up of
the system to an applied wind stress (e.g., Allen et al. 1995, Allen and Newberger 1996), observed velocities are
frequently much more complex than those predicted by model studies (Brink et al. 1994). The relatively short
alongshore coherence scales (~10-20 km) appear to belie the large-scale nature of the atmospheric forcing mentioned
above. In general, the cross-shelf flows appear to be highly three-dimensional, thus including effects of smaller-scale
features in the bottom topography and the coastline as well as in the wind field.
Meandering jets and an energetic eddy field carry much of the variance in the California Current off
northern and central California (Strub et al. 1991). These jets, which extend over at least the upper 200 m of the
water column carry recently upwelled coastal water and associated biological production seaward of the shelf to
distances of several hundred kilometers. The strongest jets are generated near coastal promontories where flow
separates from the coast, the resulting jet becoming unstable (see model studies in, e.g., Batteen 1997). The
meandering jet that separates from the coast near southern Oregon can be traced southward along the whole
California coast (Barth et al. 2000). In contrast, most of the PNW coastal region is not dominated by meandering
jets. Satellite-derived patterns of sea surface temperature show only one region where upwelling appears to be
enhanced off the Washington coast, and the colder water upwelled in this area flows southward down the shelf rather
than across the shelf and coastal margin (Fig. 3).
2. Nutrient supply in the Pacific Northwest coastal zone
Hickey and Banas 8
The CCS contains water of three types: Pacific Subarctic, North Pacific Central and Southern (sometimes
termed "Equatorial"). Pacific Subarctic water, characterized by low salinity and temperature and high oxygen and
nutrients, is advected southward in the CCS (Hickey 1979, 1998). North Pacific central water, characterized by high
salinity and temperature and low oxygen and nutrients, enters the CCS from the west. Southern water, characterized
by high salinity, temperature and nutrients, and low oxygen, enters the CCS from the south with the northward
undercurrent. In general, salinity and temperature increase southward in the CCS and salinity also increases with
depth.
Upwelling along the coast brings colder, saltier and nutrient richer water to the surface adjacent to the coast
all along the U.S. West Coast (Huyer 1983). In general, the strength and duration of upwelling (as seen at the sea
surface) increases to the south in the PNW. Maximum upwelling occurs in spring and summer. With the exception of
regions affected by the Columbia plume, stratification in the CCS is remarkably similar at most locations and is
largely controlled by the large-scale advection and upwelling of water masses as described above (Huyer 1983).
In contrast to most U.S. East Coast environments, the shelf is relatively narrow and the nutricline is
fortuitously positioned so that nutrient-rich deeper water can be effectively brought to the surface by the wind-driven
upwelling. In contrast to most East Coast coastal areas, nitrate input to the ocean from coastal rivers is negligible
even from the Columbia, which accounts for the majority of the drainage in this region (Barnes et al. 1972). Both
seasonal and event-scale patterns of all macronutrients on the continental shelf are dominated by seasonal and event-
scale patterns in upwelling processes (Fig. 4) (Landry et al. 1989, Hickey 1989). Wind-driven upwelling of nutrients
from deeper layers fuels coastal productivity, resulting in both a strong seasonal cycle and several-day fluctuations in
productivity that follow changes in the wind direction and, hence, upwelling. During an upwelling event,
phytoplankton respond to the infusion of nutrients near the coast and this "bloom" is moved offshore, continuing to
grow while depleting the nutrient supply. When winds reverse (as occurs during storms), the bloom moves back
toward shore where it can contact the coast or enter coastal estuaries (Roegner et al. 2002).
3. Mesoscale features and along-coast gradients
Hickey and Banas 9
The large-scale nature of oceanographic processes in the PNW has been described in the preceding
sections. In the section below we describe several mesoscale features that cause local variation in flow patterns and
response to forcing. These features may play a role in ecosystem variability that is as significant or even more
significant than the large-scale processes. Important mesoscale features include river plumes, submarine canyons,
banks and coastal promontories. Such features can modulate the local upwelling response, they can alter flow
patterns and they can change environmental characteristics such as turbidity, mixed layer depth, stratification and
mixing rates. For these reasons, such features are likely of particular importance to phytoplankton/zooplankton
production, growth and retention as well as larval transport and retention.
a) Along-coast gradients in productivity and forcing
Time series of vertically integrated chlorophyll for the Washington and Oregon shelves suggest that
chlorophyll is greater on the Washington shelf (Fig. 5). This result, derived from averaging a number of unrelated
surveys in the 1970s and 1980s (hence, data are both temporally and spatially aliased), is confirmed by satellite-
derived images of ocean color (Strub et al. 1990) and also by recent surveys of the Columbia plume region (Peterson
pers. comm.). Off Oregon, only over Heceta Bank does chlorophyll approach values seen off the Washington coast.
Limited studies also suggest higher primary productivity off the Washington coast (Anderson 1972), suggesting that
the alongshore difference is not simply due to greater retention on the Washington shelf. The greater productivity is
also observed higher in the food chain; e.g., in euphausiids and copepods (Landry and Lorenzen 1989). Moreover,
juvenile salmon are observed more frequently off the Washington coast (Pearcy 1992).
The apparently greater richness of the Washington coast is particularly surprising because the gradients in
the primary forcing, alongshore wind stress, increase in the opposite direction; i.e., the amplitude of upwelling-
favorable coastal winds decreases northward in the PNW (Hickey, 1979). Southward stress frequently differs by
almost a factor of two between southern Oregon and northern Washington. In addition, the duration of coastal
upwelling also decreases seasonally towards the north. In spite of the alongshore decrease in wind stress we note that
the seasonal variation in macronutrient supply to the mid-shelf does not differ substantially between the two regions
(Fig. 4). This result can be attributed to several processes: differences in circulation patterns due to differences in
shelf structure; upwelling enhancement by canyons (see subsection c below); and influences of the Columbia plume
(see subsection d below).
Hickey and Banas 10
The width and shape of the continental shelf varies substantially in the PNW (Fig. 6). Most important, the
width of the shallow nearshore region (arbitrarily defined as the area shallower than 100 m) is greater by more than a
factor of two (~50 km) off Washington than off Oregon, with the exception of Heceta/Stonewall Bank in southern
Oregon. The shallow nearshore region is favored by the juveniles and/or returning larvae of many species (e.g.,
Rooper et al. 2002). Model studies in Allen et al. (1995) show that a wider, gently sloping shelf results in slower
circulation (i.e., possibly greater retention). Also, on such a shelf the upwelling flow tends to be more concentrated
in the bottom boundary layer than in steeper regions. This might explain the apparently similar levels of
macronutrients on the Washington and Oregon shelves in spite of the substantially weaker wind stress to the north.
b) Banks
On a West Coast-wide survey of domoic acid in surface waters in 1998, high values of this toxin were
measured only in the vicinity of known topographic features such as banks or offshore islands (Fig. 7) (Trainer et al.
2000). Domoic acid frequently results in closures of razor clam beaches along the Washington coast, it has been
measured in Dungeness crabs, and it has been responsible for a number of mortalities in seabirds and marine
mammals in California (Trainer et al. 2002). Domoic acid is associated with the diatom Pseudo-nitzschia. It seems
likely that in regions where large coastal promontories occur, such as off southern Oregon and northern and central
California, plankton and larvae can be swept offshore and southward by the meandering jets and/or eddies that form
where coastal jets detach from the shelf. These plankton and larvae likely return to the coast rarely, if at all.
