Wondzell and Gooseff: Treatise in Fluvial …...Wondzell and Gooseff: Treatise in Fluvial...

Wondzell and Gooseff: Treatise in Fluvial Geomorpholgy – Geomorphic Controls on Hyporheic Exchange

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9.13 Geomorphic Controls on Hyporheic Exchange Across Scales - Watersheds to Particles 1

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Steven M. Wondzell 3

U.S. Forest Service, 4

Pacific Northwest Research Station, 5

Olympia Forest Sciences Laboratory, 6

Olympia, WA 98512 USA. 7

Phone: 360-753-7691 8

E-mail: [email protected] 9

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Michael N. Gooseff 11

Civil & Environmental Engineering Department, 12

Pennsylvania State University, 13

University Park, PA 16802 USA 14

Phone: 814- 867-0044 15

E-mail: [email protected] 16


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Abstract 17

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We examined the relationship between fluvial geomorphology and hyporheic exchange flows. 19

We use geomorphology as a framework to understand hyporheic process and how these 20

processes change with location within a stream network, and over time in response to changes in 21

stream discharge and catchment wetness. We focus primarily on hydostatic and hydrodynamic 22

processes – the processes where linkages to fluvial geomorphology are most direct. Hydrostatic 23

processes result from morphologic features that create elevational head gradients whereas 24

hydrodynamic processes result from the interaction between stream flow and channel 25

morphologic features. We provide examples of the specific morphologic features that drive or 26

enable hyporheic exchange and we examine how these processes interact in real stream networks 27

to create complex subsurface flow nets through the hyporheic zone. 28

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Key words 31

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Hyporheic, step-pool sequence, pool-riffle sequence, meander bends, back channels, floodplain 33

spring brooks, mid-channel islands, stream bedforms, pumping exchange, saturated hydraulic 34

conductivity. 35


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9.13.1. Introduction 36

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Hyporheic exchange flow (HEF) is the movement of stream water from the surface channel into 38

the subsurface and back to the stream (Figure 1). Stream water in hyporheic flow paths may mix 39

with groundwater so that the relative proportion of stream-source water in the hyporheic zone is 40

highly variable, ranging from 100% stream water to nearly 100% groundwater. Also the 41

residence time distribution of stream water in the hyporheic zone tends to be highly skewed, with 42

most of the stream water moving along short flow paths and thus having short residence times 43

(hours), but some water either moving on long flow paths or encountering relatively immobile 44

regions having very extended residence times (weeks to months, or longer). The boundaries of 45

the hyporheic zone are arbitrary, usually defined by the amount of stream-source water present in 46

the subsurface. Triska et al. (1989) set a threshold of 10% stream-source water to define the 47

limits of the hyporheic zone so that regions with <10% stream-source water were defined as 48

groundwater. Alternatively, the extent of the hyporheic zone can be delimited by water residence 49

time, for example, the subsurface zone delineated by hyporheic exchange flows with residence 50

times less than 24 hours (the 24-h hyporheic zone; Gooseff, in press). 51

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The objective of this chapter is to examine the relation between geomorphology and hyporheic 53

processes. The two primary controls on hyporheic exchange are the gradients in total head 54

established along and across streambeds and the hydraulic conductivity of the streambed and 55

adjacent aquifer, both of which are significantly influenced by geomorphology. Total head (also 56

known as potential) is the sum of pressure head, elevation head, and velocity head. Pressure head 57

represents height of a column of fluid to produce pressure. Velocity head represents the vertical 58

distance needed for the fluid to fall freely (neglecting friction) to reach a particular velocity from 59

rest. Elevation head represents the potential energy of a fluid particle in terms of its height from 60

reference datum. Hydrostatic head is referred to as the sum of elevation and pressure head. 61

Groundwater tables in unconfined aquifers represent the spatial gradients in hydrostatic head. A 62

number of processes either drive or enable HEF, several of which are based on changes in head 63

gradients. We follow the organizational structure presented by Käser et al. (2009), who divided 64

these processes into five distinct classes: 65

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1. Transient exchange – the temporary movement of stream water into stream banks due to 67

short-term increases in stream stage (i.e., bank storage processes due to changes in 68

hydrostatic head gradients between stream and lateral riparian aquifer; Lewandowski et al. 69

