Wondzell and Gooseff: Treatise in Fluvial Geomorpholgy – Geomorphic Controls on Hyporheic Exchange 1 9.13 Geomorphic Controls on Hyporheic Exchange Across Scales - Watersheds to Particles
2 3 Steven M. Wondzell
4 5 6 7 8 9 10 U.S. Forest Service,
Pacific Northwest Research Station,
Olympia Forest Sciences Laboratory,
Olympia, WA 98512 USA.
Phone: 360-753-7691
E-mail: swondzell@fs.fed.us
11 Michael N. Gooseff
12 13 14 15 16 Civil & Environmental Engineering Department,
Pennsylvania State University,
University Park, PA 16802 USA
Phone: 814- 867-0044
E-mail: mgooseff@engr.psu.edu
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Wondzell and Gooseff: Treatise in Fluvial Geomorpholgy – Geomorphic Controls on Hyporheic Exchange 17 18 19 Abstract
20 We use geomorphology as a framework to understand hyporheic process and how these
21 processes change with location within a stream network, and over time in response to changes in
22 stream discharge and catchment wetness. We focus primarily on hydostatic and hydrodynamic
23 processes – the processes where linkages to fluvial geomorphology are most direct. Hydrostatic
24 processes result from morphologic features that create elevational head gradients whereas
25 hydrodynamic processes result from the interaction between stream flow and channel
26 morphologic features. We provide examples of the specific morphologic features that drive or
27 enable hyporheic exchange and we examine how these processes interact in real stream networks
28 to create complex subsurface flow nets through the hyporheic zone.
We examined the relationship between fluvial geomorphology and hyporheic exchange flows.
29 30 31 Key words
32 33 Hyporheic, step-pool sequence, pool-riffle sequence, meander bends, back channels, floodplain
34 spring brooks, mid-channel islands, stream bedforms, pumping exchange, saturated hydraulic
35 conductivity.
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Wondzell and Gooseff: Treatise in Fluvial Geomorpholgy – Geomorphic Controls on Hyporheic Exchange 36 9.13.1. Introduction
37 38 Hyporheic exchange flow (HEF) is the movement of stream water from the surface channel into
39 the subsurface and back to the stream (Figure 1). Stream water in hyporheic flow paths may mix
40 with groundwater so that the relative proportion of stream-source water in the hyporheic zone is
41 highly variable, ranging from 100% stream water to nearly 100% groundwater. Also the
42 residence time distribution of stream water in the hyporheic zone tends to be highly skewed, with
43 most of the stream water moving along short flow paths and thus having short residence times
44 (hours), but some water either moving on long flow paths or encountering relatively immobile
45 regions having very extended residence times (weeks to months, or longer). The boundaries of
46 the hyporheic zone are arbitrary, usually defined by the amount of stream-source water present in
47 the subsurface. Triska et al. (1989) set a threshold of 10% stream-source water to define the
48 limits of the hyporheic zone so that regions with <10% stream-source water were defined as
49 groundwater. Alternatively, the extent of the hyporheic zone can be delimited by water residence
50 time, for example, the subsurface zone delineated by hyporheic exchange flows with residence
51 times less than 24 hours (the 24-h hyporheic zone; Gooseff, in press).
52 53 The objective of this chapter is to examine the relation between geomorphology and hyporheic
54 processes. The two primary controls on hyporheic exchange are the gradients in total head
55 established along and across streambeds and the hydraulic conductivity of the streambed and
56 adjacent aquifer, both of which are significantly influenced by geomorphology. Total head (also
57 known as potential) is the sum of pressure head, elevation head, and velocity head. Pressure head
58 represents height of a column of fluid to produce pressure. Velocity head represents the vertical
59 distance needed for the fluid to fall freely (neglecting friction) to reach a particular velocity from
60 rest. Elevation head represents the potential energy of a fluid particle in terms of its height from
61 reference datum. Hydrostatic head is referred to as the sum of elevation and pressure head.
62 Groundwater tables in unconfined aquifers represent the spatial gradients in hydrostatic head. A
63 number of processes either drive or enable HEF, several of which are based on changes in head
64 gradients. We follow the organizational structure presented by Käser et al. (2009), who divided
65 these processes into five distinct classes:
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Wondzell and Gooseff: Treatise in Fluvial Geomorpholgy – Geomorphic Controls on Hyporheic Exchange 67 1. Transient exchange – the temporary movement of stream water into stream banks due to
68 short-term increases in stream stage (i.e., bank storage processes due to changes in
69 hydrostatic head gradients between stream and lateral riparian aquifer; Lewandowski et al.
70 2009; Sawyer et al. 2009a).
71 72 2. Turn-over exchange – the trapping of stream water in the streambed during times of
73 significant bed mobility (Elliot and Brooks, 1997b; Packman and Brooks 2001).
74 75 3. Turbulent diffusion – exchange driven by slip velocity that is created at the surface of the
76 porous medium of the bed where streamwise velocity vectors continue to propagate into the
77 surface layers of the bed (Packman and Bencala, 2000).
78 79 4. Hydrostatic-driven exchange – exchange driven by static hydraulic gradients which are
80 determined by changes in water surface elevation (Harvey and Bencala, 1993), spatial
81 heterogeneity in saturated hydraulic conductivity, or changes in the saturated cross-sectional
82 area of floodplain alluvium through which hyporheic flow occurs.
83 84 5. Hydrodynamic-driven exchange – exchange driven by the velocity head component of the
85 total head gradient on the bed surface (i.e., pumping exchange; Elliott and Brooks, 1997a,b)
86 and exchange induced by momentum gradients across beds and banks.
87 88 These classes of HEF processes are coupled to geomorphic processes in many ways. This is most
89 obvious for hydrostatic effects, which are directly dependent on channel and valley-floor
90 morphology and the depositional environment that controls spatial heterogeneity in saturated
91 hydraulic conductivity (K). However, turnover of streambed sediment is also related to fluvial
92 geomorphic processes. Similarly, hydrodynamic effects result from the interaction of flow over
93 stream bedforms. Geomorphic processes build stream bedforms and determine channel
94 morphology, especially longitudinal gradient, bed roughness, and water depth all of which
95 influence flow velocity. The relationship between geomorphology and the other classes of
96 processes is less direct, but still plays a role in controlling these processes through channel form
97 and the size distribution of sediment that makes up the streambed. This chapter focuses primarily
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Wondzell and Gooseff: Treatise in Fluvial Geomorpholgy – Geomorphic Controls on Hyporheic Exchange 98 on the hydostatic and hydrodynamic processes where linkages to geomorphic processes are most
99 direct.
100 101 We organize our discussion of the interactions between geomorphology and HEF using a
102 hierarchical scaling framework developed for river networks (Frissell et al. 1986; Bisson and
103 Montgomery, 1996), starting at the whole network, through the stream segment, to the stream
104 reach, to the channel unit, and down to the sub-channel unit scale. We recognize that describing
105 any given process or related flow path at a single “scale” is somewhat arbitrary because of the
106 nested structure of the hyporheic flow net and dispersion among HEF flow paths. Despite that,
107 the concept of scale is an important heuristic tool to organize our understanding of hyporheic
108 processes. In many senses, the reach scale is the most informative scale at which to consider
109 HEF. A single reach, by definition, has characteristic channel morphology so that the factors
110 driving HEF within the reach are relatively consistent. However, only a few of the geomorphic
111 factors driving HEF actually operate at this scale. Most of the drivers work at the channel unit or
112 smaller scales. And to understand the importance of HEF in stream ecosystem processes, the
113 cumulative effects of HEF must be evaluated at scales much larger than a single reach.
