9.13 Geomorphic Controls on Hyporheic Exchange Across Scales:
Watersheds to Particles
SM Wondzell, Olympia Forestry Sciences Laboratory, Olympia, WA, USA
MN Gooseff, Pennsylvania State University, University Park, PA, USA
r 2013 Elsevier Inc. All rights reserved.
9.13.1
9.13.2
9.13.2.1
9.13.2.2
9.13.2.2.1
9.13.2.2.2
9.13.2.3
9.13.2.3.1
9.13.2.3.2
9.13.2.3.3
9.13.2.3.4
9.13.2.3.5
9.13.2.4
9.13.2.5
9.13.3
9.13.3.1
9.13.3.2
9.13.4
References
Introduction
The Effect of Geomorphology on HEFs
The Whole Network to Segment Scale
The Reach Scale – Setting the Potential for Hyporheic Exchange
Losing and gaining reaches
Changes in saturated cross-sectional area
The Subreach to Channel-Unit Scale – Hydrostatic Processes
Step–pool and pool–riffle sequences
Meander bends and point bars
Back channels and floodplain spring brooks
Secondary channels and islands
Spatial heterogeneity in saturated hydraulic conductivity
The Bedform Scale – Hydrodynamic Processes
The Particle Scale – Turbulent Diffusion
Discussion
Multiple Features Acting in Concert
Change in Processes Driving HEF through the Stream Network
Conclusion
Glossary
Bedform streambed features, such as ripples, sand-waves,
or dunes, formed in streambed sediment through the
interaction between flowing water and sediment. Bedforms
can range in size from a few centimeters (ripples) to many
meters (dunes).
Channel unit relatively homogeneous areas of a channel
with characteristic substrate, depth, and flow pattern,
typically bounded by other channel areas with different
features. The length of channel units generally ranges
from one to a few times the wetted width of the stream.
Pools and riffles are common channel units in small
streams.
Hydrodynamic exchange (also known as pumping
exchange) exchange of stream water with the subsurface
driven by the velocity head component of the total head
gradient on the bed surface and exchange induced by
momentum gradients across beds and banks.
Hydrostatic exchange exchange of stream water with
the subsurface driven by static hydraulic gradients
that are determined by changes in water surface
elevation, spatial heterogeneity in saturated hydraulic
conductivity, or changes in the saturated cross-sectional
204
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205
206
206
208
209
209
210
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211
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212
213
213
213
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215
area of floodplain alluvium through which hyporheic flow
occurs.
Hyporheic exchange flow the flow of stream water from
the surface stream channel, through the streambed and into
the shallow unconfined aquifer beneath or adjacent to the
stream channel and back into the surface stream over
relatively short spatial (1-100s of meters) and temporal
(hours to weeks) scales.
Hyporheic zone the zone of saturated sediments
underlying the surface stream or in the floodplain adjacent to
the stream perfused with stream-source water that has recently
left the stream channel and will return to the stream channel
relatively quickly. Stream water in hyporheic flow paths may
mix with groundwater so that the relative proportion of
stream-source water in the hyporheic zone is highly variable,
ranging from 100% stream water to nearly 100% groundwater.
The extent of the hyporheic zone can be arbitrarily defined on
the basis of the proportion of stream water present in the
subsurface (e.g., 410%) or the residence time of stream water
(e.g., areas with residence times 24 h).
Stream reach a length of stream channel in which
topographic features and sequences of channel units are
relatively homogeneous. The length of stream reaches are
Wondzell, S.M., Gooseff, M.N., 2013. Geomorphic controls on hyporheic
exchange across scales: watersheds to particles. In: Shroder, J. (Editor in
Chief), Wohl, E. (Ed.), Treatise on Geomorphology. Academic Press, San
Diego, CA, vol. 9, Fluvial Geomorphology, pp. 203–218.
Treatise on Geomorphology, Volume 9
http://dx.doi.org/10.1016/B978-0-12-374739-6.00238-4
203
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Geomorphic Controls on Hyporheic Exchange Across Scales: Watersheds to Particles
generally many 10s of times the width of the wetted stream
channel.
Transient exchange the temporary movement of stream
water into stream banks due to short-term increases in
stream stage (i.e., bank storage processes due to changes in
hydrostatic head gradients between stream and lateral
riparian aquifer).
Abstract
We examined the relationship between fluvial geomorphology and hyporheic exchange flows. We use geomorphology as a
framework to understand hyporheic process and how these processes change with location within a stream network, and,
over time, in response to changes in stream discharge and catchment wetness. We focus primarily on hydrostatic and
hydrodynamic processes – the processes where linkages to fluvial geomorphology are most direct. Hydrostatic processes
result from morphologic features that create elevational head gradients, whereas hydrodynamic processes result from the
interaction between stream flow and channel morphologic features. We provide examples of the specific morphologic
features that drive or enable hyporheic exchange and we examine how these processes interact in real stream networks to
create complex subsurface flow nets through the hyporheic zone.
9.13.1
Introduction
Hyporheic exchange flow (HEF) is the movement of stream
water from the surface channel into the subsurface and back to
the stream (Figure 1). Stream water in hyporheic flow paths
may mix with groundwater so that the relative proportion of
stream-source water in the hyporheic zone (HZ) is highly
variable, ranging from 100% stream water to nearly 100%
Pool−riffle−pool sequence
(a)
Riffle
Pool
Pool
(b)
Floodplain
Riffle
Active
channel
Subsurface
flow path
Wetted
channel
Figure 1 Idealized conceptual model of nested hyporheic flow paths
as influenced by step–pool or pool–riffle sequences. (a) Plan view
showing arcuate HEF flow paths through the adjacent floodplain
created by the change in the longitudinal gradient over the pool–riffle
sequence where the amount of HEF is proportional to the head
gradient. (b) Longitudinal section along the thalweg of the stream
showing the vertical component of HEF flows through the streambed.
groundwater. Also, the residence time distribution of stream
water in the hyporheic zone tends to be highly skewed, with
most of the stream water moving along short flow paths and
thus having short residence times (hours), but some water
either moving on long flow paths or encountering relatively
immobile regions having very extended residence times
(weeks to months, or longer). The boundaries of the hyporheic zone are arbitrary, commonly defined by the amount of
stream-source water present in the subsurface. Triska et al.
(1989) set a threshold of 10% stream-source water to define
the limits of the hyporheic zone so that regions with o10%
stream-source water were defined as groundwater. Alternatively, the extent of the hyporheic zone can be delimited by
water residence time, for example, the subsurface zone delineated by HEFs with residence times less than 24 h (the 24-h
hyporheic zone; Gooseff, 2010).
The objective of this chapter is to examine the relation
between geomorphology and hyporheic processes. The two
primary controls on hyporheic exchange are the gradients in
total head established along and across streambeds and the
hydraulic conductivity of the streambed and adjacent aquifer,
both of which are significantly influenced by geomorphology.
Total head (also known as potential) is the sum of pressure
head, elevation head, and velocity head. Pressure head represents height of a column of fluid to produce pressure. Velocity
head represents the vertical distance needed for the fluid to fall
freely (neglecting friction) to reach a particular velocity from
rest. Elevation head represents the potential energy of a fluid
particle in terms of its height from reference datum. Hydrostatic head is referred to as the sum of elevation and pressure
head. Groundwater tables in unconfined aquifers represent
the spatial gradients in hydrostatic head. A number of processes either drive or enable HEF, several of which are based
on changes in head gradients. We follow the organizational
structure presented by Käser et al. (2009), who divided these
processes into five distinct classes:
1. Transient exchange – the temporary movement of stream
water into stream banks due to short-term increases in
stream stage (i.e., bank storage processes due to changes in
Geomorphic Controls on Hyporheic Exchange Across Scales: Watersheds to Particles
2.
3.
4.
5.
hydrostatic head gradients between stream and lateral riparian aquifer; Lewandowski et al., 2009; Sawyer et al., 2009).
Turnover exchange – the trapping of stream water in the
streambed during times of significant bed mobility (Elliott
and Brooks, 1997b; Packman and Brooks, 2001).
Turbulent diffusion – exchange driven by slip velocity that is
created at the surface of the porous medium of the bed
where streamwise velocity vectors continue to propagate into
the surface layers of the bed (Packman and Bencala, 2000).