2009; Sawyer et al. 2009a). 70

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2. Turn-over exchange – the trapping of stream water in the streambed during times of 72

significant bed mobility (Elliot and Brooks, 1997b; Packman and Brooks 2001). 73

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3. Turbulent diffusion – exchange driven by slip velocity that is created at the surface of the 75

porous medium of the bed where streamwise velocity vectors continue to propagate into the 76

surface layers of the bed (Packman and Bencala, 2000). 77

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4. Hydrostatic-driven exchange – exchange driven by static hydraulic gradients which are 79

determined by changes in water surface elevation (Harvey and Bencala, 1993), spatial 80

heterogeneity in saturated hydraulic conductivity, or changes in the saturated cross-sectional 81

area of floodplain alluvium through which hyporheic flow occurs. 82

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5. Hydrodynamic-driven exchange – exchange driven by the velocity head component of the 84

total head gradient on the bed surface (i.e., pumping exchange; Elliott and Brooks, 1997a,b) 85

and exchange induced by momentum gradients across beds and banks. 86

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These classes of HEF processes are coupled to geomorphic processes in many ways. This is most 88

obvious for hydrostatic effects, which are directly dependent on channel and valley-floor 89

morphology and the depositional environment that controls spatial heterogeneity in saturated 90

hydraulic conductivity (K). However, turnover of streambed sediment is also related to fluvial 91

geomorphic processes. Similarly, hydrodynamic effects result from the interaction of flow over 92

stream bedforms. Geomorphic processes build stream bedforms and determine channel 93

morphology, especially longitudinal gradient, bed roughness, and water depth all of which 94

influence flow velocity. The relationship between geomorphology and the other classes of 95

processes is less direct, but still plays a role in controlling these processes through channel form 96

and the size distribution of sediment that makes up the streambed. This chapter focuses primarily 97


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on the hydostatic and hydrodynamic processes where linkages to geomorphic processes are most 98

direct. 99

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We organize our discussion of the interactions between geomorphology and HEF using a 101

hierarchical scaling framework developed for river networks (Frissell et al. 1986; Bisson and 102

Montgomery, 1996), starting at the whole network, through the stream segment, to the stream 103

reach, to the channel unit, and down to the sub-channel unit scale. We recognize that describing 104

any given process or related flow path at a single “scale” is somewhat arbitrary because of the 105

nested structure of the hyporheic flow net and dispersion among HEF flow paths. Despite that, 106

the concept of scale is an important heuristic tool to organize our understanding of hyporheic 107

processes. In many senses, the reach scale is the most informative scale at which to consider 108

HEF. A single reach, by definition, has characteristic channel morphology so that the factors 109

driving HEF within the reach are relatively consistent. However, only a few of the geomorphic 110

factors driving HEF actually operate at this scale. Most of the drivers work at the channel unit or 111

smaller scales. And to understand the importance of HEF in stream ecosystem processes, the 112

cumulative effects of HEF must be evaluated at scales much larger than a single reach. 113

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9.13.2. The effect of geomorphology on hyporheic exchange flows 115

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9.13.2.1. The whole network to segment scale 117

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The geologic setting of the stream network is an important factor determining the likely 119

occurrence of HEF, but there have been few attempts to study HEF at this broad scale. Rather, 120

our expectations are pieced together by drawing comparisons among HZ studies that have been 121

conducted in widely varying geologic settings, at different locations in the stream network, or 122

under widely varying flow conditions. We expect that geomorphic-hyporheic relationships will 123

differ substantially among different geologic settings. 124

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Fluvial geomorphic studies have examined the factors that determine the types of channel 126

morphologies present within stream networks (Montgomery and Buffington, 1997; Wohl and 127

Merritt, 2005; Brardinoni and Hassan, 2007). Montgomery and Buffington (1997) presented one 128