Available nutrients would have been depleted well before the meander could return to the coast (meander scales are
several weeks to months). On the other hand, in regions where banks and more complex mesoscale topography
occur, such as offshore of the Strait of Juan de Fuca (the Juan de Fuca eddy) or Heceta/Stonewall Bank off the
central Oregon coast, flow patterns favoring retention, and perhaps more continuous macronutrient supply as well,
are more likely. Maps of ocean pigment clearly show that chlorophyll is greater and located father offshore in the
vicinity of both of these features (Strub et al. 1990). Under weak-southward-wind conditions or during periods of
northward winds associated with storms, plankton and larvae in these retention areas can move inshore to settle on
the coast or enter coastal estuaries (e.g., Trainer et al. 2002).
c) Submarine canyons
Hickey and Banas 11
The Washington coast is indented by a number of submarine canyons (Fig. 6). Upwelling of nutrient-rich
water is enhanced several-fold in the presence of such canyons (see model study in Allen 1996 and observations in
Hickey 1989). The upwelling from canyons may at least partially compensate for the generally weaker upwelling
winds that occur off the Washington coast relative to Oregon.
Canyons also alter regional circulation patterns in a manner that increases local retention (Hickey 1995,
1997). In particular, counterclockwise circulation patterns are generally observed both within and over submarine
canyons (although not necessarily extending to the sea surface) (Fig. 8). Such eddies provide an effective mechanism
for trapping particles such as suspended sediment or food for organic detritus (Hickey 1995). Zooplankton densities
are frequently denser over the submarine canyons off the Washington coast (Swartzman and Hickey 2001).
d) The Columbia River plume
The Columbia River provides over 77% of the drainage between the Strait of Juan de Fuca and San
Francisco Bay (Barnes et al. 1972). The plume from the Columbia River likely has major ecological effects in the
PNW. On a seasonal basis, the plume from the Columbia flows northward over the shelf and slope in fall and winter,
and southward well offshore of the shelf in spring and summer. In winter, the plume has a dramatic effect on the
Washington coast, producing time-variable currents as large as the wind-driven currents (Hickey et al. 1998). In
summer, fresh water from the Columbia gives rise to the low-salinity signal and associated front used to trace the
meandering jet that separates from the shelf at Cape Blanco (Huyer 1983). Both observational and modeling studies
show that the plume is a "moving target," changing direction, thickness and width with every change in local wind
strength or direction (Fig. 9) (Hickey et al. 1998, Garcia-Berdeal et al. 2002).
River plumes are generally turbid, thereby providing less light for plankton growth, while at the same time
providing better cover from grazing for higher trophic levels. Plumes also provide retention areas; eddy-like features
are generated within a plume under both steady (Garcia-Berdeal et al. 2002) and unsteady (Yankovsky et al. 2001)
outflow conditions. Inshore of the Columbia plume on the Washington coast in winter, a retentive circulation pattern
occurs during periods of upwelling-favorable winds (Hickey et al. 1998). Deep mixing is inhibited by high
stratification at the base of the plume, thus tending to keep plankton within the euphotic zone. Plumes alter regional
current patterns in the upper layers, providing along-plume jets for rapid transport, and convergences and trapping at
frontal boundaries on the edges. The model example shown in Figure 10 illustrates these retentive circulation
Hickey and Banas 12
patterns and also the along-plume jets. Information on stratification enhancement is given in Garcia-Berdeal et al.
(2002) using a numerical model and Hickey et al. (1998) from observations of the Columbia plume. Recent studies
suggest that plume edges are preferred feeding sites for zooplankton. The fact that juvenile salmonids are frequently
found near the Columbia plume (Pearcy 1992) may be due to the local retention patterns or to frontal convergences,
either of which might enhance food availability in this region.
Other than the Columbia, river plumes on the PNW coast are relatively small, and satellite imagery suggests
that their traceable effects are confined to within one or two tidal excursions of the mouth of the river or estuary (not
shown). Other river or estuarine plumes include those from Grays Harbor and Willapa Bay, Washington and Coos
Bay, Oregon.
Both the structure and magnitude of the Columbia River plume have significant interannual variability.
During years of high snowpack in the Pacific Northwest (such as 1999), very fresh water from the plume can flood
the major coastal estuaries north of the Columbia estuary for prolonged periods, reversing the normal estuarine
density and salinity gradients over much of the estuaries. Because such plume intrusions would not occur in estuaries
off the Oregon coast, the presence or absence of the plume may provide an important environmental distinction
between these estuaries as well as between nearshore coastal regions (see Section 4d below).
e) The Strait of Juan de Fuca
The counterclockwise cold eddy off the Strait of Juan de Fuca (also called the "Tully" eddy; Tully 1942) is
situated southwest of Vancouver Island and offshore of northern Washington. The eddy, which has a diameter of
about 50 km, forms in spring and declines in fall (Freeland and Denman 1982). The eddy is a dominant feature of
circulation patterns off the northern Washington coast and is visible in summertime satellite imagery as a relative
minimum in sea surface temperature (Fig. 3) and, generally, a relative maximum in chlorophyll a. The seasonal eddy
is a result of the interaction between effluent from the Strait, southward wind-driven currents along the continental
slope and the underlying topography, a spur of the Juan de Fuca submarine canyon. A connection between the eddy
and the Washington coast was demonstrated in July 1991, when oil that spilled in the eddy was found on the
Washington coast 6 days later (Venkatesh and Crawford 1993). Recent preliminary studies with drifters introduced
into a diagnostic numerical model for a summer period in 1998 suggest that drifting particles can escape from the
eddy to flow southeastward along the Washington shelf (MacFadyen et al. 2002). During storms, onshore flow in the
Hickey and Banas 13
surface Ekman layer moves drifter pathways closer to the coast and even reverses the path to a northward direction.
Pathways of drifters deployed in 2001 in the field were consistent with these modeled pathways; a drifter that had
moved southward from its deployment site in the eddy reversed direction during a storm and moved back up the
coast, approaching as close as 1 km from the beach. Thus, it seems likely that marine organisms residing in the Juan
de Fuca eddy can, under certain ocean conditions, impact the Washington coast.
The photic zone in the Juan de Fuca eddy region is characterized by high ambient macronutrients supplied
by wind mixing, episodic wind-driven upwelling, topographically controlled upwelling (Freeland and Denman 1982)
and the outflow from Juan de Fuca Strait where deep, nutrient-rich water is brought to the surface by estuarine
circulation and tidal mixing (Mackas et al. 1980). Thus, although the ultimate source of nutrients for the eddy is the
same as that in a nearshore coastal upwelling region (California Undercurrent water), infusion of upwelled nutrients
into the eddy likely occurs on different time scales and with different rates than in regions adjacent to the coast.
Repeated surveys on the northern Washington coast have demonstrated that when domoic acid is present off
the coast it is usually within or near the Juan de Fuca eddy (Trainer et al. 2002). The diatom Pseudo-nitzschia is
always present in significant numbers when the acid is present and these diatoms are known toxin producers. A
relationship between toxification of clams and onshore water movement in storms has been demonstrated in a time
series (Trainer et al. 2002) for at least one toxic event. Growing conditions in this mesoscale feature must differ from
the large scale conditions along the coast where toxin is not usually produced; and thus, in this case, the mesoscale
dynamics are as important as the large scale dynamics in determining the nature of the ecosystem.