Hydrostatic-driven exchange – exchange driven by static
hydraulic gradients that are determined by changes in
water surface elevation (Harvey and Bencala, 1993), spatial
heterogeneity in saturated hydraulic conductivity, or
changes in the saturated cross-sectional area of floodplain
alluvium through which hyporheic flow occurs.
Hydrodynamic-driven exchange – exchange driven by the
velocity head component of the total head gradient on the
bed surface (i.e., pumping exchange; Elliott and Brooks,
1997a, 1997b) and exchange induced by momentum gradients across beds and banks.
These classes of HEF processes are coupled to geomorphic
processes in many ways. This is most obvious for hydrostatic
effects, which are directly dependent on channel and valleyfloor morphology and the depositional environment that
controls spatial heterogeneity in saturated hydraulic conductivity (K). However, turnover of streambed sediment is also
related to fluvial geomorphic processes. Similarly, hydrodynamic effects result from the interaction of flow over stream
bedforms. Geomorphic processes build stream bedforms and
determine channel morphology, especially longitudinal gradient, bed roughness, and water depth, all of which influence
flow velocity. The relationship between geomorphology and
the other classes of processes is less direct, but still plays a role
in controlling these processes through channel form and the
size distribution of sediment that makes up the streambed.
This chapter focuses primarily on the hydrostatic and hydrodynamic processes where linkages to geomorphic processes
are most direct.
We organize our discussion of the interactions between
geomorphology and HEF using a hierarchical scaling framework developed for river networks (Frissell et al., 1986;
Bisson and Montgomery, 1996), starting at the whole network, through the stream segment, to the stream reach, to
the channel unit, and down to the subchannel unit scale.
We recognize that describing any given process or related
flow path at a single scale is somewhat arbitrary because of
the nested structure of the hyporheic flow net and dispersion
among HEF flow paths. Despite this, the concept of scale is
an important heuristic tool to organize our understanding
of hyporheic processes. In many senses, the reach scale is
the most informative scale at which to consider HEF. A single
reach, by definition, has characteristic channel morphology so
that the factors driving HEF within the reach are relatively
consistent. However, only a few of the geomorphic factors
driving HEF actually operate at this scale. Most of the drivers
work at the channel unit or smaller scales. Moreover, to
understand the importance of HEF in stream ecosystem processes, the cumulative effects of HEF must be evaluated at
scales much larger than a single reach.
9.13.2
9.13.2.1
205
The Effect of Geomorphology on HEFs
The Whole Network to Segment Scale
The geologic setting of the stream network is an important
factor determining the likely occurrence of HEF, but there have
been few attempts to study HEF at this broad scale. Rather, our
expectations are pieced together by drawing comparisons
among HZ studies that have been conducted in widely varying
geologic settings, at different locations in the stream network,
or under widely varying flow conditions. We expect that
geomorphic–hyporheic relationships will differ substantially
among different geologic settings.
Fluvial geomorphic studies have examined the factors that
determine the types of channel morphologies present within
stream networks (Montgomery and Buffington, 1997; Wohl
and Merritt, 2005; Brardinoni and Hassan, 2007). Montgomery and Buffington (1997) presented one such description
of the distribution of channel morphologies typical of many
mountainous landscapes. They showed that catchment area
and channel longitudinal gradient controlled the development of distinct channel types such that the channel types
tended to follow a characteristic sequence within a catchment
(Figure 2(a)). In their example, this sequence starts with
bedrock and colluvial channels in the steepest, uppermost
headwaters. As longitudinal gradients decrease, channels
change to cascades, to step–pool, to plane–bed, to pool–riffle,
and the largest, lowest-gradient rivers were typified as
dune–ripple channels. Along with these changes in channel
morphology, the following would be expected: decreased
longitudinal gradient and mean grain size of streambed sediment, and increased depth, width, hydraulic radius, and
flow velocity (Leopold and Maddock, 1953; Wohl and
Merritt, 2008).
In this chapter, we use Montgomery and Buffington’s
(1997) description of the sequence of channel types within
a catchment as a simple heuristic model to organize our
examination of the relative importance of the different processes that drive HEF within stream networks. We recognize
that local controlling factors commonly interrupt simple sequencing of channel types. For example, landslides may block
large mainstem channels, creating locally steep gradients over
the landslide debris and uncharacteristically low gradients in
the depositional reach directly upstream (Benda et al., 2003).
We also recognize that regional differences in geology and
geomorphology will lead to dramatically different spatial organization of channel types (see, e.g., characteristic channel
type in glaciated mountainous regions as described by Brardinoni and Hassan (2007)). Our descriptions of the spatial
organization of stream types and the resulting HEF processes
will have to be modified for any specific landscape.
Most hyporheic exchange results from head gradients
pushing water through the streambed. The amount of stream
water entering the hyporheic zone is thus a function of the
steepness of the head gradient and the saturated hydraulic
conductivity of the streambed and underlying aquifer. The
head gradients can be induced in many ways, but the two of
primary influence are the hydrostatic and hydrodynamic
processes. The relative importance of each of these processes is
expected to vary among channel types and with longitudinal
206
Geomorphic Controls on Hyporheic Exchange Across Scales: Watersheds to Particles
Hillslope
Hollow
Colluvialbedrock
Cascade
Step−
pool
Plane−
bed
Pool−riffle
Dune−ripple
Relative HEF
(a)
(b)
Hydrostatic
contribution
Hydrodynamic
contribution
Turbulent
diffusion
Longitudinal stream profile through stream network
Figure 2 (a) Hypothetical distribution of channel types along a stream profile in a mountainous stream catchment, and (b) the corresponding
relative contribution of turbulent diffusion and both hydrostatic and hydrodynamic processes to the total amount of HEF occurring within a
stream reach. (Note that boundaries between channel types are generally less distinct than shown here and that a range of conditions occurs
within each category, thus the contribution of each process varies both among and within each channel type.) (a) Redrawn with permission from
Montgomery, D.R., Buffington, J.M., 1997. Channel-reach morphology in mountain drainage basins. Geological Society of America Bulletin 109,
596–611.
gradient. In high-gradient streams, channel forms, such as
step–pool sequences or pool–riffle sequences, can create
very steep hydrostatic head gradients. Further, because of high
bed roughness and relatively shallow water depth, flow
velocities tend to be lower in small steep streams than in
larger, low-gradient streams (Leopold and Maddock, 1953;
Wondzell et al., 2007). By contrast, it is difficult for natural
processes to create steep changes in the longitudinal gradient
in low-gradient streams. Instead, stream flow interacts with
stream bedforms, such as dunes or ripples, such that hydrodynamic forces dominate the development of head gradients
through the streambed. Thus, we expect that hydrostatic effects
will dominate in high-gradient channels and that hydrodynamic processes will dominate in low-gradient channels
(Figure 2(b)). Further, because channel types and longitudinal gradients generally vary systematically within stream
networks, we further expect that hydrostatic effects will tend
to dominate in the upper portions of stream networks and
that the relative importance of hydrodynamic processes will
increase down the stream network.
9.13.2.2
The Reach Scale – Setting the Potential for
Hyporheic Exchange
The potential for HEF to occur varies within any given stream
reach. Roughly speaking, this potential is determined by
the factors that generate head differences that drive HEF,
the properties of the subsurface alluvium through which
HEF occurs, and the potential effect of lateral groundwater
inputs from adjacent hillslopes that might limit hyporheic
expression.
9.13.2.2.1
Losing and gaining reaches
Hyporheic exchange is likely to be more limited in strongly
gaining reaches than in neutral reaches because of steep
streamward hydrologic gradients surrounding the channel
(Wroblicky et al., 1998; Storey et al., 2003; Malcolm et al.,
2003, 2005; Cardenas, 2009). Similarly, where water is lost
to regional aquifers in strongly losing reaches, return flows
of stream water back to the stream are likely to be severely
restricted and thus also limit the expression of the hyporheic
zone (Cardenas, 2009). These patterns of gains and/or losses
are controlled, at some level, by regional groundwater and
catchment characteristics interacting with smaller-scale effects.