114 115 9.13.2. The effect of geomorphology on hyporheic exchange flows
116 117 9.13.2.1. The whole network to segment scale
118 119 The geologic setting of the stream network is an important factor determining the likely
120 occurrence of HEF, but there have been few attempts to study HEF at this broad scale. Rather,
121 our expectations are pieced together by drawing comparisons among HZ studies that have been
122 conducted in widely varying geologic settings, at different locations in the stream network, or
123 under widely varying flow conditions. We expect that geomorphic-hyporheic relationships will
124 differ substantially among different geologic settings.
125 126 Fluvial geomorphic studies have examined the factors that determine the types of channel
127 morphologies present within stream networks (Montgomery and Buffington, 1997; Wohl and
128 Merritt, 2005; Brardinoni and Hassan, 2007). Montgomery and Buffington (1997) presented one
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Wondzell and Gooseff: Treatise in Fluvial Geomorpholgy – Geomorphic Controls on Hyporheic Exchange 129 such description of the distribution of channel morphologies typical of many mountainous
130 landscapes. They showed that catchment area and channel longitudinal gradient controlled the
131 development of distinct channel types such that the channel types tended to follow a
132 characteristic sequence within a catchment (Figure 2A). In their example, this sequence starts
133 with bedrock and colluvial channels in the steepest, upper-most headwaters. As longitudinal
134 gradients decrease, channels change to cascades, to step-pool, to plane-bed, to pool-riffle, and the
135 largest, lowest gradient rivers were typified as dune-ripple channels. Along with these changes in
136 channel morphology, the following would be expected: decreased longitudinal gradient and
137 mean grain size of streambed sediment, and increased depth, width, hydraulic radius, and flow
138 velocity (Leopold and Maddock, 1953; Wohl and Merritt, 2008).
139 140 In this paper, we use Montgomery and Buffington’s (1997) description of the sequence of
141 channel types within a catchment as a simple heuristic model to organize our examination of the
142 relative importance of the different processes that drive HEF within stream networks. We
143 recognize that local controlling factors often interrupt simple sequencing of channel types. For
144 example, landslides may block large mainstem channels, creating locally steep gradients over the
145 landslide debris and uncharacteristically low gradients in the depositional reach immediately
146 upstream (Benda et al. 2003). We also recognize that regional differences in geology and
147 geomorphology will lead to dramatically different spatial organization of channel types (see for
148 example characteristic channel type in glaciated mountainous regions as described by Brardinoni
149 and Hassan, 1997). Our descriptions of the spatial organization of stream types and the resulting
150 HEF processes will have to be modified for any specific landscape.
151 152 Most hyporheic exchange results from head gradients pushing water through the streambed. The
153 amount of stream water entering the hyporheic zone is thus a function of the steepness of the
154 head gradient and the saturated hydraulic conductivity of the streambed and underlying aquifer.
155 The head gradients can be induced in many ways, but the two of primary influence are the
156 hydrostatic and hydrodynamic processes. The relative importance of each of these processes is
157 expected to vary among channel types and with longitudinal gradient. In high gradient streams,
158 channel forms such as step-pool sequences or pool-riffle sequences can create very steep
159 hydrostatic head gradients. Further, because of high bed roughness and relatively shallow water
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Wondzell and Gooseff: Treatise in Fluvial Geomorpholgy – Geomorphic Controls on Hyporheic Exchange 160 depth, flow velocities tend to be lower in small steep streams than in larger, low gradient streams
161 (Leopold and Maddock, 1953; Wondzell et al. 2007). In contrast, it is difficult for natural
162 processes to create steep changes in the longitudinal gradient in low gradient streams. Instead,
163 stream flow interacts with stream bedforms, such as dunes or ripples, such that hydrodynamic
164 forces dominate the development of head gradients through the streambed. Thus, we expect that
165 hydrostatic effects will dominate in high gradient channels and that hydrodynamic processes will
166 dominate in low gradient channels (Figure 2B). Further, because channel types and longitudinal
167 gradients generally vary systematically within stream networks, we further expect that
168 hydrostatic effects will tend to dominate in the upper portions of stream networks and that the
169 relative importance of hydrodynamic processes will increase down the stream network.
170 171 9.13.2.2. The reach scale – setting the potential for hyporheic exchange
172 173 The potential for HEF to occur varies within any given stream reach. Roughly speaking, this
174 potential is determined by the factors that generate head differences that drive HEF, the
175 properties of the subsurface alluvium through which HEF occurs, and the potential effect of
176 lateral groundwater inputs from adjacent hillslopes that might limit hyporheic expression.
177 178 9.13.2.2.1. Losing and gaining reaches
179 180 Hyporheic exchange is likely to be more limited in strongly gaining reaches than in neutral
181 reaches because of steep streamward hydrologic gradients surrounding the channel (Wroblicky et
182 al. 1998; Storey et al. 2003; Malcolm et al. 2003 and 2005; Cardenas, 2009). Similarly, where
183 water is lost to regional aquifers in strongly losing reaches, return flows of stream water back to
184 the stream are likely to be severely restricted and thus also limit the expression of the hyporheic
185 zone (Cardenas, 2009). These patterns of gains and/or losses are controlled, at some level, by
186 regional groundwater and catchment characteristics interacting with smaller scale effects. In
187 large gaining rivers, Larkin and Sharp (1992) demonstrated that the relative dominance of cross-
188 valley vs. down valley flow paths through valley-floor aquifers varied depending on the
189 longitudinal gradient of the valley floor and the hydraulic conductivity of the valley floor
190 alluvium. In higher gradient reaches (>0.004 m/m) and in areas with coarser substrate, flow was
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Wondzell and Gooseff: Treatise in Fluvial Geomorpholgy – Geomorphic Controls on Hyporheic Exchange 191 predominantly down valley. Conversely, where valley floor gradients were shallower or
192 sediment more finely textured, flow tended to be toward the stream. Thus, the way in which
193 lateral inputs influence hyporheic exchange is not solely a function of their magnitude, but also a
194 function of the ability of subsurface water to move down-valley (Storey et al. 2003). The ratio
195 between these two factors – the magnitude of the inputs relative to down valley flow –
196 determines how hyporheic exchange is affected.
197 198 As a first approximation, the potential for down valley flow can be estimated using the
199 relationships summarized in Darcy’s Law – that is, the product of the longitudinal valley
200 gradient, the saturated cross-sectional area of the floodplain perpendicular to the direction of
201 subsurface flow, and the hydraulic conductivity of the alluvium. As lateral inputs increase,
202 several factors may change: (1) water tables may rise, thus increasing the saturated thickness and
203 the cross-sectional area through which water flows allowing the transmission of more water, or
204 (2) flow paths may begin to turn obliquely toward the stream, which also increases the saturated
205 cross-sectional area and may also increase head gradients. Consequently, under dry conditions
206 when lateral inputs are relatively small, the potential extent and magnitude of hyporheic
207 exchange can be fully expressed (Figure 3A). As subsurface flows turn toward the channel they
208 begin to limit the extent of the hyporheic zone with only minor effect on the HEF (Wondzell and
209 Swanson, 1996). If sufficiently large, lateral inputs can severely limit both the spatial extent and
210 magnitude of hyporheic exchange (Figure 3B; Harvey and Bencala, 1993; Wroblicky et al. 1998;
211 Storey et al. 2003; Cardenas and Wilson, 2007; Malcolm et al. 2003, Soulsby et al. 2009).
212 213 Simple generalizations of where and when lateral inputs will limit HEF are difficult because of
214 the wide range of geomorphic settings in which HEF occurs and because the magnitude of lateral
215 inputs changes with catchment wetness. Lateral inputs are expected to be high when catchments
216 are wet and decrease as catchments dry out. However, lateral inputs are not spatially uniform. In
217 steep mountainous settings, the size of the upslope area draining directly to the valley floor is
218 important, concentrating lateral inputs in zones at the base of hillslope hollows (Jencso et al.