4. Pacific Northwest estuaries
We have shown that the oceanic environment of the Northwest coast is broadly coherent, although
mesoscale features may be important to local circulation as well as to the ecosystem. In the following section we
describe the physical characteristics and dynamics of estuaries linked to that coastal ocean. The geomorphology,
freshwater forcing, and tidal regime of these estuaries is first described, with comparison to the more commonly
studied coastal-plain estuaries of the East Coast (a). This is followed by a brief description of the PNCERS estuary
dataset (b). Coupling between processes in the coastal ocean and the estuaries is next described including a
discussion of alongshore coherence between the estuaries and the important effects of Columbia plume intrusions (c).
Last, significant differences between intertidal bank and channel water properties are illustrated (d).
Hickey and Banas 14
a) Physical characteristics and forcing
The four estuaries studied in PNCERS (Grays Harbor, Willapa Bay, Yaquina Bay, and Coos Bay, see
detailed map in Figure 11) are members of a chain of small estuaries that begins along the Washington coast, spans
the coast of Oregon, and continues into northern California. Most of these estuaries are drowned river valleys,
formed from sea level rise during the last 10,000 y. Some have also been shaped by ocean-built bars, either partially
(e.g., Willapa Bay, Washington) or entirely (e.g., Netarts Bay, Oregon). Emmett et al. (2000) reviews the geography
of this system in detail.
Indices of geomorphology and tide and river forcing for the PNCERS estuaries are given in Table 1. For
comparison, the same parameters are included for the Columbia River estuary; San Francisco Bay and South San
Francisco Bay alone; Naragansett Bay; Chesapeake Bay and its tributary the James River; and Plum Island Sound, a
small embayment on the Massachusetts coast. Except where otherwise marked, data are from the NOAA National
Estuarine Inventory Data Atlas (NOAA 1985). Volume parameters, which are particularly difficult to define and
measure (e.g., Malamud-Roam 2000), are here calculated by simple, approximate methods for the sake of
uniformity, and thus only gross patterns among the area and volume parameters are significant. Volume is calculated
as the product of mean depth and surface area at mean sea level (MSL), a method which gives errors up to ~20% in
comparison with other published figures (NOAA/EPA 1991). Mean tidal prism volume is reported as a percentage of
volume at high water, which is calculated as MSL volume plus half the tidal prism itself.
Coos Bay is only a few times larger than tiny Plum Island Sound, but is nevertheless the largest of the
Oregon estuaries. Grays Harbor and Willapa Bay, the two coastal-plain estuaries north of the Columbia, are an order
of magnitude larger, comparable in volume and morphology to South San Francisco Bay. Both Washington estuaries
consist of multiply-connected channels 10-20 m deep surrounded by wide mud and sand flats. Half or more of the
surface area of these estuaries lies in the intertidal zone. Significantly, even the smaller, narrower estuaries of Oregon
have similar percentages of intertidal area (Table 1, Percy et al. 1974).
Tides on this coast are mixed-semidiurnal, with spring-neap amplitude variation on the order of 50%
(Emmett et al. 2000). Mean tidal ranges, as shown in Table 1, are generally twice as large as on the outer Atlantic
Coast. The combination of large tidal range with broad, open intertidal surface area yields tidal prisms that are large
fractions (30-50%) of total volume. This result holds very generally for Northwest coast estuaries, and is a marked
Hickey and Banas 15
difference between these systems and all but the smallest of their counterparts on other North American coasts.
These large tidal prisms suggest that flushing by tidal action is probably important in all these estuaries, even those
that receive significant riverflow (Dyer 1973). Tidal excursions, as estimated from current measurements in Willapa,
Grays, and Coos, are 12-15 km, significant fractions (25-50%) of the length of the estuaries.
Table 1 includes long-term mean flows for the lowest- and highest-flow months of the year, and, as a
measure of the strength of river forcing relative to estuary size, the "river-filling time," volume divided by flow rate.
The output from the Columbia River is two orders of magnitude larger than riverflow into the other coastal estuaries.
With the exception of the Columbia, these estuaries receive freshwater input from local rainfall only, not from
snowmelt, and their riverflows show strong seasonality in concert with the winter storm and summer dry- and fair-
weather seasons. Local riverflow peaks during winter storms and is negligible during late summer. The seasonal
variation is generally several times greater than in East Coast estuaries, though flood and drought events beyond the
mean seasonal cycle have not been considered here. As a result we might expect the hydrodynamic classification of
Northwest estuaries to change dramatically between seasons, or even—where flushing and adjustment times are
short—between individual wind events.
This riverflow pattern yields a seasonal hydrographic cycle that contrasts strongly with traditional models of
temperate partially mixed estuaries, with important ecological implications. Tyler and Seliger (1980), for example,
show that primary production in Chesapeake Bay is controlled by "stratification dependent pathways" reminiscent of
the seasonal dynamics of the open-ocean mixed layer. In that estuary, in winter, mixing by wind and tide erases
stratification and resuspends nutrients, while in spring and summer increased riverflow and solar heating produce
strong stratification and reduced vertical exchange. In such a system, stratification controls on vertical mixing are
crucial to determining plankton growth rates and the potential for phytoplankton blooms, as in San Francisco Bay
(Lucas et al. 1999). In sharp contrast, in Willapa Bay stratification is in general very low during summer, when
riverflows are low, and high during the winter, when riverflow peaks (Banas et al. 2002). Vertical, one-dimensional,
stratification-centered models of primary productivity thus would not apply here even at the coarsest level. Rather,
during the growing season in Pacific Northwest estuaries, hydrography, nutrient levels, and biomass all appear to be
controlled less by in situ processes than by mesoscale processes in the coastal ocean (Hickey et al. 2002, Roegner et
al. 2002). Below we consider this ocean-estuary coupling in more detail.
b) PNCERS observations
Hickey and Banas 16
During PNCERS, arrays of moored sensors were maintained in three coastal estuaries (Grays Harbor,
Willapa Bay and Coos Bay) as well as at two sites in the nearshore costal ocean, one off Washington, the other off
southern Oregon. Locations are shown in Figure 11. Moored sensors included S4 current meters or ADCPs, and
Seabird C-T sensors or Aanderaa current meters equipped with conductivity and temperature sensors. Estuarine
instruments were set in the lower water column, sustained by a taut wire mooring. Sampling interval was less than 30
minutes in the estuaries and hourly on the coast. Two arrays were maintained in Willapa Bay and Coos Bay and one
in Grays Harbor. The longest time series (temperature) spans 3 years. In general, salinity time series are much shorter
due to fouling and clogging problems. Hydrographic sections were made with a SeaBird 19 CTD at sporadic
intervals when mooring instruments were exchanged or cleaned. At the same time as PNCERS, under the direction of
Dr. Jan Newton, Washington State Department of Ecology maintained sensors at a number of locations in Willapa
Bay and also sampled hydrographic sections. More detailed analysis of these time series can be found in Banas et al.
2002 and Siegel et al. 2002. NCAR NCEP six-hourly winds from the Reanalysis project (Kalnay et al. 1996) were
obtained at 2.5 degree intervals and interpolated to Grays Harbor and Coos Bay mid shelf locations. These data are
provided by the NOAA-CIRES Climate Diagnostics Center, Boulder, Colorado at http://www.cdc.noaa.gov/. These
winds are generated from an atmospheric model that is primarily driven by atmospheric pressure but also include
data assimilated from both coastal buoys and land stations. The NCEP winds, being a spatial average, provide a
more accurate representation of alongshore gradients in wind than in situ buoys, which are frequently biased by
cross-shelf wind structure. Data from NDBC Buoy 46029 (the “Columbia River buoy”) were used in Figure 15a
(only).
Data were edited for outliers. Hydrographic section data were used to validate data from the moored arrays.