In large gaining rivers, Larkin and Sharp (1992) demonstrated
that the relative dominance of cross-valley versus down valley
flow paths through valley-floor aquifers varied depending on
the longitudinal gradient of the valley floor and the hydraulic
conductivity of the valley-floor alluvium. In higher-gradient
reaches (40.004 m m–1) and in areas with coarser substrate,
flow was predominantly down valley. Conversely, where
valley-floor gradients were shallower or sediment more finely
textured, flow tended to be toward the stream. Thus, the way
in which lateral inputs influence hyporheic exchange is not
solely a function of their magnitude, but also a function of the
ability of subsurface water to move down valley (Storey et al.,
Geomorphic Controls on Hyporheic Exchange Across Scales: Watersheds to Particles
2003). The ratio between these two factors – the magnitude
of the inputs relative to down-valley flow – determines how
hyporheic exchange is affected.
As a first approximation, the potential for down-valley
flow can be estimated using the relationships summarized in
Darcy’s law – that is, the product of the longitudinal valley
gradient, the saturated cross-sectional area of the floodplain
perpendicular to the direction of subsurface flow, and the
hydraulic conductivity of the alluvium. As lateral inputs increase, several factors may change: (1) water tables may rise,
thus increasing the saturated thickness and the cross-sectional
area through which water flows allowing the transmission of
more water, or (2) flow paths may begin to turn obliquely
toward the stream, which also increases the saturated crosssectional area and may also increase head gradients.
Under dry conditions when lateral inputs are relatively
small, the potential extent and magnitude of hyporheic exchange can be fully expressed (Figure 3(a)). As subsurface
flows turn toward the channel, they begin to limit the extent
of the hyporheic zone with only a minor effect on the HEF
(Wondzell and Swanson, 1996). If sufficiently large, lateral
inputs can severely limit both the spatial extent and magnitude of hyporheic exchange (Figure 3(b); Harvey and Bencala,
1993; Wroblicky et al., 1998; Storey et al., 2003; Cardenas and
Wilson, 2007; Malcolm et al., 2003; Soulsby et al., 2009).
Higher-gradient, low lateral inputs
(a)
Lower-gradient, high lateral inputs
(b)
Figure 3 Idealized conceptual model of the influence of lateral
inflows on hyporheic exchange flows. (a) A high-gradient stream
where floodplain alluvium has relatively high saturated hydraulic
conductivity under relatively dry conditions when lateral inputs are
low and easily transported down valley via subsurface flow. Lateral
inputs still reach the stream, but are diverted toward zones with
hyporheic upwelling. (b) A low-gradient stream where floodplain
alluvium has relatively low saturated hydraulic conductivity under
relatively wet conditions when lateral inputs are sufficiently large to
overwhelm down-valley transport, causing lateral inputs to cross the
valley toward the stream. Lateral inputs severely restrict hyporheic
exchange flows. Legend follows Figure 1.
207
Simple generalizations of where and when lateral inputs
will limit HEF are difficult because of the wide range of geomorphic settings in which HEF occurs and because the magnitude of lateral inputs changes with catchment wetness.
Lateral inputs are expected to be high when catchments are
wet and decrease as catchments dry out. However, lateral inputs are not spatially uniform. In steep mountainous settings,
the size of the upslope area draining directly to the valley floor
is important, concentrating lateral inputs in zones at the base
of hillslope hollows (Jencso et al., 2009). Lateral inputs may
persist the entire year at the bases of the largest hillslope
hollows. Most hillslope hollows are small, however, so that
most of the stream network would be disconnected from lateral inputs except for short periods of time when catchments
are very wet, for example, after large storms or during peak
snowmelt. We are unaware of similar studies relating topography to spatial patterns of hillslope inputs in areas of
low relief with humid climates. However, Storey et al. (2003)
reported that an extensive shallow surficial aquifer was present
along their lowland, low-gradient study reach and that lateral
inputs of groundwater substantially reduced both the extent
and the amount of HEFs except during summer base flow.
Clearly, the influence of lateral inputs in lowland catchments
may be much different from that in steep mountainous
catchments.
Changes in lateral inputs to streams do not occur in isolation. Rather, they are likely to be accompanied by corresponding changes in stream stage (and discharge). The change
in water-table elevations resulting from changed lateral inputs
must be considered relative to the accompanying changes
in stream stage. Although the number of studies examining
changes in hyporheic flow paths with changing catchment
wetness is limited, studies in small mountain streams suggest
that water-table elevations in the floodplain increase more
than stream stage so that HEF is typically more restricted
when catchments are wet (Figures 4(a) and 4(b); Harvey and
Bencala, 1993; Wondzell and Swanson, 1996; Stednick and
Fernald, 1999). Storey et al. (2003) reported similar results for
a lowland, low-gradient river.
In some cases, however, stream stage may change markedly
without corresponding changes in precipitation recharge or
changes in lateral inputs. Most examples of these processes
come from large, lowland rivers because river stage is controlled by processes far upstream. These ‘bank storage’ processes (Pinder and Sauer, 1971) have been recognized as a
form of transient hyporheic exchange (Figures 4(c) and 4(d))
that can result from both in-bank or over-bank floods (Bates
et al., 2000; Burt et al., 2002). In some situations, increased
stream stage may even lead to groundwater ridging in the
floodplain, reversing head gradients and limiting lateral
groundwater inputs. Similarly, hyporheic exchange through
stream banks can result from diel variations in stream stage
(and discharge) during snowmelt periods (Loheide and
Lundquist, 2009) or from tidally induced changes in water
elevations in coastal streams and rivers (Bianchin et al., 2010).
Transient hyporheic exchange may be especially evident in
regulated rivers where releases from dams (or other control
structures) can result in large and rapid changes in river stage
without corresponding local precipitation to recharge floodplain aquifers (e.g., Fritz and Arntzen, 2007; Lewandowski
208
Geomorphic Controls on Hyporheic Exchange Across Scales: Watersheds to Particles
Precipitation
Lateral inputs
Stage increase
(a)
(b)
Stage increase
Stage decrease
(c)
(d)
Figure 4 Idealized conceptual model of the influence changing stream stage on transient hyporheic exchange. (a, b) A losing reach at low base
flow is converted to a gaining reach during a storm because precipitation recharge and lateral inputs of hillslope water increase watertable elevations more than the corresponding increase in the stream stage. The original stream and water table position from (a) is shown in (b)
for reference (light-gray line). (c, d) An example of a river where changes in stream stage result from snowmelt, tidal influences, or dam releases
far upstream. Increased stream stage causes stream water to flow into the adjacent aquifer creating a losing stream reach. Conversely,
decreased stage leads to drainage of the aquifer creating a gaining reach. Alternating increases and decreases in stream stage lead to transient
hyporheic exchange. The neutral condition (where stream stage is equal to the water-table elevation) is shown for reference (black-and-white
dashed line). Legend follows Figure 1.
et al., 2009; Sawyer et al., 2009; Francis et al., 2010). However,
transient hyporheic exchange may not always result from
fluctuations in river stage. For example, Hanrahan (2008)
studied vertical HEF through the streambed of a large, regulated gravel bed river where stage sometimes changed by
nearly 2 m in an hour. For the most part, they did not observe
transient hyporheic exchange related to changes in stage. They
concluded that hydrostatic and hydrodynamic processes remained the dominant control on HEF. Notably, Hanrahan
(2008) did not examine lateral exchanges through the stream
banks, which can be more responsive to changes in stage than
are locations in the stream channel itself (Storey et al., 2003).
Water-table fluctuations in the floodplain at long distances
from the stream are not necessarily indicative of extensive HEF
because pressure fluctuations can propagate through surficial
(unconfined) aquifers much faster than does the actual flow of
stream water. This was clearly demonstrated by Lewandowski
et al. (2009) who showed that river water penetrated, at
most, only 4 m into the stream bank even though watertable fluctuations were observed more than 300 m from
the river.
HEF can occur in strongly gaining and losing reaches because of the nested structure of hyporheic flow paths, and
because HEF can occur at a variety of spatial scales. Thus, an
envelope of the HZ can be set within larger nonhyporheic flow
paths (Figure 3(b); Cardenas and Wilson, 2007). Similarly,
smaller-scale HEF can occur as a result of smaller-scale geomorphic drivers, even within a reach that is, overall, strongly
losing (Payn et al., 2009). Further, because HEF is dominated
by relatively near-stream flow paths that are short in length
and residence time (Kasahara and Wondzell, 2003), the
magnitude of HEF can be substantial, even in strongly gaining
reaches where the spatial extent of the hyporheic zone is
greatly restricted (Wondzell and Swanson, 1996; Cardenas
and Wilson, 2007; Payn et al., 2009).