219 2009). Lateral inputs may persist the entire year at the bases of the largest hillslope hollows.
220 Most hillslope hollows are small, however, so that most of the stream network would be
221 disconnected from lateral inputs except for short periods of time when catchments are very wet,
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Wondzell and Gooseff: Treatise in Fluvial Geomorpholgy – Geomorphic Controls on Hyporheic Exchange 222 for example after large storms or during peak snowmelt. We are unaware of similar studies
223 relating topography to spatial patterns of hillslope inputs in areas of low relief with humid
224 climates. However, Storey et al. (2003) reported that an extensive shallow surfical aquifer was
225 present along their lowland, low-gradient study reach and that lateral inputs of groundwater
226 substantially reduced both the extent and the amount of hyporheic exchange flows except during
227 summer baseflow. Clearly, the influence of lateral inputs may be much different in lowland
228 catchments than in steep mountainous catchments.
229 230 Changes in lateral inputs to streams do not occur in isolation. Rather, they are likely to be
231 accompanied by corresponding changes in stream stage (and discharge). The change in water
232 table elevations resulting from changed lateral inputs must be considered relative to the
233 accompanying changes in stream stage. Although the number of studies examining changes in
234 hyporheic flow paths with changing catchment wetness is limited, studies in small mountain
235 streams suggest that water table elevations in the floodplain increase more than stream stage so
236 that HEF is typically more restricted when catchments are wet (Figure 4A and 4B; Harvey and
237 Bencala, 1993; Wondzell and Swanson, 1996; Stednick and Fernald, 1999). Storey et al. (2003)
238 reported similar results for a lowland, low-gradient river.
239 240 In some cases, however, stream stage may change markedly without corresponding changes in
241 precipitation recharge or changes in lateral inputs. Most examples of these processes come from
242 large, lowland rivers because river stage is controlled by processes far upstream. These “bank
243 storage” processes (Pinder and Sauer, 1971) have been recognized as a form of transient
244 hyporheic exchange (Figure 4C and 4D) that can result from both in-bank or over-bank floods
245 (Bates et al. 2000; Burt et al. 2002). In some situations, increased stream stage may even lead to
246 groundwater ridging in the floodplain, reversing head gradients and limiting lateral groundwater
247 inputs. Similarly, hyporheic exchange through stream banks can result from diel variations in
248 stream stage (and discharge) during snow melt periods (Loheide and Lundquist, 2009) or from
249 tidally induced changes in water elevations in coastal streams and rivers (Bianchin et al. in
250 press).
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Wondzell and Gooseff: Treatise in Fluvial Geomorpholgy – Geomorphic Controls on Hyporheic Exchange 252 Transient hyporheic exchange may be especially evident in regulated rivers where releases from
253 dams (or other control structures) can result in large and rapid changes in river stage without
254 corresponding local precipitation to recharge floodplain aquifers (e.g., Fritz and Arntzen, 2007;
255 Lewandowski et al. 2009; Sawyer et al. 2009a; Francis et al., in press). However, transient
256 hyporheic exchange may not always result from fluctuations in river stage. For example,
257 Hanrahan (2008) studied vertical HEF through the streambed of a large, regulated gravel bed
258 river where stage sometimes changed by nearly 2 m in an hour. For the most part, they did not
259 observe transient hyporheic exchange related to changes in stage. They concluded that
260 hydrostatic and hydrodynamic processes remained the dominant control on HEF. Notably,
261 Hanrahan (2008) did not examine lateral exchanges through the stream banks, which can be
262 more responsive to changes in stage than are locations in the stream channel itself (Storey et al.
263 2003). Water table fluctuations in the floodplain at long distances from the stream are not
264 necessarily indicative of extensive HEF because pressure fluctuations can propagate through
265 surficial (unconfined) aquifers much faster than does the actual flow of stream water. This was
266 clearly demonstrated by Lewandowski et al. (2009) who showed that river water penetrated, at
267 most, only 4 m into the stream bank even though water table fluctuations were observed more
268 than 300 m from the river.
269 270 HEF can occur in strongly gaining and losing reaches because of the nested structure of
271 hyporheic flow paths, and because HEF can occur at a variety of spatial scales. Thus an envelope
272 of the HZ can be set within larger non-hyporheic flow paths (Figure 3B; Cardenas and Wilson,
273 2007). Similarly, smaller-scale HEF can occur as a result of smaller scale geomorphic drivers,
274 even within a reach that is, overall, strongly losing (Payn et al. 2009). Further, because HEF is
275 dominated by relatively near-stream flow paths that are short in length and residence time
276 (Kasahara and Wondzell, 2003), the magnitude of HEF can be substantial, even in strongly
277 gaining reaches where the spatial extent of the hyporheic zone is greatly restricted (Wondzell and
278 Swanson, 1996; Cardenas and Wilson, 2007; Payn et al. 2009).
279 280 9.13.2.2.2. Changes in saturated cross-sectional area
281 10
Wondzell and Gooseff: Treatise in Fluvial Geomorpholgy – Geomorphic Controls on Hyporheic Exchange 282 The saturated cross sectional area of the floodplain (orthogonal to groundwater flow path
283 direction) is one of the factors determining the amount of groundwater transmitted down valley
284 through the valley floor alluvium. Thus, any change in the cross sectional area along the length
285 of a stream reach will lead to parallel changes in the down valley flow of water through the
286 floodplain, thereby driving downwelling from, or upwelling to the stream (Stanford and Ward,
287 1993). Downwelling occurs where valley floors increase in width, for example, downstream of
288 bedrock-constrained reaches (Figure 5A; Poole et al. 2004 and 2006; Acuna and Tockner, 2009).
289 Conversely, upwelling occurs where valley floors narrow at the lower end of wide unconstrained
290 reaches (Figure 5A; Baxter and Hauer, 2000; Acuna and Tockner, 2009). Similarly, variations in
291 the thickness of the surficial aquifer, caused by variations in depth to bedrock or other confining
292 layers drive similar patterns of upwelling and downwelling. For example, upwelling commonly
293 occurs just upstream of bedrock sills with a subsequent transition to downwelling just
294 downstream of such bedrock sills as the surficial aquifer again thickens (Figure 5B; Valett,
295 1993). This is easily observed in streams in arid regions during the dry season, where perennial
296 flow may only occur above bedrock sills, which force the subsurface flow to the surface.
297 298 9.13.2.3. The sub-reach to channel-unit scale – hydrostatic processes
299 300 Geomorphic features of the stream channel and valley floor within stream reaches control the
301 elevation of surface water and can thereby create significant head gradients through the valley
302 floor alluvium, driving HEF. Because these geomorphic features are static on the time scales
303 typical of hyporheic exchange (hours to weeks) they are broadly recognized as “hydrostatic
304 processes”.