For subtidal time series, data were filtered with a Butterworth low pass filter and smoothed to hourly intervals.
c) Links to the coastal ocean
Response to upwelling/downwelling
As suggested above, an account of event-scale and seasonal variations in ocean-estuary coupling is critical
to a description of the physical dynamics of Northwest coast estuaries as well as their ecosystems. In this section we
Hickey and Banas 17
review recent analyses and observations of time-dependent coupling mechanisms, and present new data comparing
the responses of the PNCERS estuaries to ocean forcing.
The properties of the ocean water presented at the mouth of the estuary are governed by whether upwelling
or downwelling is occurring along the coast at that time (Hickey et al. 2002). During upwelling, surface waters move
offshore and cold, saltier, nutrient-rich water is moved upward within a few km of the coast; phytoplankton seed
stock are also upwelled into the euphotic zone and, fueled by the high nutrient level, begin to grow (Roegner et al.
2002). The growing phytoplankton move offshore as new seed stock is upwelled so that the highest biomass may be
situated some distance from the coastal wall and the mouths of the estuaries. During downwelling, warmer, fresher,
nutrient-depleted surface waters move inshore and downward, and offshore phytoplankton blooms likewise are
advected back to the coast.
Oceanic phytoplankton can enter a coastal estuary by two routes. During upwelling events, seed stock can
be pulled into the estuary, where a local bloom is fueled by the high nutrients brought in with the upwelled water (de-
Angelis and Gordon 1985). During downwelling events, phytoplankton from a prior offshore bloom can be pulled
directly into the estuary (Roegner et al. 2002). This biomass, although nutrient-poor and declining rather than
growing, may provide a direct food source to secondary production, particularly near the mouth of the estuary. In the
Pacific Northwest, transitions between upwelling and downwelling occur at 2-10 day intervals (Hickey 1989), and so
the ocean end-member of estuarine water properties can change significantly over just a few tidal cycles.
In Willapa Bay both upwelling and downwelling water presented at the mouth of the estuary have been
observed to travel up-estuary in the lower water column at a rate on the order of 10 km d-1, modifying the
gravitational circulation of the estuary as it passes (Hickey et al. 2002). These modulations of circulation and water
properties lag local wind stress fluctuations (hence, upwelling or downwelling) by more than a day (Fig. 12). This
mode of up-estuary propagation is consistent with the suggestion by Duxbury (1979) that modulation of the
gravitational circulation by upwelling and downwelling is responsible for increased mean flushing rates in summer
months in Grays Harbor. Such a baroclinic coupling between ocean and estuary is schematized in Fig. 13, with
values taken from typical early-summer conditions in Willapa Bay. During upwelling events, high ocean salinities
increase the along-channel salinity gradient and hence the magnitude of baroclinic exchange (Hansen and Rattray
1965, Monteiro and Largier 1999); during downwelling events, the salinity contrast between ocean and estuary, and
hence the strength of the exchange flow, are reduced.
Hickey and Banas 18
At the same time, the propagation of oceanic signals into the estuary appears to have, in addition to this
baroclinic, density-driven component, a significant diffusive, primarily barotropic, density-independent component.
In the long-term average in Willapa Bay, oceanic signals appear to propagate upstream in the surface layer (opposing
the mean gravitational circulation) at a rate similar to that of their propagation in the lower layer. The strength and
along-channel profile of this diffusive process suggest lateral stirring by wide tidal residual eddies tied to bathymetry
(Banas et al. 2002). Lateral wind-driven circulations (e.g., Wang 1979, Geyer 1997) may also be important, but as
yet have not been quantified.
The effectiveness of ocean-estuary exchange by tidal stirring depends not only upon processes within the
estuary, but also upon net transport and mixing on the shelf within a tidal excursion of the mouth. Shelf processes on
the tidal time scale presumably determine, largely, the fraction of an ebb tidal prism that does not simply re-enter the
estuary on the following flood tide (the "tidal exchange ratio"). From measurements of horizontal tidal diffusivity
Banas et al. (2002) estimate a tidal exchange ratio ≥ 0.6 for Willapa Bay as a whole, with the possibility of much
lower exchange ratios across cross-sections in the landward reaches of the bay. The exchange ratio at the mouth is at
the upper limit of the range reported by Dyer (1973), and thus is consistent with the active, highly advective inner-
shelf environment on this coast.
The studies reviewed above demonstrate that both baroclinic, density-driven exchange and diffusive, tide-
or wind-driven exchange can contribute significantly to the flushing of Northwest estuaries. Whichever of these
exchange mechanisms dominates determines the overall rate of ocean-estuary exchange, and the spatial pathways
along which oceanic nutrients, phytoplankton, and planktonic larvae enter an estuary on tidal or subtidal time scales.
In general, tide- and wind-driven mechanisms are expected to dominate in small, shallow, well-mixed estuaries, and
density-driven exchange is expected to dominate in deeper, partially stratified systems (e.g., Hansen and Rattray
1966). The Northwest coast estuaries span both of these broad categories, and we can expect the relative role of
density-driven and density-independent mechanisms to vary significantly between systems, and over time in a single
system.
On seasonal time scales, Willapa Bay, for example, appears to vary between river-controlled and tide-
controlled flushing. Banas et al. (2002) show that river-controlled exchange dominates in the narrow, landward,
southeastern reach of Willapa during all but the lowest late-summer riverflows. At the same time, in the seaward
reach of the estuary, flushing by tidal stirring dominates during all but the largest winter storms, even when
stratification is sustained at several psu. Thus tidal stirring appears to control the nutrient and biomass budget for
Hickey and Banas 19
most of the volume of Willapa during the spring-to-summer growing season. The morphological and forcing
properties that place Willapa at this transition point between river- and tide-controlled flushing are not particular to
Willapa, but rather general characteristics of Northwest coast estuaries: large tidal prisms, complex bathymetry, and
highly variable riverflow that is frequently large during winter but close to zero during much of the growing season.
Alongshore correlation between estuaries
As discussed in Section 2, wind-driven coastal processes in summer have scales of several hundred
kilometers. Thus, if the estuarine-ocean coupling processes described above are uniformly valid, we might expect
water properties in many PNW estuaries to vary coherently. The PNCERS data confirm that this is indeed the case:
time series of temperature data collected simultaneously in Grays Harbor, Willapa Bay, and Coos Bay demonstrate
that in the spring-fall growing season all three estuaries, which span 400 km of the Northwest coast are highly
coherent (Fig. 14a). Coherence is highest in late summer-early fall (day 201-271, r = 0.95 between Grays and
Willapa, 0.91 between Coos and Grays). In spring-early summer (day 130-200) a few peaks occur in the northern
estuaries that do not occur in the southern estuary, reducing coherence (r = 0.74). The spring-summer period is
further analyzed in the next three panels of Figure 14, showing temperature (b), alongshore wind (c) and salinity (d).
Comparison of temperature and wind illustrates the general response to upwelling and downwelling favorable winds,
with the wind/property lag (1.5 d at both Grays Harbor and Coos Bay) discussed in the preceding section. Significant
differences between the northern and southern estuaries are observed. For example, from day 165-175 temperature
decreases in Coos Bay while it increases in Grays Harbor (Fig. 14b). Comparison with alongshore wind (Fig. 14c)
demonstrates that this difference is caused by the fact that downwelling favorable winds are stronger during this
period near Grays Harbor than near Coos Bay. Several other similar examples of alongshore differences in estuary
water properties caused by alongshore wind differences can be seen in the records.