9.13.2.2.2
Changes in saturated cross-sectional area
The saturated cross-sectional area of the floodplain (orthogonal to groundwater flow path direction) is one of the
factors determining the amount of groundwater transmitted
down valley through the valley-floor alluvium. Thus, any
change in the cross-sectional area along the length of a stream
reach will lead to parallel changes in the down-valley flow of
water through the floodplain, thereby driving downwelling
from, or upwelling to, the stream (Stanford and Ward, 1993).
Downwelling occurs where valley floors increase in width, for
example, downstream of bedrock-constrained reaches
(Figure 5(b); Poole et al., 2004, 2006; Acuna and Tockner,
2009). Conversely, upwelling occurs where valley floors
narrow at the lower end of wide unconstrained reaches
(Figure 5(a); Baxter and Hauer, 2000; Acuna and Tockner,
2009). Similarly, variations in the thickness of the surficial
aquifer, caused by variations in depth to bedrock or other
confining layers, drive similar patterns of upwelling and
downwelling. For example, upwelling commonly occurs just
upstream of bedrock sills with a subsequent transition to
downwelling just downstream of such bedrock sills as the
surficial aquifer again thickens (Figure 5(b); Valett, 1993).
This is easily observed in streams in arid regions during the dry
season, where perennial flow may only occur above bedrock
sills, which force the subsurface flow to the surface.
Geomorphic Controls on Hyporheic Exchange Across Scales: Watersheds to Particles
Unconstrained stream reach
Bedrock
gorge
(a)
(b)
Bedrock sill
Figure 5 Idealized conceptual model of the influence of the change
in saturated cross-sectional area of the floodplain on hyporheic
exchange flows. (a) The influence of change in valley constraint with
downwelling at the head of an unconstrained reach and upwelling at
the downstream end of the reach caused by the transition from
narrow bedrock gorges to wide alluvial valley floors. (b) The influence
of variations in depth to bedrock forcing upwelling upstream of a
bedrock sill and downwelling downstream, where the depth of
alluvium again increases. Legend follows Figure 1.
9.13.2.3
The Subreach to Channel-Unit Scale –
Hydrostatic Processes
Geomorphic features of the stream channel and valley floor
within stream reaches control the elevation of surface water
and can thereby create significant head gradients through the
valley-floor alluvium, driving HEF. Because these geomorphic
features are static on the timescales typical of hyporheic exchange (hours to weeks), they are broadly recognized as
‘hydrostatic processes’.
9.13.2.3.1
Step–pool and pool–riffle sequences
One of the best-studied examples of hydrostatic processes
involves the changes in water surface elevation along a
step–pool sequence and the resulting head gradients that drive
HEF (Figure 1; Harvey and Bencala, 1993). Harvey and
Bencala (1993) showed that the change in the longitudinal
gradient of the stream channel (which approximates the
stream energy profile) drove HEF. They also observed that HEF
flow paths tended to be curved – first curving away from
the stream above the step or riffle and then curving back to
the stream below the step or riffle. Building from their observations, model analyses showed that along an idealized
straight channel with homogeneous isotropic porous sediment, hyporheic flow paths around a change in the longitudinal gradient will exploit the full three-dimensional (3D)
saturated volume along the channel, thus extending both
vertically beneath the streambed and horizontally through the
streambanks and near-stream aquifer (Figures 1(a) and 1(b)).
Real streams are substantially more complicated, however,
such that changes in hydraulic conductivity of the alluvium,
bends in the channel, and the spatial location of lateral
groundwater inputs lead to the development of a complicated
flow net through the valley floor (e.g., Cardenas and Zlotnik,
2003). Despite these complexities, the steepness of the
hydraulic head gradient imposed by the change in the
209
longitudinal gradient and the saturated hydraulic conductivity
control the amount of stream water exchanged with the
subsurface.
Many factors can modify the effect of steps or riffles on
HEF. For example, the height of the step (or steepness of the
riffle) determines the head gradient available to drive HEF so
that a single very large step has the potential to drive more
HEF than if the same amount of elevational change is spread
over several smaller steps (Kasahara, 2000). Because of this,
large wood can be important in determining the amount of
HEF in forest streams. Single logs tend to create frequent,
small obstructions that collect and store small amounts of
sediment, forming pool–step sequences in which the extent
of the hyporheic zone tends to be small (Wondzell, 2006).
Although logjams are less common, they can create large obstructions storing sediment in wedges several meters deep and
Z10 m in length, and significantly widen constrained stream
channels. Consequently, logjams can form extensive hyporheic zones in steep, confined mountain streams (Wondzell,
2006).
Large, channel-spanning logs can wedge into steep
narrow channels, forcing the accumulation of sediment in
channels, converting bedrock reaches to alluvial reaches with a
step–pool morphology (Montgomery et al., 1996), thereby
greatly enhancing HEF. Similarly, large wood can force plane–
bed channels into a pool–riffle morphology (Montgomery
et al., 1996), which should lead to more HEF than would
be present in a comparable wood-free channel. Large wood
can have the opposite effect in channels that would have a
free-formed pool–riffle morphology. In one documented case,
accumulations of large wood tended to force a pool–riffle
channel toward a step–pool morphology (Wondzell et al.,
2009). The channel adjusted to removal of all large, in-stream
wood by developing a better-defined pool–riffle structure
around meander bends, leading to increased sediment storage.
Continued channel adjustment over time, following the
removal of large wood, eventually led to substantial increases
in HEF.
The size, spacing, and sequence of channel units (e.g.,
pools and riffles) along the stream longitudinal profile can
also affect HEF (Anderson et al., 2005; Gooseff et al., 2006).
Anderson et al. (2005) made detailed measurements of
channel profiles and patterns of HEF, and showed that channel unit size and spacing increased as did the length of
channel characterized by downwelling with increasing drainage area in a mountainous stream catchment. Gooseff
et al. (2006) built on these results, examining HEF using
2D groundwater models of idealized longitudinal profiles of
mountain streams. The modeling results of Gooseff et al.
(2006) confirmed that both channel unit spacing and size
were important in determining hyporheic exchange patterns
of upwelling and downwelling. Perhaps more surprising,
however, was the observation that the sequence of channel
units also affected simulated HEF. Gooseff et al. (2006)
compared pairs of idealized stream reaches that varied only by
the way the longitudinal gradient changed over the pool–riffle
sequence – that is, the slope of the riffle was gradual on its
upstream end and steepest at its downstream end (described
as a pool–riffle–step sequence) versus riffles that were initially
steep with the slope decreasing toward the downstream end
210
Geomorphic Controls on Hyporheic Exchange Across Scales: Watersheds to Particles
(described as a pool–step–riffle sequences). Simulated downwelling lengths were substantially longer for pool–riffle–step
sequences than for pool–step–riffle sequences.
9.13.2.3.2
Meander bends and point bars
A variety of channel and valley-floor morphologic features, in
addition to changes in the longitudinal gradient, create head
gradients with the potential to drive HEF. These include
channel meander bends and associated point bars, back
channels or floodplain spring brooks, and islands set between
main and secondary channels. In all these cases, differences in
the elevational head of surface water between two channels,
between different points in a single channel around a meander
bend, or between points on opposite sides of an island create
head gradients that drive HEF. For example, head gradients
through the point bar in a meander bend are steeper than the
longitudinal gradient of the stream channel around the point
bar (Peterson and Sickbert, 2006) so that stream water infiltrates the upper end of the point bar and is returned to the
channel at the lower end of the point bar (Figure 6(a); Vervier
and Naiman, 1992). More generally, these exchange flows
occur across the full length of meander bends and are
(a)
(b)
Pool/Run
(c)
Figure 6 Idealized conceptual model of the influence of meander
bends on hyporheic exchange flow. (a) Simple, low-radius meander
with HEF traversing the point bar and floodplain. (b) High-radius
meander with incipient meander-cutoff, where the short distance
across the neck leads to much higher head gradients and thus
greater HEF through the neck than the remainder of the meander bar.