305 306 9.13.2.3.1. Step-pool and pool-riffle sequences
307 308 One of the best-studied examples of hydrostatic processes involves the changes in water surface
309 elevation along a pool-step sequence and the resulting head gradients that drive HEF (Figure 1;
310 Harvey and Bencala, 1993). Harvey and Bencala (1993) showed that the change in the
311 longitudinal gradient of the stream channel (which approximates the stream energy profile) drove
312 HEF. They also observed that HEF flow paths tended to be curved – first curving away from the
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Wondzell and Gooseff: Treatise in Fluvial Geomorpholgy – Geomorphic Controls on Hyporheic Exchange 313 stream above the step or riffle and then curving back to the stream below the step or riffle.
314 Building from their observations, model analyses show that along an idealized straight channel
315 with homogeneous isotropic porous sediment, hyporheic flow paths around a change in the
316 longitudinal gradient will exploit the full 3-dimensional saturated volume along the channel, thus
317 extending both vertically beneath the streambed and horizontally through the streambanks and
318 near stream aquifer (Figure 1A and 1B). Real streams are substantially more complicated,
319 however, such that changes in hydraulic conductivity of the alluvium, bends in the channel, and
320 the spatial location of lateral groundwater inputs lead to the development of a complicated flow
321 net through the valley floor (e.g., Cardenas and Zlotnik, 2003). Despite these complexities, the
322 steepness of the hydraulic head gradient imposed by the change in the longitudinal gradient and
323 the saturated hydraulic conductivity control the amount of stream water exchanged with the
324 subsurface.
325 326 Many factors can modify the effect of steps or riffles on HEF. For example, the height of the step
327 (or steepness of the riffle) determines the head gradient available to drive HEF so that a single
328 very large step has the potential to drive more HEF than if the same amount of elevational
329 change is spread over several smaller steps (Kasahara, 2000). Because of this, large wood can be
330 important in determining the amount of HEF in forest streams. Single logs tend to create
331 frequent, small obstructions that collect and store small amounts of sediment, forming pool-step
332 sequences in which the extent of the hyporheic zone tends to be small (Wondzell, 2006).
333 Although log jams are less common, they can create large obstructions storing sediment in
334 wedges several meters deep and 10 or more meters in length, and significantly widen constrained
335 stream channels. Consequently, log jams can form extensive hyporheic zones in steep, confined
336 mountain streams (Wondzell, 2006).
337 338 Large, channel-spanning logs can wedge into steep narrow channels, forcing the accumulation of
339 sediment in channels, converting bedrock reaches to alluvial reaches with a step-pool
340 morphology (Montgomery et al. 1996), thereby greatly enhancing HEF. Similarly, large wood
341 can force plane-bed channels into a pool-riffle morphology (Montgomery et al. 1996) which
342 should lead to more HEF than would be present in a comparable wood-free channel. Large wood
343 can have the opposite effect in channels that would have a free-formed pool-riffle morphology.
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Wondzell and Gooseff: Treatise in Fluvial Geomorpholgy – Geomorphic Controls on Hyporheic Exchange 344 In one documented case, accumulations of large wood tended to force a pool-riffle channel
345 toward a step-pool morphology (Wondzell et al. 2009). The channel adjusted to removal of all
346 large, in-stream wood by developing a better defined pool-riffle structure around meander bends,
347 leading to increased sediment storage. Continued channel adjustment over time following the
348 removal of large wood eventually led to substantial increases in HEF.
349 350 The size, spacing, and sequence of channel units (e.g., pools and riffles) along the stream
351 longitudinal profile can also affect HEF (Anderson et al. 2005; Gooseff et al. 2006). Anderson et
352 al. (2005) made detailed measurements of channel profiles and patterns of HEF, and showed that
353 channel unit size and spacing increased as did the length of channel characterized by
354 downwelling with increasing drainage area in a mountainous stream catchment. Gooseff et al.
355 (2006) built on these results, examining HEF using 2-D groundwater models of idealized
356 longitudinal profiles of mountain streams. Gooseff et al.’s (2006) modeling results confirmed
357 that both channel unit spacing and size were important in determining hyporheic exchange
358 patterns of upwelling and downwelling. Perhaps more surprising, however, was the observation
359 that the sequence of channel units also affected simulated HEF. Gooseff et al. (2006) compared
360 pairs of idealized stream reaches that varied only by the way the longitudinal gradient changed
361 over the pool-riffle sequence – i.e., the slope of the riffle was gradual on its upstream end and
362 steepest at its downstream end (described as a pool–riffle–step sequence) versus riffles that were
363 initially steep with the slope decreasing toward the downstream end (described as a pool–step–
364 riffle sequences). Simulated downwelling lengths were substantially longer for pool–riffle–step
365 sequences than for pool–step–riffle sequences.
366 367 9.13.2.3.2. Meander bends and point bars
368 369 A variety of channel and valley floor morphologic features, in addition to changes in the
370 longitudinal gradient, create head gradients with the potential to drive HEF. These include
371 channel meander bends and associated point bars, back channels or floodplain spring brooks, and
372 islands set between main and secondary channels. In all these cases, differences in the
373 elevational head of surface water between two channels, between different points in a single
374 channel around a meander bend, or between points on opposite sides of an island create head
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Wondzell and Gooseff: Treatise in Fluvial Geomorpholgy – Geomorphic Controls on Hyporheic Exchange 375 gradients that drive HEF. For example, head gradients through the point bar in a meander bend
376 are steeper than the longitudinal gradient of the stream channel around the point bar (Peterson
377 and Sickbert, 2006) so that stream water infiltrates the upper end of the point bar and is returned
378 to the channel at the lower end of the point bar (Figure 6A; Vervier and Naiman, 1992). More
379 generally, these exchange flows occur across the full length of meander bends and are influenced
380 by both the change in stream water elevation around the meander bend and the plan-view shape
381 of the meander bend. Highly evolved meander bends support steep head gradients across the
382 mender neck because of the close proximity of the stream channels (Figure 6B; Boano et al.
383 2006; Revelli et al. 2008) so that HEF is dominantly located in the meander neck, with much
384 reduced HEF across the remainder of the meander where head gradients are much lower. In other
385 cases, meanders develop a characteristic pattern of alternating pools and riffles, with riffles
386 located at the thalweg cross-overs in the inflections between adjacent meanders and pools or low
387 gradient runs wrapping around the point bar (Figure 6C). This combination of channel
388 morphologic features can create complex HEF flow paths within meander bends. The residence
389 times of HEF traversing meander bends can be quite short where meanders are small and
390 saturated hydraulic conductivities are high (Pinay et al. 2009). Conversely, residence times of
391 HEF may be extremely long in meander bends of low gradient rivers with fine textured sediment
392 (Boano et al. 2006; Peterson and Sickbert, 2006).
393 394 9.13.2.3.3. Back channels and floodplain spring brooks
395 396 Channel planforms are often complex in wide floodplains, including a network of old or
397 abandoned channels. If the upstream ends of these channels are plugged with sediment and if the
398 downstream ends are sufficiently incised to intercept the water table and are connected back to
399 the river at their downstream ends, they will act as drains, imposing head gradients from the
400 stream to the old channel (Figure 7A; Wondzell and Swanson, 1996; Poole et al. 2006). These
401 channels are also known as floodplain spring brooks because water upwells into the channel,
402 forming a spring at its head. In addition to creating HEF, these channels will capture whatever
403 water is in the surficial aquifer of the floodplain, including down valley flows from upstream
404 locations, and lateral inputs of groundwater or hillslope water from the valley margin. However,
405 because lateral inputs tend to be small and spatially isolated (Jencso et al. 2009; and as discussed
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Wondzell and Gooseff: Treatise in Fluvial Geomorpholgy – Geomorphic Controls on Hyporheic Exchange 406 above), floodplain spring brooks will most often be fed by HEF (Wondzell and Swanson, 1996;
407 Jones et al. 2007).
408 409 Abandoned channels can also be plugged at their downstream ends and open to the river at their
410 upstream ends. In this case, stream water can flow into the abandoned channel, infiltrate the
411 channel bed and raise the water table in the middle of the floodplain, thereby creating head
412 gradients and driving HEF from the abandoned channel back to the main stream channel (Figure
413 7B). More complex situations arise when the longitudinal gradients in either the back channel or
414 mainstem channel are interrupted by steeper riffles or steps. Figure 7C shows the interactions
415 between a back channel and riffle. Above the riffle, water in the main channel is higher than the
416 back channel so water flows towards the spring brook. Downstream of the riffle, the main
417 channel is lower than the back channel so that the back channel loses water over its downstream
418 extent, eventually going dry before reaching the main channel.