Another example is illustrated in salinity records from northern and southern estuaries from the same period
(Fig. 14d). In the southern estuary (Coos Bay) the upwelling signal is of significantly longer duration than in Grays
Harbor. Note that the greater salinity range and overall lower salinity at the northern estuary is consistent with the
generally lower regional salinity due to the proximity of the Columbia plume (see next section).
Columbia River plume intrusions
Hickey and Banas 20
Although the estuaries of Washington and Oregon generally respond to ocean forcing coherently as
discussed above, the Columbia River plume can cause major asymmetries between the estuaries. Since the plume
moves offshore when it flows southward past Oregon during periods of upwelling-favorable winds, it does not
impinge upon most Oregon estuaries directly. When the plume flows north under downwelling winds, however, it
fills the nearshore water column north of the river mouth past the depth of the estuary mouths (Garcia-Berdeal et al.
2002, Roegner et al. 2002). The plume may also impact estuaries on the northern Oregon coast during downwelling,
when the southwest-tending plume formed under the preceding upwelling conditions and seasonal southward
ambient flow moves shoreward. However, mixing during the downwelling event would result in much less density
contrast than off the Washington coast, where the plume is relatively new and thus fresher. The effect of the plume
on the estuaries is most dramatic and sustained in late spring and early summer, when local riverflow has slackened
but the Columbia is still running high with snowmelt.
Lower water column salinities from moorings inside the mouths of Willapa Bay and Grays Harbor during
April and May 2000 are shown in Fig. 15. For each station, the along-channel salinity gradient has also been
calculated as a subtidal time series, by dividing the difference between high- and low-water salinities by the tidal
excursion for each semidiurnal tidal cycle, and then filtering the resulting discrete series. This method, also used by
Banas et al. (2002), takes advantage of the fact that each station effectively samples ~ 15 km of the channel through
tidal advection. This allows us to calculate along channel gradients without requiring pairs of stations to obtain
differences. An upwelling event, which brings ~32 psu water into the estuaries and produces strong along-channel
gradients (on April 19, ~5 psu over one tidal excursion), is followed by a plume intrusion, indicated by a dramatic
decrease in salinity and weak along-channel salinity gradients. When downwelling-favorable winds slacken after
~April 27, salinity and the along-channel salinity gradient increase again. The five-month wind time series shown in
Figure 15a suggests that this intermittent alternation of upwelling and plume intrusion continues from late winter
through early summer.
During the onset of plume intrusions the along-channel salinity gradient in the estuary can reverse for
sustained periods. In Figure 15b, for example, as the plume intrusion intensifies during the period April 20-28,
salinity at the Willapa Bay mooring at high slack water (indicated by dots) is generally lower than the subtidal
average, indicating that each flood tide is bringing somewhat fresher water into the estuary. This reversal of the
expected gradient between mid-estuary and ocean water is illustrated in a CTD transect along the main channel of
Hickey and Banas 21
Willapa Bay on May 3, 2000, during the recovery from the plume intrusion (Fig. 16a). Salinity increases downstream
from the head to > 21.8 psu, drops to < 21.4 psu, and then increases again within one tidal excursion of the mouth.
Vertical gradients weaken during plume intrusions along with the longitudinal gradients. The vertical
salinity difference in the interior of the estuary in the May 3 transect is on the order of 0.1 psu. In comparison, a
transect on May 30 during the onset of an upwelling event after a period of intermittent winds (Fig. 16b) shows
vertical salinity differences ~2-4 psu within a tidal excursion of the mouth. During a plume intrusion, the reduced
salinity contrast between the river and ocean end-members of the estuary presumably weakens baroclinic pressure
gradients and thus stratification to the point where vertical shear dispersion can completely homogenize the water
column. Thus in contrast to input of freshwater from the local rivers, which tends to increase stratification and
gravitational exchange (Banas et al. 2002), input of freshwater from the Columbia River via the coastal ocean tends
to produce near-complete mixing in Washington estuaries.
As described in the preceding section, downwelling conditions tend to reduce estuarine salinity gradients
even in the absence of plume intrusions (Hickey et al. 2002), though to a much lesser extent. The effect of the
Columbia River plume, then, is to greatly intensify the contrast between spring and summer upwelling and
downwelling conditions in the Washington estuaries in comparison with Oregon estuaries. This asymmetry between
the two coasts would likely be observed not just on the event scale, but on interannual scales as well. Following wet
(La-Niña-like) winters like 1998-1999, but not following dry (El-Niño-like) winters like 1997-1998, sustained plume
intrusions would be expected in the Washington estuaries during May and June. Anecdotal evidence suggests that in
years prior to the man-made reduction in spring freshets from the Columbia, the surface of Willapa Bay would freeze
sufficiently to support walking, suggesting very low wintertime salinities (Proc. Oyster Grow. Conf. 2002).
d) Spatial variability in the intertidal zone
We have argued that much of the physical forcing important to estuarine productivity is coherent over tens
or hundreds of kilometers on this coast, and that the response of the coastal estuaries to this forcing may be coherent
and generalizable on this scale as well. At the same time, pervasive, significant variation in currents and hydrography
is possible on much smaller scales—as short as ~ 100 m—in estuaries with complex bathymetry, particularly in very
shallow regions, which often are most important biologically. These small-scale variations, which can be thought of
Hickey and Banas 22
as creating estuarine microenvironments, easily confound attempts to generalize from measurements that do not
integrate over larger scales.
In this section we use new data to describe the two mechanisms of lateral variability best resolved by tidal
scale observations in the Washington estuaries: 1) direct solar heating of bank water, and 2) the creation of persistent
lateral gradients by tidal advection. A full account of the transverse structure of these estuaries—which must
consider competition and interaction between tidal currents, density-driven flows, rotational effects, and wind-driven
circulations, all of which are shaped by bathymetry (e.g., Friedrichs et al. 1992, Valle-Levinson and O'Donnell
1996)—is beyond the scope of available data.
Solar heating
Coordinated longitudinal (along-channel) and transverse (bank-to-channel-to-bank) CTD transects were
obtained in Willapa Bay and Grays Harbor during the summers of 1999 and 2000. These observations frequently
suggest solar heating of water on shallow intertidal flats: either direct heating of the water at high tide, or transfer to
the water of heat stored in the mud flats themselves from insolation at low tide. Consider, for example, a late-
afternoon, early-flood transect along the main channel of Grays Harbor during a period of fair weather in June 1999
(Fig. 17). The warmest water in the channel is associated with neither the ocean nor the river end-member, but rather
appears near the surface over a broad middle reach of the channel. CTD casts along this transect were separated by
~4 km, and therefore the spatial structure of this warm water may be patchier than contouring between casts allows.
We interpret this signal as evidence of water warmed during the midday high tide that has circulated back into the
main channel on the following ebb. A temperature-salinity (T-S) diagram of this transect (Fig. 17c) shows clearly
that this signal represents warming of water at intermediate salinity, and effectively constitutes a third mixing end-
member, toward which the T-S profile of the channel is inflected. Furthermore, transverse, channel-to-shoal surveys
on the day of the along-channel transect and over the next four days locate a similar warm water mass in depths < 5
m at higher stages of the tide (dots in Figure 17c).
Surveys in Willapa Bay from June and July 2000 (Fig. 18) show similar results: warmest temperatures on
banks in the interior of the estuary, inflection of the main-channel T-S profile that lifts intermediate water above the
mixing line between the ocean and river end-members. The warmest points in the June 2000 survey, more than 4°C
Hickey and Banas 23
warmer than main-channel water of the same salinity, represent the shallowest water sampled, water < 0.5 m deep
sampled by foot with a hand-held meter.