(c) Meander bend with riffles located at the inflections between
adjacent meanders so that head gradients through the point bar are
low and much of the HEF occurs around the riffles, driven by
longitudinal changes in gradient. Legend follows Figure 1.
influenced by both the change in stream water elevation
around the meander bend and the plan-view shape of the
meander bend. Highly evolved meander bends support steep
head gradients across the mender neck because of the close
proximity of the stream channels (Figure 6(b); Boano et al.,
2006; Revelli et al., 2008) so that HEF is dominantly located in
the meander neck, with much reduced HEF across the remainder of the meander where head gradients are much lower.
In other cases, meanders develop a characteristic pattern of
alternating pools and riffles, with riffles located at the thalweg
crossovers in the inflections between adjacent meanders and
pools or low-gradient runs wrapping around the point bar
(Figure 6(c)). This combination of channel morphologic features can create complex HEF flow paths within meander
bends. The residence times of HEF traversing meander bends
can be quite short where meanders are small and saturated
hydraulic conductivities are high (Pinay et al., 2009). Conversely, residence times of HEF may be extremely long in meander bends of low-gradient rivers with fine-textured sediment
(Boano et al., 2006; Peterson and Sickbert, 2006).
9.13.2.3.3
Back channels and floodplain spring brooks
Channel planforms are generally complex in wide floodplains,
including a network of old or abandoned channels. If the
upstream ends of these channels are plugged with sediment
and if the downstream ends are sufficiently incised to intercept
the water table and are connected back to the river at
their downstream ends, they will act as drains, imposing head
gradients from the stream to the old channel (Figure 7(a);
Wondzell and Swanson, 1996; Poole et al., 2006). These
channels are also known as floodplain spring brooks because
water upwells into the channel, forming a spring at its head. In
addition to creating HEF, these channels will capture whatever
water is in the surficial aquifer of the floodplain, including
down-valley flows from upstream locations, and lateral inputs
of groundwater or hillslope water from the valley margin.
However, because lateral inputs tend to be small and spatially
isolated (Jencso et al., 2009; and as discussed above), floodplain spring brooks will most generally be fed by HEF
(Wondzell and Swanson, 1996; Jones et al., 2007).
Abandoned channels can also be plugged at their downstream ends and open to the river at their upstream ends. In
this case, stream water can flow into the abandoned channel,
infiltrate the channel bed, and raise the water table in the
middle of the floodplain, thereby creating head gradients and
driving HEF from the abandoned channel back to the main
stream channel (Figure 7(b)). More complex situations arise
when the longitudinal gradients in either the back channel or
mainstem channel are interrupted by steeper riffles or steps.
Figure 7(c) shows the interactions between a back channel
and riffle. Above the riffle, the elevation of water in the main
channel is higher than the back channel so water flows toward
the spring brook. Downstream of the riffle, water in the main
channel is lower than the back channel so that the back
channel loses water over its downstream extent, eventually
going dry before reaching the main channel.
The channel planform features that drive HEF can occur
over a range of spatial scales, and their influence may change
through time as the stage height of water in the main channel
Geomorphic Controls on Hyporheic Exchange Across Scales: Watersheds to Particles
(a)
211
(a)
(b)
(b)
C
(c)
(c)
Figure 7 Idealized conceptual model of the influence of back
channels on hyporheic exchange flows. (a) A back channel is incised
below the water table, acts as a drain, and creates head gradients
from the main channel to the back channel. (b) A back channel is
plugged near its downstream end, conducts water onto the
floodplain, raises the water table, and creates head gradients toward
the main channel. (c) Complex pattern of HEF caused by interactions
between a riffle in the main channel and a back channel.
Paleochannels (dashed lines) support preferential flow. Legend
follows Figure 1.
changes. For example, a small gravel bar may have low points
along the stream bank. At high stage, the entire gravel bar may
be submerged. As the stage decreases, the center of the bar
may become exposed, creating a secondary channel along the
bank. As the stage decreases further, flow may become discontinuous through the secondary channel such that it functions as a drain if it is plugged at the upstream end, or functions
as a conduit allowing stream water to infiltrate the surface of
the gravel bar if it is plugged at its downstream end. Old
channels in large floodplains may act similarly, with continuous flow along their full length during floods, but becoming
disconnected at intermediate to low stage, or even dry completely during periods of minimum discharge. In large floodplain reaches, these channels can be hundreds of meters to
kilometers in length, extending nearly the full length of the
stream reach (Poole et al., 2006; Arrigoni et al., 2008).
9.13.2.3.4
Secondary channels and islands
Islands present a special case of back channels in which the
channel is continuously connected to the main channel over
Figure 8 Idealized conceptual model of the influence of mid-stream
islands on hyporheic exchange flows. (a) Parallel and smooth
longitudinal gradients in the channels on both sides of the island
create HEF flow paths that parallel stream flow. (b) Riffles at the
head of the island enhance head gradients leading to greater HEF.
(c) Offset riffles create strong cross-island head gradients and flow
paths, resulting in more HEF but with shorter flow path lengths and
residence times. Legend follows Figure 1.
its full length. The hyporheic hydrology of islands has not
been extensively studied. However, we expect that the surface
water elevations in channels bounding the island create
boundary conditions for total head and control HEF through
islands as is generally indicated by the available literature
(Dent et al., 2007; Francis et al., 2010). If channels along both
sides of the island are parallel and symmetric with constant
longitudinal gradient, then flow through the island will parallel the channels and the head gradient driving flow will
equal the overall longitudinal gradient of the stream reach
(Figure 8(a)). If riffles are present in the channels, the head
gradient through the island adjacent to the riffles can be much
steeper than the reach averaged longitudinal gradient
(Figure 8(b)). Also, if riffles are displaced along the primary
and secondary channels surrounding an elongated island such
that a riffle is located near the head of the island in one
channel and near the tail of the island in the second channel,
the resulting head gradients would tend to drive flows laterally
through the island, leading to very large cross-sectional areas
experiencing HEF, and therefore large amounts of HEF, albeit,
with shorter-length flow paths (Figure 8(c)). Although islands
may be uncommon in most channel types, they may dominate HEF in braided and anastomosing stream reaches (Ward
et al., 1999; Arscott et al., 2001). Given the complexities
212
Geomorphic Controls on Hyporheic Exchange Across Scales: Watersheds to Particles
of potential sizes and shapes of islands and patterns in
longitudinal gradients in the bounding channels, the resulting
flow nets, residence times, and amounts of HEF are likely to
vary widely.
9.13.2.3.5
Spatial heterogeneity in saturated hydraulic
conductivity
Fluvial processes control the depositional environment on the
streambed and across the floodplain creating spatial heterogeneity in the texture of deposited and reworked sediment
across a range of scales, from the surface of the streambed to
the entire floodplain. Because sediment texture is closely related to saturated hydraulic conductivity (K), these processes
can substantially influence HEF. However, because of the
difficulties in quantifying these patterns at the scales at
which they influence HEF, they have been relatively little
studied. At fine scales, streambed roughness can control the
depositional environment across the streambed (Buffington
and Montgomery, 1999), which lead to spatial patterns in the
distribution of K within the streambed (Genereux et al.,
2008), which in turn can influence both the location and
amount of HEF. HEF will be restricted where the streambed is
clogged with fine sediment and preferentially located in zones
with higher K. Experiments in flumes have also shown that
HEF can also influence patterns of fine-sediment deposition,
with fine sediment preferentially deposited in downwelling
zones (Packman and MacKay, 2003; Rehg et al., 2005), which
may explain differences in K between upwelling and downwelling zones observed in a steep headwater stream (Scordo
and Moore, 2009).
Spatially heterogeneous patterns in K influence HEF. For
example, groundwater flow modeling studies using homogeneous versus heterogeneous K showed that spatial heterogeneity may add substantial complexity to the spatial patterns
of the hyporheic flow net (Woessner, 2000). Where relatively
high K regions are aligned parallel with head gradients, they
create preferential flow pathways (Wagner and Bretschko,
2002) that can increase the total amount of HEF (Cardenas
and Zlotnik, 2003; Cardenas et al., 2004). Results from
Cardenas et al. (2004) showed that influence of heterogeneity
in K was relatively greater in lower-gradient streams and where
head gradients driving HEF were reduced. To our knowledge,
the influence of fine-grained heterogeneity has not been studied in steeper channels where hydrostatic processes dominate.