419 420 The channel planform features that drive HEF can occur over a range of spatial scales, and their
421 influence may change through time as the stage height of water in the main channel changes. For
422 example, a small gravel bar may have low points along the stream bank. At high stage, the entire
423 gravel bar may be submerged. As stage decreases the center of the bar may become exposed,
424 creating a secondary channel along the bank. As stage decreases further, flow may become
425 discontinuous through the secondary channel such that it functions as a drain if it is plugged at
426 the upstream end, or functions as a conduit allowing stream water to infiltrate the surface of the
427 gravel bar if it is plugged at its downstream end. Old channels in large floodplains may act
428 similarly, with continuous flow along their full length during floods, but becoming disconnected
429 at intermediate to low stage, or even dry completely during periods of minimum discharge. In
430 large floodplain reaches, these channels can be 100’s of meters to kilometers in length, extending
431 nearly the full length of the stream reach (Poole et al. 2006; Arrigoni et al. 2008).
432 433 9.13.2.3.4. Secondary channels and islands
434 435 Islands present a special case of back channels in which the channel is continuously connected to
436 the main channel over its full length. Hyporheic hydrology of islands has not been extensively
15
Wondzell and Gooseff: Treatise in Fluvial Geomorpholgy – Geomorphic Controls on Hyporheic Exchange 437 studied. However, we expect that the surface water elevations in channels bounding the island
438 create boundary conditions for total head and control HEF through islands as is generally
439 indicated by the available literature (Dent et al. 2007; Francis et al., in press). If channels along
440 both sides of the island are parallel and symmetric with constant longitudinal gradient, then flow
441 through the island will parallel the channels and the head gradient driving flow will equal the
442 overall longitudinal gradient of the stream reach (Figure 8A). If riffles are present in the
443 channels, the head gradient through the island adjacent to the riffles can be much steeper than the
444 reach averaged longitudinal gradient (Figure 8B). Also, if riffles are displaced along the primary
445 and secondary channels surrounding an elongated island such that a riffle is located near the head
446 of the island in one channel and near the tail of the island in the second channel, the resulting
447 head gradients would tend to drive flows laterally through the island, leading to very large cross-
448 sectional areas experiencing HEF, and therefore large amounts of HEF, albeit, with shorter
449 length flow paths (Figure 8C). While islands may be uncommon in most channel types, they may
450 dominate HEF in braided and anastomosing stream reaches (Ward et al. 1999; Arscott et al.
451 2001). Given the complexities of potential sizes and shapes of islands and patterns in
452 longitudinal gradients in the bounding channels, the resulting flow nets, residence times, and
453 amounts of HEF are likely to vary widely.
454 455 9.13.2.3.5. Spatial heterogeneity in saturated hydraulic conductivity
456 457 Fluvial processes control the depositional environment on the streambed and across the
458 floodplain creating spatial heterogeneity in the texture of deposited and reworked sediment
459 across a range of scales, from the surface of the streambed to the entire floodplain. Because
460 sediment texture is closely related to saturated hydraulic conductivity (K), these processes can
461 substantially influence HEF. However, because of the difficulties in quantifying these patterns at
462 the scales at which they influence HEF, they have been relatively little studied. At fine scales,
463 streambed roughness can control the depositional environment across the streambed (Buffington
464 and Montgomery, 1999), which lead to spatial patterns in the distribution of K within the
465 streambed (Genereaux et al. 2008), which in turn can influence both the location and amount of
466 HEF. HEF will be restricted where the streambed is clogged with fine sediment and
467 preferentially located in zones with higher K. Experiments in flumes have also shown that HEF
16
Wondzell and Gooseff: Treatise in Fluvial Geomorpholgy – Geomorphic Controls on Hyporheic Exchange 468 can also influence patterns of fine-sediment deposition, with fine sediment preferentially
469 deposited in downwelling zones (Packman and MacKay, 2003; Rehg et al. 2005) which may
470 explain differences in K between upwelling and downwelling zones observed in a steep
471 headwater stream (Scordo and Moore, 2009).
472 473 Spatially heterogeneous patterns in K influence HEF. For example, groundwater flow modeling
474 studies using homogeneous vs. heterogeneous K showed that spatial heterogeneity may add
475 substantial complexity to the spatial patterns of the hyporheic flow net (Woessner 2000). When
476 relatively high K regions are aligned parallel with head gradients they create preferential flow
477 pathways (Wagner and Bretschko, 2002) that can increase the total amount of HEF (Cardenas
478 and Zlotnik, 2003; Cardenas et al. 2004). Results from Cardenas et al. (2004) showed that
479 influence of heterogeneity in K was relatively greater in lower gradient streams and where head
480 gradients driving HEF were reduced. To our knowledge, the influence of fine-grained
481 heterogeneity has not been studied in steeper channels where hydrostatic processes dominate.
482 483 Fluvial processes also influence spatial patterns in K at the scale of the entire floodplain.
484 Especially important is the layering of stream and floodplain alluvium. Layering can create
485 strong vertical anisotropy (Chen, 2004), limiting vertical exchange and promoting lateral flows
486 through the streambed and floodplain (Packman et al. 2006; Marion et al. 2008). Overbank
487 deposition can also bury back channels creating “paleochannels” where coarse streambed
488 alluvium is buried under finer floodplain soils (Stanford and Ward, 1993; Stanford et al. 1994;
489 Poole et al. 2004). If these paleochannels intercept the water table, they will function as large
490 preferential-flow pathways that can route water the full length of a floodplain. In this regard they
491 function much like a subsurface version of back channels or floodplain spring brooks – either
492 acting as drains lowering the water table in the floodplain and imposing head gradients from the
493 stream to the paleochannel, or acting as distributaries, routing water into the floodplain and
494 imposing head gradients from the paleochannel to the stream. Locations of paleochannels are
495 sometimes evident from shallow depressions along the floodplain. In other cases, over-bank
496 deposition will have completely filled old channels so that there is no surficial indication on the
497 flat floodplain surface. The influence of paleochannels is difficult to discern because networks of
17
Wondzell and Gooseff: Treatise in Fluvial Geomorpholgy – Geomorphic Controls on Hyporheic Exchange 498 widely spaced wells are unlikely to find and trace the location of these features along the length
499 of the floodplain. As a consequence, their influence on HEF has not been widely studied.