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such description of the distribution of channel morphologies typical of many mountainous 129

landscapes. They showed that catchment area and channel longitudinal gradient controlled the 130

development of distinct channel types such that the channel types tended to follow a 131

characteristic sequence within a catchment (Figure 2A). In their example, this sequence starts 132

with bedrock and colluvial channels in the steepest, upper-most headwaters. As longitudinal 133

gradients decrease, channels change to cascades, to step-pool, to plane-bed, to pool-riffle, and the 134

largest, lowest gradient rivers were typified as dune-ripple channels. Along with these changes in 135

channel morphology, the following would be expected: decreased longitudinal gradient and 136

mean grain size of streambed sediment, and increased depth, width, hydraulic radius, and flow 137

velocity (Leopold and Maddock, 1953; Wohl and Merritt, 2008). 138

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In this paper, we use Montgomery and Buffington’s (1997) description of the sequence of 140

channel types within a catchment as a simple heuristic model to organize our examination of the 141

relative importance of the different processes that drive HEF within stream networks. We 142

recognize that local controlling factors often interrupt simple sequencing of channel types. For 143

example, landslides may block large mainstem channels, creating locally steep gradients over the 144

landslide debris and uncharacteristically low gradients in the depositional reach immediately 145

upstream (Benda et al. 2003). We also recognize that regional differences in geology and 146

geomorphology will lead to dramatically different spatial organization of channel types (see for 147

example characteristic channel type in glaciated mountainous regions as described by Brardinoni 148

and Hassan, 1997). Our descriptions of the spatial organization of stream types and the resulting 149

HEF processes will have to be modified for any specific landscape. 150

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Most hyporheic exchange results from head gradients pushing water through the streambed. The 152

amount of stream water entering the hyporheic zone is thus a function of the steepness of the 153

head gradient and the saturated hydraulic conductivity of the streambed and underlying aquifer. 154

The head gradients can be induced in many ways, but the two of primary influence are the 155

hydrostatic and hydrodynamic processes. The relative importance of each of these processes is 156

expected to vary among channel types and with longitudinal gradient. In high gradient streams, 157

channel forms such as step-pool sequences or pool-riffle sequences can create very steep 158

hydrostatic head gradients. Further, because of high bed roughness and relatively shallow water 159


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depth, flow velocities tend to be lower in small steep streams than in larger, low gradient streams 160

(Leopold and Maddock, 1953; Wondzell et al. 2007). In contrast, it is difficult for natural 161

processes to create steep changes in the longitudinal gradient in low gradient streams. Instead, 162

stream flow interacts with stream bedforms, such as dunes or ripples, such that hydrodynamic 163

forces dominate the development of head gradients through the streambed. Thus, we expect that 164

hydrostatic effects will dominate in high gradient channels and that hydrodynamic processes will 165

dominate in low gradient channels (Figure 2B). Further, because channel types and longitudinal 166

gradients generally vary systematically within stream networks, we further expect that 167

hydrostatic effects will tend to dominate in the upper portions of stream networks and that the 168

relative importance of hydrodynamic processes will increase down the stream network. 169

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9.13.2.2. The reach scale – setting the potential for hyporheic exchange 171

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The potential for HEF to occur varies within any given stream reach. Roughly speaking, this 173

potential is determined by the factors that generate head differences that drive HEF, the 174

properties of the subsurface alluvium through which HEF occurs, and the potential effect of 175

lateral groundwater inputs from adjacent hillslopes that might limit hyporheic expression. 176

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9.13.2.2.1. Losing and gaining reaches 178

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Hyporheic exchange is likely to be more limited in strongly gaining reaches than in neutral 180

reaches because of steep streamward hydrologic gradients surrounding the channel (Wroblicky et 181

al. 1998; Storey et al. 2003; Malcolm et al. 2003 and 2005; Cardenas, 2009). Similarly, where 182

water is lost to regional aquifers in strongly losing reaches, return flows of stream water back to 183

the stream are likely to be severely restricted and thus also limit the expression of the hyporheic 184

zone (Cardenas, 2009). These patterns of gains and/or losses are controlled, at some level, by 185

regional groundwater and catchment characteristics interacting with smaller scale effects. In 186

large gaining rivers, Larkin and Sharp (1992) demonstrated that the relative dominance of cross-187

valley vs. down valley flow paths through valley-floor aquifers varied depending on the 188

longitudinal gradient of the valley floor and the hydraulic conductivity of the valley floor 189

alluvium. In higher gradient reaches (>0.004 m/m) and in areas with coarser substrate, flow was 190