Fluvial processes also influence spatial patterns in K at
the scale of the entire floodplain. Especially important is the
layering of stream and floodplain alluvium. Layering can
create strong vertical anisotropy (Chen, 2004), limiting
vertical exchange and promoting lateral flows through the
streambed and floodplain (Packman et al., 2006; Marion
et al., 2008). Overbank deposition can also bury back channels creating ‘paleochannels’ where coarse streambed alluvium
is buried under finer floodplain soils (Stanford and Ward,
1993; Stanford et al., 1994; Poole et al., 2004). If these
paleochannels intercept the water table, they will function as
large preferential-flow pathways that can route water the full
length of a floodplain. In this regard, they function much like
a subsurface version of back channels or floodplain spring
brooks – either acting as drains lowering the water table in the
floodplain and imposing head gradients from the stream to
the paleochannel, or acting as distributaries routing water into
the floodplain and imposing head gradients from the paleochannel to the stream. Locations of paleochannels are sometimes evident from shallow depressions along the floodplain.
In other cases, over-bank deposition will have completely
filled old channels so that there is no surficial indication on
the flat floodplain surface. The influence of paleochannels is
difficult to discern because networks of widely spaced wells are
unlikely to find and trace the location of these features along
the length of the floodplain. As a consequence, their influence
on HEF has not been widely studied.
9.13.2.4
The Bedform Scale – Hydrodynamic Processes
Channel hydraulics, and the spatial and temporal distribution
of velocity (kinetic energy) across streambeds, are significantly
influenced by the form of the channel and the bedforms that
occur in channels. The continuous feedback between pressure
distribution and shear stress across the bed surface and the
potential to erode the bed will cause turnover exchange to
occur during times of high flows. During lower flows, when
bed sediment is relatively stable, bedforms cause some level of
form drag on the flows, inducing pressure distributions across
the bedforms, thereby driving HEF at a scale smaller than the
bedform (Figure 9). The size of the bedform is set by both the
energy regime of the reach and the material that makes up
the reach, and the form drag induced on the water column by
the bedform is of course partly controlled by its size. Thus, the
scale of HEF flowpaths induced by hydrodynamic exchange
across the bedforms will scale in part with the size of bedforms
present (Cardenas et al., 2004). Finally, the heterogeneity
of the bed material that makes up the bedforms will have a
distinct control on the flux rate and actual flowpaths through
and around the bedforms (Sawyer and Cardenas, 2009).
In sand-bed streams, hydrodynamic HEF has been extensively studied both theoretically and empirically. Typical
bedforms in sand-bed streams are dunes and ripples, which
have a fairly predictable geometry and spacing, based on bed
sediment composition and flow rate. Thibodeaux and Boyle
(1987) pioneered investigations of the hydrodynamic pressure
distribution across dunes, noting the penetration of channel
water into the porous bedforms. Further development of a
‘pumping exchange’ model by Elliott and Brooks (1997a,
1997b) expanded the ability to predict HEF and associated
Stream flow
Figure 9 Idealized longitudinal-section in the center of a straight
stream channel with bedforms (triangular dunes) showing the
interaction with stream flow that creates regions of low and high
pressure on the streambed that drive HEF. Nonhyporheic subsurface
flows, known as underflow (dashed arrows), are present beneath the
hyporheic zone. Legend follows Figure 1.
Geomorphic Controls on Hyporheic Exchange Across Scales: Watersheds to Particles
solute dynamics in channel–bed systems. Whereas most
studies of hydrodynamic exchange processes were generally
carried out in or applied to flume studies, there has been at
least one application of incorporating the pumping exchange
model to tracer transport in field studies. Salehin et al. (2003)
studied the transport of tracer along several kilometers of Sava
Brook in Sweden and successfully applied a solute transport
model to the observed data to explain long residence-time
distributions using the pumping exchange model theory. The
predictability of dune and ripple sizing and spacing makes the
pumping exchange model a useful tool to explore HEF in
sand-bed streams and rivers.
In gravel-bed streams, bedform types may be generally
predictable (i.e., Montgomery and Buffington, 1997; Wohl
and Merritt, 2008; Chin, 2002), but the exact geometry and
spacing of bedforms are less predictable, particularly at a scale
that will directly influence head distributions across and along
the channel. Hence, the velocity distribution in the channel
and around the bedform, which contributes to hydrodynamic
exchange, is also unpredictable. Tonina and Buffington (2007)
conducted careful studies of total pressure distribution
across streambeds in flumes that had realistic geometry of a
pool–riffle sequence in a gravel-bed channel. Their results
indicated that total head distribution (i.e., incorporating velocity head in addition to hydrostatic head) was important to
exchange at focused points in the channel where high velocity
occurred. Further, they confirmed that, in general, there was
little or no contribution of velocity head to parts of the bed
that were overlain by deeper, slower flow, and therefore a
hydrostatic representation of exchange will likely be more
applicable in these locations.
Regardless of the predictability of bedform geometry and
spacing, the associated hydrodynamic HEF may induce only
limited lengths of exchange in the subsurface because much of
the exchange dynamics are expected to be vertical rather than
lateral. Exchange lateral to the channel is more likely to be
driven by hydrostatic gradients set up across meander bends or
bars (as described above). Hydrodynamic HEF will contribute
to, but be only one component of, total HEF in natural
channels, and its importance will be dictated by both channel
hydraulics and, if present, competing hydrostatic factors that
can create steeper head gradients.
9.13.2.5
The Particle Scale – Turbulent Diffusion
At the particle scale on streambeds, turbulent diffusion is
significantly influenced by the size and arrangement of surface
sediment. Because turbulent diffusion is induced by the momentum transfer between the water column and the porous
media, HEF due to turbulent diffusion is a function of the
decreasing velocity profile within the surface layers of the
porous media (Shimizu et al., 1990). Thus, the distribution of
sediment at the surface will greatly influence the potential for
energy and mass transfer within this zone. Turbulent diffusion
HEF is prominent in gravel-bed streams where surface pores
are more likely to accommodate such open exchanges of
momentum across the bed (Tonina and Buffington, 2009).
Beds composed of sand particle sizes and smaller provide too
much resistance to the momentum exchange between the
water column and the bed. Hence, turbulent diffusion is more
213
likely to be an important component of HEF in low-order,
high-gradient streams (Figure 2(b)). Careful theoretical and
empirical research on turbulent diffusion has been conducted
largely on planar beds (Shimizu et al., 1990; Habel et al.,
2002). Therefore, in the complex bed topography of typical
gravel channels, turbulent diffusion will be a component of
HEF, likely not the singular driver of HEF.
9.13.3
9.13.3.1
Discussion
Multiple Features Acting in Concert
In the examples presented above (Figures 1 and 3–9), we
have mostly focused on single types of channel morphologic
features that drive or enable hydrostatic and hydrodynamic
HEF. However, these features never occur in isolation. Rather, a
single stream reach will typically contain many of the morphologic features described above. Interactions among these
features are likely to be important in determining the actual
HEF in any given stream reach. In some cases, the effects
of multiple features could be additive and result in higher
HEF than if they did not co-occur. For example, cross-valley
flow paths between main channels and floodplain spring
brooks can be accentuated by riffles (Figure 7(c)). However,
interactive effects could also cancel, for example, where riffles
at the inflection points of meander bends reduce head gradients through point bars (Figure 6(c)). The interactions between different processes driving or enabling HEF is complex,
and, to some degree, site specific, making it difficult to
quantify the effects of these interactions. Because of these
difficulties, there are relatively few comparative studies that
have examined multiple processes concurrently, within natural
stream channels, and attempted to evaluate the net effect of
each process on the total HEF within stream reaches.