Note that since the fair-weather events which bring increased insolation also bring cold, upwelled water, an
estuary's response to direct heating may be masked on the event scale, and better resolved by an integration over
many events. In Willapa Bay, where time series of along-channel transects exist, the inflection of the T-S profile
tends to increase as the summer proceeds, though not monotonically (not shown).
Differential tidal advection
Not all bank-to-channel hydrographic variations result from solar heating or other transformation of water
properties. Consider the along- and cross-channel flood-tide transects from July 1999 in Willapa Bay shown in Fig.
19. The along-channel salinity gradient is ~5 psu over one tidal excursion (15 km); across a shallow, narrow bank
adjacent to the main channel during late flood, the salinity gradient is ~ 4 psu over only 1.3 km. Huzzey (1988)
likewise found that in the York River, which like Willapa consists of a deep central channel flanked by shoals, the
freshest water in a cross-section at high slack water was located on the banks. A T-S diagram of the July 1999
transects (Fig. 19c) shows that the bank and channel water masses, unlike those shown in Figs. 17 and 18, are
indistinguishable. The lateral variation in salinity and temperature thus must have arisen from advective
rearrangement, not transformation, of main-channel water in the intertidal zone.
Large lateral gradients can arise solely from differential advection by tidal currents (Huzzey and Brubaker
1988, O'Donnell 1993); i.e., the fact that on a shallow bank tidal motion is slowed by friction so that a given flood or
ebb moves water parcels farther longitudinally in a channel than on an adjacent shoal. This shearing of the flow
effectively transfers the along-channel gradient over one tidal excursion, or some fraction thereof, into a cross-
channel gradient. In support of this explanation for the lateral variation seen in Willapa in July 1999, repeated
channel-to-bank surveys in the same location have shown that the transverse gradient there at high water follows the
along-channel variation. On November 1-2, 1999, for example, the along-channel salinity gradient in the central
reach of the estuary was much weaker than that shown in Fig. 19, only ~ 0.5 psu over one tidal excursion (Banas et
al. 2002), and the salinity variation over the bank was likewise ~ 0.5 psu.
The differential-advective effect would be expected to be strongest on the shallowest banks (like that shown
in Figure 19b) where the effect of friction is presumably greatest, and less important on deeper, subtidal shoals. Such
Hickey and Banas 24
lateral structure in tidal advection may have important local, biological consequences. For example, sessile
organisms in a shallow region with strong lateral gradients may experience mean temperatures or rates of nutrient or
food supply appreciably different—more like conditions a large fraction of a tidal excursion up-estuary—than
organisms in deeper water a short distance away. At the same time, differential tidal advection may contribute to
overall estuarine flushing if these lateral shears are a lateral-dispersion mechanism similar to the models of "tidal
trapping" reviewed by Fischer (1976).
5. Discussion
New information on oceanographic processes has been synthesized to provide a framework for better
understanding some aspects of ecosystem variability. In addition new data from three coastal estuaries have been
presented to identify dominant processes and scales of variability, and to illustrate how the estuaries interact with the
coastal ocean.
From an ecological point of view the study presents the important idea that in spite of the fact that
upwelling wind stress is as much as a factor of two weaker off Washington than off Oregon, productivity in general
is higher off Washington (with the exception, at times, of Heceta Bank in southern Oregon). Two possible
mechanisms were discussed: the difference in shelf structure and the presence of submarine canyons on the
Washington coast. An additional possibility to account for these differences is the role of micronutrients. The few
primary productivity measurements in the literature illustrate enhanced productivity in the Columbia plume
(Anderson 1972), a rich source of iron. Moreover, because of the Columbia, iron-rich sediments overlie the
Washington shelf in the mid shelf silt deposit (Nittrouer 1978). These sediments could provide iron to upwelling
phytoplankton via contact in the bottom boundary layer. Off Oregon, plume sediments are much reduced.
Furthermore, because of the steeper shelf, upwelling water may cross the shelf more in the interior and less in the
bottom boundary layer (Allen et al. 1995). Unfortunately, existing data are not sufficient to separate the several
plausible mechanisms that could produce the observed alongshore productivity differences.
Within the coastal estuaries, the physical questions of primary ecological importance involve ocean-estuary
exchange. The construction of budgets for nutrients or phytoplankton in these estuaries requires better knowledge of
the bulk rate of exchange, and its variation on seasonal and event time scales. Likewise, delineation of the pathways
Hickey and Banas 25
of larval recruitment into the estuaries requires that we better understand variations in ocean-estuary matter transfer
on event and tidal time scales, as well as the fine-grained spatial structure of residual circulations on these scales.
Bulk coupling between ocean and estuary is necessarily central to the dynamics of any small embayment;
but its importance is perhaps amplified in PNW estuaries for two reasons. First, the ocean rather than local rivers is
the dominant source of nutrients and biomass along this coast. Second, oceanic water properties are extremely
variable on the scale of the residence time or adjustment time of the estuaries themselves (i.e., on the event scale).
This coincidence of time scales makes the estuaries highly variable and unsteady themselves, far more so, during
summer, than fluctuations in local river input would force on their own.
The example of Willapa Bay suggests that we may, in fact, be able to neglect the influence of river-driven
circulations on the overall flushing rate of these estuaries during some portion of the year (roughly speaking, the
growing season) and on sufficiently long time scales (several events or a few weeks). This is a very approximate
model of estuarine dynamics during this time period, but a powerful simplification, if tenable. It should be noted that
even if one can parameterize estuary flushing in terms of tide- or wind-driven lateral stirring as we are suggesting,
freshwater influences may still be important in determining the distribution of heat, salt, and other tracers within the
estuary.
The data collected to date, while providing a framework for beginning to understand the processes in the
important estuaries of the PNW, have also demonstrated the complexities that they include. In our ongoing research,
a three dimensional numerical model is being used to separate the dominant forcing mechanisms, Lagrangain
pathways and estuary-ocean fluxes. The PNW coast estuaries constitute a fruitful set for dynamical or ecological
comparison: they have enough commonalities (similar tidal forcing, similar riverflow patterns, event-scale
synchrony) to be usefully compared, and at the same time enough diversity (in overall size and cross-sectional shape,
relative riverflow magnitude, and relation to the Columbia plume) to form an interesting natural experiment.
Acknowledgments
Data collection was supported by the Pacific Northwest Ecosystem Research Study (PNCERS) (grant #
NA76RG0485 and NA96OP0238 from the Coastal Ocean program of the National Oceanic and Atmospheric
Administration). Analysis was supported by PNCERS, Washington Sea Grant (grant # NA16RG1044-R/ES-42 and
Hickey and Banas 26
NA16RG1044-R/F-137) and by a grant (#OCE-0001034) to B. Hickey from the National Science Foundation as part
of GLOBEC. This is contribution number XXXX of the U.S. GLOBEC program, jointly funded by the National
Science Foundation and National Oceanic and Atmospheric Administration. Last, we would like to thank Dr. Jan
Newton and Mr. Eric Siegel of the Washington State Department of Ecology for graciously providing additional
CTD data for Willapa Bay and Dr. John Klinck and M. Dinneman for providing modeled circulation to illustrate the
canyon effects.
.
Hickey and Banas 27
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Malamud-Roam, K.P. 2000. Muted tidal regimes in marshes of the San Francisco estuary: theory and implications
for ecological restoration. Ph.D. Dissertation, University of California at Berkeley. Berkeley, CA 94720.