500 501 9.13.2.4. The bedform scale – hydrodynamic processes
502 503 Channel hydraulics, and the spatial and temporal distribution of velocity (kinetic energy) across
504 streambeds are significantly influenced by the form of the channel and the bedforms that occur in
505 channels. The continuous feedback between pressure distribution and shear stress across the bed
506 surface and the potential to erode the bed will cause turn-over exchange to occur during times of
507 high flows. During lower flows, when bed sediment is relatively stable, bedforms cause some
508 level of form drag on the flows, inducing pressure distributions across the bedforms, thereby
509 driving HEF at a scale smaller than the bedform (Figure 9). The size of the bedform is set by
510 both the energy regime of the reach and the material that makes up the reach, and the form drag
511 induced on the water column by the bedform is of course partly controlled by its size. Thus, the
512 scale of HEF flowpaths induced by hydrodynamic exchange across the bedforms will scale in
513 part with the size of bedforms present (Cardenas et al. 2004). Finally, the heterogeneity of the
514 bed material that makes up the bedforms will have a distinct control on the flux rate and actual
515 flowpaths through and around the bedforms (Sawyer et al. 2009b).
516 517 In sand bed streams, hydrodynamic HEF has been extensively studied both theoretically and
518 empirically. Typical bed forms in sand bed streams are dunes and ripples, which have a fairly
519 predictable geometry and spacing, based on bed sediment composition and flow rate.
520 Thibodeaux and Boyle (1987) pioneered investigations of the hydrodynamic pressure
521 distribution across dunes, noting the penetration of channel water into the porous bed forms.
522 Further development of a ‘pumping exchange’ model by Elliot and Brooks (1997a,b) expanded
523 the ability to predict HEF and associated solute dynamics in channel-bed systems. Whereas most
524 studies of hydrodynamic exchange processes were generally carried out in or applied to flume
525 studies, there has been at least one application of incorporating the pumping exchange model to
526 tracer transport in field studies. Salehin et al. (2003) studied the transport of tracer along several
527 km of Sava Brook in Sweden and successfully applied a solute transport model to the observed
528 data to explain long time residence time distributions using the pumping exchange model theory.
18
Wondzell and Gooseff: Treatise in Fluvial Geomorpholgy – Geomorphic Controls on Hyporheic Exchange 529 The predictability of dune and ripple sizing and spacing makes the pumping exchange model a
530 useful tool to explore HEF in sand bed streams and rivers.
531 532 In gravel bed streams, bed form types may be generally predictable (i.e., Montgomery and
533 Buffington, 1997; Wohl and Merritt, 2008; Chin, 2002), but the exact geometry and spacing of
534 bed forms is less predictable, particularly at a scale that will directly influence head distributions
535 across and along the channel. Hence, the velocity distribution in the channel and around the bed
536 form, which contributes to hydrodynamic exchange, is also unpredictable. Tonina and
537 Buffington (2007) conducted careful studies of total pressure distribution across streambeds in
538 flumes that had ‘realistic’ geometry of a pool-riffle sequence in a gravel bed channel. Their
539 results indicated that total head distribution (i.e., incorporating velocity head in addition to
540 hydrostatic head) was important to exchange at focused points in the channel where high velocity
541 occurred. Further, they confirmed that in general, there was little or no contribution of velocity
542 head to parts of the bed that were overlain by deeper, slower flow, and therefore a hydrostatic
543 representation of exchange will likely be more applicable in these locations.
544 545 Regardless of the predictability of bed form geometry and spacing, the associated hydrodynamic
546 HEF may induce only limited lengths of exchange in the subsurface because much of the
547 exchange dynamics are expected to be vertical rather than lateral. Exchange lateral to the channel
548 is more likely to be driven by hydrostatic gradients set up across meander bends or bars (as
549 described above). Hydrodynamic HEF will contribute to, but be only one component of, total
550 HEF in natural channels, and its importance will be dictated by both channel hydraulics and, if
551 present, competing hydrostatic factors that can create steeper head gradients.
552 553 9.13.2.4. The particle scale – turbulent diffusion
554 555 At the particle scale on streambeds, turbulent diffusion is significantly influenced by the size and
556 arrangement of surface sediment. Because turbulent diffusion is induced by the momentum
557 transfer between the water column and the porous media, HEF due to turbulent diffusion is a
558 function of the decreasing velocity profile within the surface layers of the porous media (Shimizu
559 et al. 1990). Thus, the distribution of sediment at the surface will greatly influence the potential
19
Wondzell and Gooseff: Treatise in Fluvial Geomorpholgy – Geomorphic Controls on Hyporheic Exchange 560 for energy and mass transfer within this zone. Turbulent diffusion HEF is prominent in gravel
561 bed streams where surface pores are more likely to accommodate such open exchanges of
562 momentum across the bed (Tonina and Buffington, 2009). Beds composed of sand particle sizes
563 and smaller provide too much resistance to the momentum exchange between the water column
564 and the bed. Hence, turbulent diffusion is more likely to be an important component of HEF in
565 low order, high-gradient streams (Figure 2B). Careful theoretical and empirical research on
566 turbulent diffusion has been conducted largely on planar beds (Shimizu et al. 1990; Habel et al.
567 2002). Therefore, in the complex bed topography of typical gravel channels, turbulent diffusion
568 will be a component of HEF, likely not the singular driver of HEF.
569 570 9.13.3. Discussion
571 572 9.13.3.1 Multiple features acting in concert
573 574 In the examples presented above (Figures 1, 3–9), we have mostly focused on single types of
575 channel morphologic features that drive or enable hydrostatic and hydrodynamic HEF. However,
576 these features never occur in isolation. Rather, a single stream reach will typically contain many
577 of the morphologic features described above. Interactions among these features are likely to be
578 important in determining the actual HEF in any given stream reach. In some cases, the effects of
579 multiple features could be additive and result in higher HEF than if they did not co-occur. For
580 example, cross-valley flow paths between main channels and floodplain spring brooks can be
581 accentuated by riffles (Figure 7C). However, interactive effects could also cancel, for example
582 where riffles at the inflection points of meander bends reduce head gradients through point bars
583 (Figure 6C). The interactions between different processes driving or enabling HEF is complex,
584 and to some degree, site specific, making it difficult to quantify the effects of these interactions.
585 Because of these difficulties, there are relatively few comparative studies that have examined
586 multiple processes concurrently, within natural stream channels and attempted to evaluate the net
587 effect of each process on the total HEF within stream reaches.
588 589 Sensitivity analyses with groundwater flow models calibrated to simulate HEF in a studied
590 stream reach provide one opportunity to examine the relative importance of channel morphologic
20
Wondzell and Gooseff: Treatise in Fluvial Geomorpholgy – Geomorphic Controls on Hyporheic Exchange 591 features on HEF where multiple features are present in a single reach. For example, Kasahara
592 and Wondzell (2003) examined a number of channel morphologic features among stream reaches
593 of different sizes in a mountainous stream network under conditions of summer baseflow
594 discharge. In all cases, the single strongest driver determining the amount of HEF occurring
595 within the simulated stream reaches was the change in longitudinal gradient over step-pool
596 sequences in the 2nd-order channel (Figure 10A) and pool-riffle sequences in the 5th-order
597 channel. The shape of the hyporheic flow net in the 5th-order stream, however, was strongly
598 controlled by the presence, location, and relative elevation difference between water in the main
599 channel and the back channels (Figure 10B). Similarly, Cardenas et al. (2004) examined
600 sediment heterogeneity, size of bedforms, and both longitudinal and lateral head gradients in a
601 low gradient, sand bed stream. They found that HEF was greater where beforms had higher
602 amplitude and were more closely spaced. Spatial heterogeniety in K increased HEF relative to
603 homogeneous simulations, as did inclusion of lateral head gradients, but the effect was small
604 relative to the effect of the size and spacing of bedforms.