Sensitivity analyses with groundwater flow models calibrated to simulate HEF in a studied stream reach provide one
opportunity to examine the relative importance of channel
morphologic features on HEF where multiple features are
present in a single reach. For example, Kasahara and Wondzell
(2003) examined a number of channel morphologic features
among stream reaches of different sizes in a mountainous
stream network under conditions of summer base flow discharge. In all cases, the single strongest driver determining the
amount of HEF occurring within the simulated stream reaches
was the change in longitudinal gradient over step–pool sequences in the second-order channel (Figure 10(a)) and
pool–riffle sequences in the fifth-order channel. The shape of
the hyporheic flow net in the fifth-order stream, however, was
strongly controlled by the presence, location, and relative
elevation difference between water in the main channel and
the back channels (Figure 10(b)). Similarly, Cardenas et al.
(2004) examined sediment heterogeneity, size of bedforms,
and both longitudinal and lateral head gradients in a lowgradient, sand-bed stream. They found that HEF was greater
where bedforms had higher amplitude and were more closely
spaced. Spatial heterogeneity in K increased HEF relative to
homogeneous simulations, as did inclusion of lateral head
gradients, but the effect was small relative to the effect of the
size and spacing of bedforms.
214
Geomorphic Controls on Hyporheic Exchange Across Scales: Watersheds to Particles
S
S
I
S
S
S
S
B
S
S
S
Scale = 20 m
(a)
M
Log
Bedrock
Wetted stream channel
Valley-floor alluvium
Back channel
Hillslope or terrace
Equipotential (0.2 m)
Hyporheic flow path
M
R
B
T
R
R
I
R
R
R
Scale = 50 m
B
(b)
Figure 10 Examples of complex hyporheic flow paths resulting from interactions between channel morphologic features: (a) a steep, secondorder step–pool channel with abundant large wood and (b) a moderate-gradient, fifth-order pool–riffle channel with two major spring brooks.
Note the difference in spatial scale between the two stream reaches. Letters indicate morphologic features driving HEF: S, steps; R, riffles; M,
meander bends; B, back channels/spring brooks; I, islands; and T, a steep riffle at the mouth of a tributary. Equipotential intervals (dashed lines)
are 0.2 m. Hyporheic flow paths (arrows) are hand drawn to indicate general direction of hyporheic flow through the valley floor.
In low gradient channels, morphologic features can interact with changes in steam stage and lateral groundwater inputs
in ways that can substantially influence the amount of HEF
over time, across seasons, or within a single storm event.
Storey et al. (2003) examined HEF in a pool–riffle sequence
at both high- and low-base flow discharge. At high stage, the
stream tended to drown the riffle, substantially reducing
the change in the longitudinal gradient over the pool–riffle
sequence and thus reducing HEF. By contrast, at low stage, the
water surface more closely followed the streambed topography, thus creating steeper head gradients that supported
more HEF. Storey et al. (2003) also showed that lateral inputs
during the wet season were sufficient to eliminate most of the
HEF through the riffle. Cardenas and Wilson (2006) showed
that low rates of groundwater discharge limited the extent of
the HZ formed by the hydrodynamics of stream bedforms,
and that high rates of groundwater discharge could completely
eliminate HEF.
We know of only one study comparing the relative influence of hydrostatic and hydrodynamic effects. In a flume,
Tonina and Buffington (2007) investigated the control of total
head (i.e., including dynamic head) in driving hyporheic exchange. Their results suggested that there are specific locations
within channels where the velocity head can provide additional potential and thereby influence the pattern of hyporheic exchange.
9.13.3.2
Change in Processes Driving HEF through the
Stream Network
Hyporheic exchange will vary widely across the sequence of
channel types occuring in stream networks (Figure 2; Buffington and Tonina, 2009). Channel networks generally follow a pattern of steep headwaters to low-gradient reaches
downstream. In mountain stream networks in particular,
gradient changes are expected to be accompanied by channel
morphology changes resulting in a sequence of distinct
channel morphologies (Figure 2(a)). Obviously, bedrock
reaches have negligible hyporheic zones (Gooseff et al., 2005;
Wondzell, 2006). We are unaware of any studies of HEF in
colluvial and cascade channel morphologies; however, the
extremely high longitudinal gradients of these channels likely
result in high-velocity underflow, which has been shown to
restrict the extent of the hyporheic zone (Storey et al., 2003).
In addition, the relatively disorganized structure of the bed
sediment prevents development of stepped water surface
Geomorphic Controls on Hyporheic Exchange Across Scales: Watersheds to Particles
profiles so that hydrostatically driven exchange due to longitudinal changes in gradient will likely be low. Turbulent diffusion is likely to be a primary driver of HEF in such reaches
(Figure 2(b)).
Free-formed step–pool channels occur at slightly lower
gradients (Figure 2(a)). These channels have well-organized
structure with periodic spacing of both steps and pools (Chin,
2002; Wohl and Merritt, 2008) that have been shown to be
primary drivers of HEF (Figure 2(b); Kasahara and Wondzell,
2003). The addition of large wood can substantially increase
sediment storage (Nakamura and Swanson, 1993; Montgomery et al., 1996), the development of step–pool structure,
and the extent, amount, and residence times of HEF in these
stream reaches (Wondzell, 2006). Other hydrostatic factors
tend to have less dominance on HEF; these reaches have low
sinuosity so meander bends are uncommon and steep longitudinal gradients limit the potential for back channels to
create lateral HEF flow paths.
We are unaware of any published studies examining HEF in
plane–bed channels. However, we expect HEF to be lower than
in either step–pool or pool–riffle channels (Figure 2(b)). The
streambed tends to be smoothly graded in these channels as
suggested by their name, and there is low spatial heterogeneity
in surface texture (Buffington and Montgomery, 1999). Pools
are widely spaced, and both steps and riffles are rare. Although
these channels occur as free-formed morphologies, pool–riffle
channels can be converted to plane–bed channels by land-use
practices that increase sediment supply and through the direct
removal of large wood, with concurrent decreases in HEF.
Lower in the stream network, channels tend have lower
longitudinal gradients (Figure 2(a)), and even in mountainous areas, unconstrained stream reaches become increasingly
common. Channel planforms can be quite complex in these
rivers and, as a consequence, a wide array of channel geomorphic features influences HEF. Braided and anastomosing
channels may form where sediment loads are high and stream
banks are erodible; the complex of channels likely leads to
substantial HEF through islands. Meandering channels form
under lower sediment loads and where banks are more stable.
Meandering channels typically have pool–riffle morphologies,
although complexes of secondary channels, back channels,
and paleochannels are common, a legacy of past floods,
channel avulsions, and overbank deposition. Because most
HEF occurs along short, near-stream flow paths, riffles are the
dominant feature determining the amount of HEF (Kasahara
and Wondzell, 2003). However, the shape of the hyporheic
flow net and the residence time distribution of HEF will be
strongly influenced by the complex of channel planforms.
Finally, hydrodynamic processes are expected to dominate
in streams with relatively mobile streambeds characterized by
dune–ripple bedforms. These streams have low longitudinal
gradients and therefore channel morphologic features tend
not to create steep hydrostatic head gradients (Figure 2(b)).
Other exchange processes are likely to be related to specific
conditions. Turnover exchange will only occur when bed material is mobile – a characteristic feature of both anastomosing
and dune–ripple channels. Transient exchange will only be appreciable during wet catchment conditions, when channel stage
is high and surrounding groundwater tables are comparatively
low. However, transient exchange may be a dominant form of
215
HEF in regulated rivers where stage fluctuates over daily cycles
due to hydroelectric generation. Turbulent diffusion, on the
other hand, is likely to occur in gravel-bed sections of the network, likely with the greatest potential influence in either cascade
or plane–bed sections of stream networks where stream velocities
are expected to be high.
9.13.4
Conclusion
Hyporheic exchange results from distinct processes, and the
relations between these processes and geomorphology are
well understood from a mechanistic perspective. Thus, geomorphology provides a critical framework to understand
hyporheic processes and how they change with location
within a stream network, and, over time, in response to
changes in stream discharge and catchment wetness. To the
degree that these geomorphic patterns are predictable, they
provide the foundation for hydrologists to make general predictions of the relative importance of the hyporheic zone at
the scale of entire catchments. Reach-to-reach variability is
high in stream networks, however, so understanding HEF at
the reach scale continues to require detailed study of specific
stream reaches. These studies are difficult and current methodological approaches are insufficient to fully examine the full
suite of processes that account for patterns of HEF in any
specific stream reach. Consequently, hyporheic studies tend to
focus on a single factor, or at most a small subset of the factors
driving HEF. Hyporheic researchers recognize that such studies
are incomplete. Detailed, holistic understanding of the importance of different processes in driving HEF, how the relative
importance of these processes changes with location in the
stream network, with the specific structure of any given stream
reach, and with changes in discharge and lateral groundwater
inputs remains elusive.