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predominantly down valley. Conversely, where valley floor gradients were shallower or 191

sediment more finely textured, flow tended to be toward the stream. Thus, the way in which 192

lateral inputs influence hyporheic exchange is not solely a function of their magnitude, but also a 193

function of the ability of subsurface water to move down-valley (Storey et al. 2003). The ratio 194

between these two factors – the magnitude of the inputs relative to down valley flow – 195

determines how hyporheic exchange is affected. 196

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As a first approximation, the potential for down valley flow can be estimated using the 198

relationships summarized in Darcy’s Law – that is, the product of the longitudinal valley 199

gradient, the saturated cross-sectional area of the floodplain perpendicular to the direction of 200

subsurface flow, and the hydraulic conductivity of the alluvium. As lateral inputs increase, 201

several factors may change: (1) water tables may rise, thus increasing the saturated thickness and 202

the cross-sectional area through which water flows allowing the transmission of more water, or 203

(2) flow paths may begin to turn obliquely toward the stream, which also increases the saturated 204

cross-sectional area and may also increase head gradients. Consequently, under dry conditions 205

when lateral inputs are relatively small, the potential extent and magnitude of hyporheic 206

exchange can be fully expressed (Figure 3A). As subsurface flows turn toward the channel they 207

begin to limit the extent of the hyporheic zone with only minor effect on the HEF (Wondzell and 208

Swanson, 1996). If sufficiently large, lateral inputs can severely limit both the spatial extent and 209

magnitude of hyporheic exchange (Figure 3B; Harvey and Bencala, 1993; Wroblicky et al. 1998; 210

Storey et al. 2003; Cardenas and Wilson, 2007; Malcolm et al. 2003, Soulsby et al. 2009). 211

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Simple generalizations of where and when lateral inputs will limit HEF are difficult because of 213

the wide range of geomorphic settings in which HEF occurs and because the magnitude of lateral 214

inputs changes with catchment wetness. Lateral inputs are expected to be high when catchments 215

are wet and decrease as catchments dry out. However, lateral inputs are not spatially uniform. In 216

steep mountainous settings, the size of the upslope area draining directly to the valley floor is 217

important, concentrating lateral inputs in zones at the base of hillslope hollows (Jencso et al. 218

2009). Lateral inputs may persist the entire year at the bases of the largest hillslope hollows. 219

Most hillslope hollows are small, however, so that most of the stream network would be 220

disconnected from lateral inputs except for short periods of time when catchments are very wet, 221


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for example after large storms or during peak snowmelt. We are unaware of similar studies 222

relating topography to spatial patterns of hillslope inputs in areas of low relief with humid 223

climates. However, Storey et al. (2003) reported that an extensive shallow surfical aquifer was 224

present along their lowland, low-gradient study reach and that lateral inputs of groundwater 225

substantially reduced both the extent and the amount of hyporheic exchange flows except during 226

summer baseflow. Clearly, the influence of lateral inputs may be much different in lowland 227

catchments than in steep mountainous catchments. 228

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Changes in lateral inputs to streams do not occur in isolation. Rather, they are likely to be 230

accompanied by corresponding changes in stream stage (and discharge). The change in water 231

table elevations resulting from changed lateral inputs must be considered relative to the 232

accompanying changes in stream stage. Although the number of studies examining changes in 233

hyporheic flow paths with changing catchment wetness is limited, studies in small mountain 234

streams suggest that water table elevations in the floodplain increase more than stream stage so 235

that HEF is typically more restricted when catchments are wet (Figure 4A and 4B; Harvey and 236

Bencala, 1993; Wondzell and Swanson, 1996; Stednick and Fernald, 1999). Storey et al. (2003) 237

reported similar results for a lowland, low-gradient river. 238

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In some cases, however, stream stage may change markedly without corresponding changes in 240

precipitation recharge or changes in lateral inputs. Most examples of these processes come from 241

large, lowland rivers because river stage is controlled by processes far upstream. These “bank 242

storage” processes (Pinder and Sauer, 1971) have been recognized as a form of transient 243

hyporheic exchange (Figure 4C and 4D) that can result from both in-bank or over-bank floods 244

(Bates et al. 2000; Burt et al. 2002). In some situations, increased stream stage may even lead to 245

groundwater ridging in the floodplain, reversing head gradients and limiting lateral groundwater 246

inputs. Similarly, hyporheic exchange through stream banks can result from diel variations in 247

stream stage (and discharge) during snow melt periods (Loheide and Lundquist, 2009) or from 248

tidally induced changes in water elevations in coastal streams and rivers (Bianchin et al. in 249

press). 250

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Transient hyporheic exchange may be especially evident in regulated rivers where releases from 252

dams (or other control structures) can result in large and rapid changes in river stage without 253

corresponding local precipitation to recharge floodplain aquifers (e.g., Fritz and Arntzen, 2007; 254