605 606 Channel morphologic features can interact with changes in steam stage and lateral groundwater
607 inputs in ways that can substantially influence the amount of HEF over time, across seasons or
608 within a single storm event. Storey et al. (2003) examined HEF in a pool-riffle sequence at both
609 high- and low-baseflow discharge. At high stage, the stream tended to “drown” the riffle,
610 substantially reducing the change in the longitudinal gradient over the pool-riffle sequence and
611 thus reducing HEF. In contrast, at low stage, the water surface more closely followed the
612 streambed topography, thus creating steeper head gradients that supported more HEF. Storey et
613 al. (2003) also showed that lateral inputs during the wet season were sufficient to eliminate most
614 of the HEF through the riffle. Cardenas and Wilson (2006) showed that low rates of groundwater
615 discharge limited the extent of the HZ formed by the hydrodynamics of stream bedforms, and
616 that high rates of groundwater discharge could completely eliminate HEF.
617 618 We know of only one study comparing the relative influence of hydrostatic and hydrodynamic
619 effects. In a flume, Tonina and Buffington (2007) investigated the control of total head (i.e.,
620 including dynamic head) in driving hyporheic exchange. Their results suggested that there are
21
Wondzell and Gooseff: Treatise in Fluvial Geomorpholgy – Geomorphic Controls on Hyporheic Exchange 621 specific locations within channels where the velocity head can provide additional potential and
622 thereby influence the pattern of hyporheic exchange.
623 624 9.13.3.2 Change in processes driving HEF through the stream network
625 626 Hyporheic exchange will vary widely across the sequence of channel types found in stream
627 networks (Figure 2; Buffington and Tonina, 2009). Channel networks generally follow a pattern
628 of steep headwaters to low-gradient reaches downstream. In mountain stream networks in
629 particular, gradient changes are expected to be accompanied by channel morphology changes
630 resulting in a sequence of distinct channel morphologies (Figure 2A). Obviously, bedrock
631 reaches have negligible hyporheic zones (Gooseff et al. 2005; Wondzell, 2006). We are unaware
632 of any studies of HEF in colluvial and cascade channel morphologies, however the extremely
633 high longitudinal gradients of these channels likely result in high velocity underflow which has
634 been shown to restrict the extent of the hyporheic zone (Storey et al. 2003). Also, the relatively
635 disorganized structure of the bed sediment prevents development of stepped water surface
636 profiles so that hydrostatically driven exchange due to longitudinal changes in gradient will
637 likely be low. Turbulent diffusion is likely to be a primary driver of HEF (Figure 2B).
638 639 Free-formed step-pool channels occur at slightly lower gradients (Figure 2A). These channels
640 have well-organized structure with periodic spacing of both steps and pools (Chin, 2002; Wohl
641 and Merrit, 2008) that have been shown to be primary drivers of HEF (Figure 2B; Kasahara and
642 Wondzell, 2003). The addition of large wood can substantially increase sediment storage
643 (Nakamura and Swanson, 1993; Montgomery et al. 1996), the development of step-pool
644 structure, and the extent, amount and residence times of HEF in these stream reaches (Wondzell,
645 2006). Other hydrostatic factors tend to have less dominance on HEF; these reaches have low
646 sinuosity so meander bends are uncommon and steep longitudinal gradients limit the potential
647 for back channels to create lateral HEF flow paths.
648 649 We are unaware of any published studies examining HEF in plane-bed channels. However, we
650 expect HEF to be lower than in either step-pool or pool-riffle channels (Figure 2B). The
651 streambed tends to be smoothly graded in these channels as suggested by their name, and there is
22
Wondzell and Gooseff: Treatise in Fluvial Geomorpholgy – Geomorphic Controls on Hyporheic Exchange 652 low spatial heterogeneity in surface texture (Buffington and Montgomery, 1999). Pools are
653 widely spaced, and both steps and riffles are rare. While these channels occur as free-formed
654 morphologies, pool-riffle channels can be converted to plane-bed channels by land use practices
655 that increase sediment supply and through the direct removal of large wood, with concurrent
656 decreases in HEF.
657 658 Lower in the stream network, channels tend have lower longitudinal gradients (Figure 2A), and
659 even in mountainous areas, unconstrained stream reaches become increasingly common. Channel
660 planforms can be quite complex in these rivers and as a consequence, a wide array of channel
661 geomorphic features influences HEF. Braided and anastomosing channels may form where
662 sediment loads are high and stream banks are erodible; the complex of channels likely leads to
663 substantial HEF through islands. Meandering channels form under lower sediment loads and
664 where banks are more stable. Meandering channels typically have pool-riffle morphologies,
665 although complexes of secondary channels, back channels, and paleochannels are common, a
666 legacy of past floods, channel avulsions and overbank deposition. Because most HEF occurs
667 along short, near stream flow paths, riffles are the dominant feature determining the amount of
668 HEF (Kasahara and Wondzell, 2003). However, the shape of the hyporheic flow net and the
669 residence time distribution of HEF will be strongly influenced by the complex of channel
670 planforms. Finally, hydrodynamic processes are expected to dominate in streams with relatively
671 mobile streambeds characterized by dune-ripple bedforms. These streams have low longitudinal
672 gradients and therefore channel morphologic features tend not to create steep hydrostatic head
673 gradients (Figure 2B).
674 675 Other exchange processes are likely to be related to specific conditions. Turn-over exchange will
676 only occur when bed material is mobile – a characteristic feature of both anastomosing and dune-
677 ripple channels. Transient exchange will only be appreciable during wet catchment conditions,
678 when channel stage is high and surrounding groundwater tables are comparatively low.
679 However, transient exchange may be a dominant form of HEF in regulated rivers where stage
680 fluctuates over daily cycles due to hydroelectric generation. Turbulent diffusion, on the other
681 hand, is likely to occur in gravel bed sections of the network, likely with the greatest potential
23
Wondzell and Gooseff: Treatise in Fluvial Geomorpholgy – Geomorphic Controls on Hyporheic Exchange 682 influence in either cascade or plane-bed sections of stream networks where stream velocities are
683 expected to be high.
684 685 9.13.4. Conclusion
686 687 Hyporheic exchange results from distinct processes, and the relations between those processes
688 and geomorphology are well understood from a mechanistic perspective. Thus, geomorphology
689 provides a critical framework to understand hyporheic processes and how they change with
690 location within a stream network, and over time in response to changes in stream discharge and
691 catchment wetness. To the degree that these geomorphic patterns are predictable, they provide
692 the foundation for hydrologists to make general predictions of the relative importance of the
693 hyporheic zone at the scale of entire catchments. Reach to reach variability is high in stream
694 networks, however, so understanding HEF at the reach scale continues to require detailed study
695 of specific stream reaches. These studies are difficult and current methodological approaches are
696 insufficient to fully examine the full suite of processes that account for patterns of HEF in any
697 specific stream reach. Consequently, hyporheic studies tend to focus on a single factor, or at
698 most a small subset of the factors driving HEF. Hyporheic researchers recognize that such
699 studies are incomplete. Detailed, holistic understanding of the importance of different processes
700 in driving HEF, how the relative importance of these processes changes with location in the
701 stream network, with the specific structure of any given stream reach, and with changes in
702 discharge and lateral groundwater inputs remains elusive.