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Biographical Sketch
Steve Wondzell is a research riparian ecologist in the US Forest Service’s Pacific Northwest Research Station,
located in Olympia, Washington, USA. His research broadly focuses on both basic and applied problems in
watershed management and riparian and aquatic ecosystems. His basic research focuses on the interactions
between hydrological, geomorphological, and ecological processes that create, maintain, or modify aquatic and
riparian habitats, and the ways in which these processes either interact with, or are affected by, land-use practices.
His applied research focuses on developing models and decision support tools that synthesize the current
knowledge of aquatic and riparian systems into forms that can help inform management decisions at large spatial
and temporal scales.
218
Geomorphic Controls on Hyporheic Exchange Across Scales: Watersheds to Particles
Dr. Michael Gooseff is an associate professor in the Department of Civil & Environmental Engineering at Penn
State University. He began his research career studying hyporheic zones associated with glacial meltwater streams
in the McMurdo Dry Valleys of Antarctica for his PhD work at the University of Colorado. He continued to study
hyporheic exchange at broader scales in his postdoctoral research in the HJ Andrews Experimental Forest of
central Oregon, USA. He then moved to conducting hyporheic research in tundra streams of Alaska, mountain
and valley streams in the intermountain western USA, and in restored and unrestored streams. He continues to
have active research programs in both polar regions and continues strong collaborations to develop new techniques to study stream-groundwater interactions and implications.
TREATISE ON
GEOMORPHOLOGY
EDITOR-IN-CHIEF
John F. Shroder
University of Nebraska at Omaha, Omaha, NE, USA
VOLUME 9
FLUVIAL GEOMORPHOLOGY
VOLUME EDITOR
ELLEN WOHL
Colorado State University, Fort Collins, CO, USA
Academic Press is an imprint of Elsevier
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Copyright r 2013 Elsevier Inc. All rights reserved
The following chapters are US Government works in the public domain and not subject to copyright:
4.4 Nanoscale: Mineral Weathering Boundary
6.4 Karst Geomorphology: Sulfur Karst Processes
6.24 Cave Sediments as Geologic Tiltmeters
7.3 Stress, Deformation, Conservation, and Rheology: A Survey of Key Concepts in Continuum Mechanics
7.7 Surface Runoff Generation and Forms of Overland Flow
7.23 Mass-Movement Causes: Earthquakes
9.2 A River Runs Through It: Conceptual Models in Fluvial Geomorphology
9.9 Suspended Load
9.23 Waters Divided: A History of Alluvial Fan Research and a View of Its Future
9.29 Incised Channels: Disturbance, Evolution and the Roles of Excess Transport Capacity and Boundary Materials in Controlling
Channel Response
9.36 Geomorphic Classification of Rivers
11.9 Loess and its Geomorphic, Stratigraphic, and Paleoclimatic Significance in the Quaternary
11.20 Anthropogenic Environments
12.2 Riverine Habitat Dynamics
12.7 Vegetation Ecogeomorphology, Dynamic Equilibrium, and Disturbance
12.8 The Reinforcement of Soil by Roots: Recent Advances and Directions for Future Research
12.21 Interactions among Hydrogeomorphology, Vegetation, and Nutrient Biogeochemistry in Floodplain Ecosystems
2.10 Hillslope Soil Erosion Modeling
Copyright r 2013 R Brazier
2.12 Morphodynamic Modeling of Rivers and Floodplains
Copyright r 2013 A Nicholas
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CONTENTS
Preface
xxvii
Foreword
xxix
VOLUME 9 Fluvial Geomorphology
9.1
Treatise on Fluvial Geomorphology
E Wohl
1
Scales and Conceptual Models
9.2
A River Runs Through It: Conceptual Models in Fluvial Geomorphology
JE O’Connor, and MG Wolman
GE Grant,
6
Drainage Basin Processes and Analysis
9.3
Subsurface and Surface Flow Leading to Channel Initiation
9.4
Network-Scale Energy Distribution
SK Kampf and BB Mirus
P Molnar
22
43
Channel Processes
9.5
Reach-Scale Flow Resistance
R Ferguson
50
9.6
Turbulence in River Flows
9.7
The Initiation of Sediment Motion and Formation of Armor Layers
HE Schott
T Buffin-Bélanger, AG Roy, and S Demers
69
EM Yager and
9.8
Bedload Kinematics and Fluxes
JK Haschenburger
9.9
Suspended Load
9.10
Bedforms in Sand-Bedded Rivers
9.11
Wood in Fluvial Systems
9.12
Influence of Aquatic and Semi-Aquatic Organisms on Channel Forms and
Processes
JA Riggsbee, MW Doyle, JP Julian, R Manners, JD Muehlbauer, J Sholtes, and
MJ Small
RA Kuhnle
87
103
124
JG Venditti
137
AM Gurnell
163
189
Exchanges/Fluxes
9.13
9.14
9.15
Geomorphic Controls on Hyporheic Exchange Across Scales: Watersheds to
Particles
SM Wondzell and MN Gooseff
203
Reciprocal Relations between Riparian Vegetation, Fluvial Landforms, and Channel
Processes
DM Merritt
219
Landslides in the Fluvial System
244
O Korup
Channel Patterns
9.16
River Meandering
JM Hooke
9.17
Morphology and Dynamics of Braided Rivers
260
P Ashmore
289
v
vi
Contents
9.18
Hydraulic Geometry: Empirical Investigations and Theoretical Approaches
9.19
Anabranching and Anastomosing Rivers
9.20
Step–Pool Channel Features
9.21
Pool–Riffle
BC Eaton
GC Nanson
313
330
AE Zimmermann
346
DM Thompson
364
Fluvial Landforms
9.22
Fluvial Terraces
FJ Pazzaglia
379
9.23
Waters Divided: A History of Alluvial Fan Research and a View of Its Future
JD Stock
413
Paleohydrology
9.24
Quantitative Paleoflood Hydrology
G Benito and JE O’Connor
459
9.25
Outburst Floods
9.26
Global Late Quaternary Fluvial Paleohydrology: With Special Emphasis on Paleofloods and
Megafloods
VR Baker
JE O’Connor, JJ Clague, JS Walder, V Manville, and RA Beebee
475
511
Specific Fluvial Environments
9.27
Steep Headwater Channels
9.28
Bedrock Rivers
9.29
Incised Channels: Disturbance, Evolution and the Roles of Excess Transport Capacity and
Boundary Materials in Controlling Channel Response
A Simon and M Rinaldi
574
9.30
Streams of the Montane Humid Tropics
595
9.31
Dryland Fluvial Environments: Assessing Distinctiveness and Diversity from a Global
Perspective
S Tooth
612
Large River Floodplains
645
9.32
M Church
528
KX Whipple, RA DiBiase, and BT Crosby
550
FN Scatena and A Gupta
T Dunne and RE Aalto
Techniques of Study
9.33
Field and Laboratory Experiments in Fluvial Geomorphology
E Wohl
679
9.34
Numerical Modeling in Fluvial Geomorphology
TJ Coulthard and MJ Van De Wiel
694
9.35
Remote Data in Fluvial Geomorphology: Characteristics and Applications
T Wasklewicz, and YS Hayakawa
T Oguchi,
711
Management and Human Effects
9.36
Geomorphic Classification of Rivers
9.37
Impacts of Land-Use and Land-Cover Change on River Systems
9.38
Flow Regulation by Dams
9.39
Urbanization and River Channels
9.40
Impacts of Humans on River Fluxes and Morphology
JPM Syvitski
9.41
JM Buffington and DR Montgomery
LA James and SA Lecce
FJ Magilligan, KH Nislow, and CE Renshaw
A Chin, AP O’Dowd, and KJ Gregory
730
768
794
809
I Overeem, AJ Kettner, and
Geomorphologist’s Guide to Participating in River Rehabilitation
828
GB Pasternack
843