Grays Willapa Yaquina Coos Columbia San Fran- South Plum I. Naragansett Chesapeake James Harbora Baya Baya Baya Rivera cisco Baya S.F. Bayb Soundc Baya Baya Rivera
area at MSL 150 240 13 34 550 1170 480 7.2–15 430 9900 610 AMSL (km2)
mean depth 4.3 3.2 2.6 4.0 7.3 6.8 4.4 2.3 10 8.5 5.2 H (m)
volume below MSL 0.64 0.76 0.034 0.14 4.0 8.0 2.1 ~0.016 4.3 84 3.2 V = H · AMSL (km3) ___________________________________________________________________________________________________________ mean tidal range at 2.1 1.9 1.8 1.7 1.7 1.3 1.4 2.6 0.9 0.8 0.8 mouth (m) intertidal area — 55d 47e 47e — — — — — — — (% area at MHW) tidal prism volume 46 50 52 31 14 16a–27b 37 ~50 10 2.0 8.6 (% volume at MHW) ___________________________________________________________________________________________________________ drainage area 7.0 2.9 0.66 1.5 670 120 — 0.58 4.6 180 26 (1000 km2) monthly-mean riverflow R (m3 s-1) lowest mo. flow 56 17 ~0 2.8 4200 330 — ~0 34 950 150 highest mo. flow 880 390 68 190 10000 1800 — 10 170 4200 600 river-filling time V/R (d) lowest mo. flow 100 500 long 600 10 300 — long 2000 1000 200 highest mo. flow 8 20 6 8 5 50 — 20 300 200 60 ____________________ aNOAA 1985; bMalamud-Roam 2000; cJay et al. 1997; dAndrews 1965; ePercy et al. 1974. Table 1. Indices of morphology, tidal forcing, and river input for the four PNCERS estuaries and seven others on the U.S. East and West Coasts.
Hickey and Banas 36
Figure Captions
Figure 1. Schematic of the California Current System. Adapted from Hickey and Royer 2001.
Figure 2. Schematic of wind-driven coastal circulation in the PNW. The cartoon illustrates the offshore (onshore)-
directed surface currents that occur in response to an upwelling (downwelling)-favorable wind stress and upwelling
(downwelling) along the coast. Freshwater flows from coastal estuaries and from the Strait of Juan de Fuca are
illustrated with darker shading. The location of a persistent summertime mesoscale feature (the Juan de Fuca Eddy)
is also shown.
Figure 3. Satellite-derived sea surface temperature in the PNW. The figure illustrates enhanced upwelling along the
Washington coast and enhanced upwelling in the lee of the northern headland along the Washington coast (shown as
white arrow). The Juan de Fuca eddy is readily apparent as a coldwater feature opposite the strait of Juan de Fuca.
Figure 4. Seasonal variation of selected nutrients in the PNWat mid shelf off Washington and off Oregon. Adapted
from Landry et al. 1989.
Figure 5. Annual cycle of the vertically-integrated chlorophyll over the Washington and Oregon shelves. Adapted
from Landry et al. 1989. The data illustrate the typically higher chlorophyll on the Washington shelf.
Figure 6. Topography of the PNW illustrating important canyons and banks.
Figure 7. Particulate domoic acid in pseudo-nitzschia species on the Pacific West Coast in 1998. Maximum
concentrations of domoic acid and toxic species are indicated to the right of each area of high toxin. Each of these
areas is associated with relatively retentive circulation patterns. Adapted from Trainer et al. 2000.
Figure 8. Modeled circulation at depths ranging from 50 m to 600 m showing cyclonic eddies over and within two
submarine canyons off the Washington coast (Dinneman and Klinck pers. comm.). Note that at 50 m the flow is
Hickey and Banas 37
relatively undisturbed by the canyon topography. The circulation was forced by an upwellling-favorable wind stress
with a magnitude typical for this region.
Figure 9. Modeled response of the Columbia plume in summer to changes in wind direction. The figure illustrates
the evolution of surface salinity (psu) for southward ambient flow conditions in response to 6 days of downwelling
favorable winds, followed by 6 days of upwelling favorable winds at (a) 13 days, (b) 15 days, (c) 16 days, (d) 19
days, (e) 21 days and (f) 25 days with a southward ambient flow of 10 cm s-1. Winds change direction after 19 days
immediately after (d). The distance between tick marks is 20 km. Adapted from Garcia-Berdeal et al. 2002.
Figure 10. Modeled velocity structure of the Columbia plume illustrating potential retentive areas (eddy-like
features) and frontal jets associated with the river plume. Surface salinity (psu) contours and surface velocity vectors
(m s-1) at t = 28 days for (a) northward ambient flow of 10 cm s-1 and (b) southward ambient flow of 10 cm s-1.
River discharge for both cases is 7000 m3 s-1. Adapted from Garcia-Berdeal et al. 2002.
Figure 11. Map of the Pacific Northwest coast from Washington to Northern California, showing the location of the
four PNCERS estuaries and other estuaries in the region. Maps of Grays Harbor, Willapa Bay, and Coos Bay, with
the locations of estuarine and offshore moorings are also shown.
Figure 12. Time series of temperature at selected sites in Willapa estuary illustrating up-estuary propagation of an
upwelling-driven signal. Time series of salinity in the estuary and alongshore wind on the nearby coast illustrate the
response of the estuary water properties to coastal upwelling events-high salinity during periods of upwelling-
favorable winds and low salinity during periods of downwelling-favorable winds, with a~1.5 day lag between wind
and estuary salinity.
Figure 13. Schematic illustrating baroclinic coupling between the coastal ocean and a coastal plain estuary in the
Pacific Northwest during upwelling and downwelling events for a low riverflow, summer period in an Eastern
Boundary System. From Hickey et al. 2002.
Figure 14. (a) Time series of temperature in three estuaries in the Pacific Northwest, illustrating simultaneous
response to large scale upwelling/downwelling along the open coast during spring-fall, 1999. All data are from the
Hickey and Banas 38
lower water column and from stations near the mouth of each estuary (see locations in Figure 11). Solid line along
the x-axis indicates interval expanded in (b,c,d). (b) Time series of temperature in Grays Harbor and Coos Bay.
Arrows indicate downwelling events (warmer water) prominent only in the northern estuaries. (c) Time series of
alongshore wind at latitudes close to Grays Harbor and Coos Bay. Arrows illustrate wind events that cause
downwelling seen in panel c. (d) Time series of salinity for the same period, illustrating effects of more persistent
upwelling favorable winds at more southern locations.
Figure 15. (a) Time series of the north-south component of nearshore wind from late winter to early summer, 2000.
The dates of the two CTD transects of Willapa Bay shown in Figure 16 are indicated. (b) Salinity and (c) the local
along-channel salinity gradient near the mouths of Willapa Bay and Grays Harbor during a three-week period Apr-
May, 2000, showing a brief upwelling event, an intrusion of the Columbia River plume, and a recovery from that
intrusion. In (b), both 30-min and subtidal (48-hr-Butterworth-filtered) data are shown. Dots mark times of high
slack water in Willapa Bay. In (c), the difference between high-slack and low-slack salinity divided by the tidal
excursion for each semidiurnal tidal cycle has been filtered as above to provide a subtidal, single-station time series
of the along-channel salinity gradient.
Figure 16. Salinity from CTD transects along the main channel of Willapa Bay on (a) May 3, 2000, near the end of a
Columbia River plume intrusion, and (b) May 30, 2000, during strong upwelling event that replaces plume water
(~21.5 psu) with much saltier water (≥ 29 psu). A reversal of the along-channel salinity gradient is marked in (a).