24
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Wondzell and Gooseff: Treatise in Fluvial Geomorpholgy – Geomorphic Controls on Hyporheic Exchange 1008 Wondzell, S. M., LaNier, J., Haggerty, R., Woodsmith, R. D., and Edwards, R. T. 2009. Changes
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1019 35
Wondzell and Gooseff: Treatise in Fluvial Geomorpholgy – Geomorphic Controls on Hyporheic Exchange 1020 Figure Legends:
1021 1022 Figure 1. Idealized conceptual model of nested hyporheic flow paths as influenced by step-pool
1023 or pool-riffle sequences. A) Plan view showing arcuate HEF flow paths through the adjacent
1024 floodplain created by the change in the longitudinal gradient over the pool-riffle sequence where
1025 the amount of HEF is proportional to the head gradient. B) Longitudinal-section along the
1026 thalweg of the stream showing the vertical component of HEF flows through the streambed.
1027 1028 Figure 2. A) Hypothetical distribution of channel types along a stream profile in a mountainous
1029 stream catchment (redrawn from Montgomery and Buffington, 1997), and B) the corresponding
1030 relative contribution of turbulent diffusion and both hydrostatic or hydrodynamic processes to
1031 the total amount of HEF occurring within a stream reach. (Note that boundaries between channel
1032 types are often less distinct than shown here and that a range of conditions occurs within each
1033 category, thus the contribution of each process varies both among and within each channel type).
1034 1035 Figure 3. Idealized conceptual model of the influence of lateral inflows on hyporheic exchange
1036 flows. A) A high gradient stream where floodplain alluvium has relatively high saturated
1037 hydraulic conductivity under relatively dry conditions when lateral inputs are low and easily
1038 transported down valley via subsurface flow. Lateral inputs still reach the stream, but are
1039 diverted towards zones with hyporheic upwelling. B) A low gradient stream where floodplain
1040 alluvium has relatively low saturated hydraulic conductivity under relatively wet conditions
1041 when lateral inputs are sufficiently large to overwhelm down valley transport, causing lateral
1042 inputs to cross the valley toward the stream. Lateral inputs severely restrict hyporheic exchange
1043 flows. Legend follows Figure 1.
1044 1045 Figure 4. Idealized conceptual model of the influence changing stream stage on transient
1046 hyporheic exchange. A and B) A losing reach at low baseflow is converted to a gaining reach
1047 during a storm because precipitation recharge and lateral inputs of hillslope water increase water
1048 table elevations more than the corresponding increase in the stream stage. The original stream
1049 and water table position from 4A is shown in 4B for reference (light grey line). C & D) An
1050 example of a river where changes in stream stage result from snow melt, tidal influences, or dam
36
Wondzell and Gooseff: Treatise in Fluvial Geomorpholgy – Geomorphic Controls on Hyporheic Exchange 1051 releases far upstream. Increased stream stage causes stream water to flow into the adjacent
1052 aquifer creating a losing stream reach. Conversely, decreased stage leads to drainage of the
1053 aquifer creating a gaining reach. Alternating increases and decreases in stream stage leads to
1054 transient hyporheic exchange. The neutral condition (where stream stage is equal to the water
1055 table elevation) is shown for reference (black and white dashed line). Legend follows Figure 1.
1056 1057 Figure 5. Idealized conceptual model of the influence of the change in saturated cross-sectional
1058 area of the floodplain on hyporheic exchange flows. A) The influence of change in valley
1059 constraint with downwelling at the head of an unconstrained reach and upwelling at the
1060 downstream end of the reach caused by the transition from narrow bedrock gorges to wide
1061 alluvial valley floors. B) The influence of variations in depth to bedrock forcing upwelling
1062 upstream of a bedrock sill and downwelling downstream, where the depth of alluvium again
1063 increases. Legend follows Figure 1.
1064 1065 Figure 6. Idealized conceptual model of the influence of meander bends on hyporheic exchange
1066 flow. A) Simple, low radius meander with HEF traversing the point bar and floodplain. B) High
1067 radius meander with incipient meander-cutoff, where the short distance across the neck leads to
1068 much higher head gradients and thus greater HEF through the neck than the remainder of the
1069 meander bar. C) Meander bend with riffles located at the inflections between adjacent meanders
1070 so that head gradients through the point bar are low and much of the HEF occurs around the
1071 riffles, driven by longitudinal changes in gradient. Legend follows Figure 1.
1072 1073 Figure 7. Idealized conceptual model of the influence of back channels on hyporheic exchange
1074 flows. A) A back channel is incised below the water table, acts as a drain, and creats head
1075 gradients from the main channel to the back channel. B) A back channel is plugged near its
1076 downstream end, conducts water onto the floodplain, raises the water table and creats head
1077 gradienets from the back channel to the main channel. C) Complex pattern of HEF caused by
1078 interactions between a riffle in the main channel and a back channel. Paleochannels (dashed
1079 lines) support preferential flow. Legend follows Figure 1.
1080 37
Wondzell and Gooseff: Treatise in Fluvial Geomorpholgy – Geomorphic Controls on Hyporheic Exchange 1081 Figure 8. Idealized conceptual model of the influence of mid-stream islands on hyporheic
1082 exchange flows. A) Parallel and smooth longitudinal gradients in the channels on both sides of
1083 the island create HEF flow paths that parallel stream flow. B) Riffles at the head of the island
1084 enhance head gradients leading to greater HEF. C) Offset riffles create strong cross-island head
1085 gradients and flow paths, resulting in more HEF but with shorter flow path lengths and residence
1086 times. Legend follows Figure 1.
1087 1088 Figure 9. Idealized longitudinal-section in the center of a straight stream channel with bedforms
1089 (triangular dunes) showing the interaction with stream flow that creates regions of low- and high-
1090 pressure on the streambed which drive HEF. Non-hyporheic subsurface flows, known as
1091 underflow (dashed arrows), are present beneath the hyporheic zone. Legend follows Figure 1.
1092 1093 Figure 10. Examples of complex hyporheic flow paths resulting from interactions between
1094 channel morphologic features: A) a steep, 2nd-order step-pool channel with abundant large wood,
1095 and B) a moderate gradient, 5th-order pool-riffle channel with two major spring brooks. Note the
1096 difference in spatial scale between the two stream reaches. Letters indicate morphologic features
1097 driving HEF: S – steps; R – riffles; M – meander bends; B – back channels / spring brooks; I –
1098 islands; and T – a steep riffle at the mouth of a tributary. Equipotential intervals (dashed lines)
1099 are 0.2 m. Hyporheic flow paths (arrows) are hand drawn to indicate general direction of
1100 hyporheic flow through the valley floor.
38
A
Pool - riffle - pool sequence
B
Riffle
Pool
Pool
Floodplain
Riffle
Active
Channel
Subsurface
flow path
Wetted
Channel
A
Hillslope
Hollow
Colluvialbedrock
Cascade
Relative HEF
Steppool
Planebed
Pool-riffle
Dune-ripple
B
Hydrostatic
contribution
Turbulent
diffusion
Hydrodynamic
contribution
Longitudinal stream profile through stream network
A
Higher gradient, low lateral inputs
B
Lower gradient, high lateral inputs
B
Precipitation
Stage Increase
Lateral inputs
A
C
Stage Increase
D
Stage Decrease
A
Unconstrained stream reach
B
Bedrock sill
Bedrock
gorge
A
B
C
Pool / Run
A
B
C
A
B
C
Stream flow
A
S
I
S
S
S
S
S
B
S
S
S
Scale = 20 m
Log
B
Bedrock
Valley floor alluvium
Hillslope or terrace
Wetted stream channel
M
Back channel
Equipotential (0.2 m)
Hyporheic flow path
R
B
T
R
R
R
Scale = 50 m
R
I
R
B
M