Triangles at the top of the salinity sections give the location of CTD casts. Tidal stage and transect route are also
given for each section. These surveys were made within a five-month wind time series in Figure 15a.
Figure 17. (a,b) Temperature and salinity from a CTD transect along the main channel of Grays Harbor June 11,
1999 during a time of strong solar heating. Triangles at the top of the sections give the location of CTD casts. (c)
Temperature-salinity profile of the along-channel transect (lines) and CTD casts on shoals adjacent to the channel
June 11-15 (dots). Location and tidal stage of bank and channel surveys are also shown. Dots on inset maps indicate
location of bank stations.
Hickey and Banas 39
Figure 18. Temperature-salinity diagrams for surveys of Willapa Bay during (a) June and (b) July 2000, showing the
hydrographic signature of direct solar heating. Line segments represent CTD casts within the main channels of the
estuary; dots represent water on banks with depths < 5 m.
Figure 19. Salinity from CTD transects on July 14, 1999 (a) along the main channel of Willapa Bay and (b) from the
channel to shore across a shallow, narrow bank. Vertical line in (a) near 18 km marks the location of the cross bank
transect in (b). Location and tidal stage are shown; triangles at the top of the sections give the location of CTD casts.
The nearly identical temperature-salinity profiles of the along channel and cross channel transects are confirmed with
a T-S diagram in (c).
Figure 1.
Storms
GraysHarbor
WillapaBay
Columbiaplume
downwelling water
fresher
Near-SurfaceShelf Flow
◆warmer◆less saline◆nutrient- reduced
Fair Weather
GraysHarbor
WillapaBay
ColumbiaRiver
upwelling water◆colder◆more saline◆nutrient- enriched
winter plume
summer plume
Near-SurfaceShelf Flow
colder
warmer
J. de FucaEddy
Juan de Fuca Strait
Juan de Fuca Strait
ColumbiaRiver
N N
Figure 2.
Figure 3.
Figure 4.
Figure 5.
Figure 6.
Figure 7.
50m, Day 10.0
46.0°N
46.4°
46.8°
100m, Day 10.0
400m, Day 10.0
46.0°N
46.4°
46.8°
125° W 124°
600m, Day 10.0
125° W 124°
150m, Day 10.0
46.0°N
46.4°
46.8°
250m, Day 10.0
GuideCanyon
WillapaCanyon
AstoriaCanyon
20 cm s -1 20 cm s -1
20 cm s -120 cm s -1
20 cm s -1 20 cm s -1
NN
NN
NN
Figure 8.
Figure 9.
Figure 10.
124°40' 124°20'
43°20'
43°30'
CBOS
CR
TD
CoosBay
124°20' 124°00' 123°40'46°20'
46°30'
46°40'
46°50'
47°00'
GHOS
GH
W3
W6
WillapaBay
Grays Harbor
Grays Harbor
Willapa Bay
Coos Bay
S. F. Bay
Columbia R.
PugetSound
Yaquina Bay
N
43°10'
Figure 11.
1 0
1 2
1 4
1 6
1 8
2 0
Tem
pera
ture
(OC
)
W1W3
W4W5
Along Estuary Temperature
24
26
28
30
32
34
-1.5
-1
-0.5
0
0.5
1
1.5
Sal
inity
(ps
u) A
djus
ted
by 1
.5 d
Time (calendar day)
r = -0.64S at W1τ
Wind Stress vs. Near Mouth Salinity
150 170 190 210160 180 200 220
up-estuary propagation ~12 cm s-1
-2W
ind
Str
ess
(dyn
es c
m)
■
■
■
■W1W3
W4
W5
Figure 12.
Figure 13.
8
10
12
14
16
18
100 150 200 250 300
Tem
pera
ture
(OC
) Willapa
Coos
Grays
June Sept
8
9
10
11
12
13
14
15
130 140 150 160 170 180 190 200
Tem
pera
ture
(OC
)
Coos
Grays
-15
-10
-5
0
5
10
130 140 150 160 170 180 190 200
Alo
ngsh
ore
Win
d (m
s-1 ) Grays
Coos
Downwelling
Upwelling
20
25
30
35
130 140 150 160 170 180 190 200
Sal
inity
(ps
u)
Coos
GraysUpwelling
Calendar Day 1999
downwelling
Longer upwelling in Coos
Figure 14.
04/16 04/20 04/24 04/28 05/02 05/06
20
24
28
32
Salinity (psu)
04/16 04/20 04/24 04/28 05/02 05/06-0.2
0
0.2
0.4
Along-channel salinity gradient (psu/km)
-10
0
10
MayMar Jul 2000Feb Apr Jun
N-S wind (m s-1)
MA
Y 3
MA
Y 3
0
from the south (downwelling-favorable)
from the north (upwelling-favorable)
Grays Harbor
Willapa Bayat high slack
subtidal
Grays Harbor
Willapa Bay
(a)
(b)
(c)
UPWELLING PLUME INTRUSION WIND RELAXATION
Figure 15.
0 10 20 30
0
10
20
Distance from mouth (km)
Dep
th (
m)
22
21.5
21.4
21.4 21.5 21.521.8
21
2019 18
0 10 20 30 40
0
10
20
Distance from mouth (km)
2928
2726
25
23 22
21.5
22
21.5
20
1918
17Salinity (psu)
21.8
gradient reversal
Tidal height at W6(MLLW; m)
21
240
2
4
0h 12h 0hLocal Time
Dep
th (
m)
40
0
2
4
0h 12h 0hLocal Time
10 km
Willapa BayMay 3, 2000
May 30, 2000
(a)
(b)
20 32
Salinity (psu)
16 24 280
Figure 16.
5 10 15 20 25 30
10
12
14
16
18
20
0 10 20 30
0
4
8
12
16
20
Distance from Mouth (km)
Dep
th (
m)
26
24
2220
18 16 14
12
10 86
0 10 20 30
0
4
8
12
16
20
Distance from Mouth (km)
Dep
th (
m)
12
13
13
1415
16
Tem
pera
ture
(°C
)
Salinity (psu)
channel
banks(depth < 5 m)
(a) (b)
(c)
20 32
Salinity (psu)
16 24 280
0
2
4
0
2
4
0
2
4
0h12h0h
0h12h0h
0h12h0h
Jun 11
Jun 13
Jun 15
Tid
al h
eigh
t at m
id-e
stua
ry (
MLL
W; m
)
channel
12 1810 14 160
Temperature (°C)
10 km
Grays Harbor
banks
Figure 17.
22 24 26 28 30 32
10
12
14
16
18
20July 24-25, 2000
Tem
pera
ture
(°C
)
Salinity (psu)
channel
banks(depth < 5 m)
22 24 26 28 30 32
10
12
14
16
18
20
10 km
June 26-28, 2000
Tem
pera
ture
(°C
)
Salinity (psu)(a) (b)
Figure 18.
22 24 26 28 30 32
10
12
14
16
18
20
10 20 30
0
10
20
Distance from mouth (km)
Dep
th (
m)
24252627
2829
30
3131.5
Salinity (psu)
0
2
4
0h 12h 0h
Tidal height at W6(MLLW; m)
0
2
4
0h 12h 0h
Tidal height at W6(MLLW; m)
-1600 -1200 -800 -400 0
0
10
Distance from eastern shore (m)
Dep
th (
m)
2526272829
29.55
Salinity (psu)
20 3216 24 280T
empe
ratu
re (
°C)
Salinity (psu)
channel (a)
bank (b)
(a)
(b)
(c)
Jul 14, 1999
1 km
10 km
Local time
Figure 19.