Lewandowski et al. 2009; Sawyer et al. 2009a; Francis et al., in press). However, transient 255

hyporheic exchange may not always result from fluctuations in river stage. For example, 256

Hanrahan (2008) studied vertical HEF through the streambed of a large, regulated gravel bed 257

river where stage sometimes changed by nearly 2 m in an hour. For the most part, they did not 258

observe transient hyporheic exchange related to changes in stage. They concluded that 259

hydrostatic and hydrodynamic processes remained the dominant control on HEF. Notably, 260

Hanrahan (2008) did not examine lateral exchanges through the stream banks, which can be 261

more responsive to changes in stage than are locations in the stream channel itself (Storey et al. 262

2003). Water table fluctuations in the floodplain at long distances from the stream are not 263

necessarily indicative of extensive HEF because pressure fluctuations can propagate through 264

surficial (unconfined) aquifers much faster than does the actual flow of stream water. This was 265

clearly demonstrated by Lewandowski et al. (2009) who showed that river water penetrated, at 266

most, only 4 m into the stream bank even though water table fluctuations were observed more 267

than 300 m from the river. 268

269

HEF can occur in strongly gaining and losing reaches because of the nested structure of 270

hyporheic flow paths, and because HEF can occur at a variety of spatial scales. Thus an envelope 271

of the HZ can be set within larger non-hyporheic flow paths (Figure 3B; Cardenas and Wilson, 272

2007). Similarly, smaller-scale HEF can occur as a result of smaller scale geomorphic drivers, 273

even within a reach that is, overall, strongly losing (Payn et al. 2009). Further, because HEF is 274

dominated by relatively near-stream flow paths that are short in length and residence time 275

(Kasahara and Wondzell, 2003), the magnitude of HEF can be substantial, even in strongly 276

gaining reaches where the spatial extent of the hyporheic zone is greatly restricted (Wondzell and 277

Swanson, 1996; Cardenas and Wilson, 2007; Payn et al. 2009). 278

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9.13.2.2.2. Changes in saturated cross-sectional area 280

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The saturated cross sectional area of the floodplain (orthogonal to groundwater flow path 282

direction) is one of the factors determining the amount of groundwater transmitted down valley 283

through the valley floor alluvium. Thus, any change in the cross sectional area along the length 284

of a stream reach will lead to parallel changes in the down valley flow of water through the 285

floodplain, thereby driving downwelling from, or upwelling to the stream (Stanford and Ward, 286

1993). Downwelling occurs where valley floors increase in width, for example, downstream of 287

bedrock-constrained reaches (Figure 5A; Poole et al. 2004 and 2006; Acuna and Tockner, 2009). 288

Conversely, upwelling occurs where valley floors narrow at the lower end of wide unconstrained 289

reaches (Figure 5A; Baxter and Hauer, 2000; Acuna and Tockner, 2009). Similarly, variations in 290

the thickness of the surficial aquifer, caused by variations in depth to bedrock or other confining 291

layers drive similar patterns of upwelling and downwelling. For example, upwelling commonly 292

occurs just upstream of bedrock sills with a subsequent transition to downwelling just 293

downstream of such bedrock sills as the surficial aquifer again thickens (Figure 5B; Valett, 294

1993). This is easily observed in streams in arid regions during the dry season, where perennial 295

flow may only occur above bedrock sills, which force the subsurface flow to the surface. 296

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9.13.2.3. The sub-reach to channel-unit scale – hydrostatic processes 298

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Geomorphic features of the stream channel and valley floor within stream reaches control the 300

elevation of surface water and can thereby create significant head gradients through the valley 301

floor alluvium, driving HEF. Because these geomorphic features are static on the time scales 302

typical of hyporheic exchange (hours to weeks) they are broadly recognized as “hydrostatic 303

processes”. 304

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9.13.2.3.1. Step-pool and pool-riffle sequences 306

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One of the best-studied examples of hydrostatic processes involves the changes in water surface 308

elevation along a pool-step sequence and the resulting head gradients that drive HEF (Figure 1; 309

Harvey and Bencala, 1993). Harvey and Bencala (1993) showed that the change in the 310

longitudinal gradient of the stream channel (which approximates the stream energy profile) drove 311

HEF. They also observed that HEF flow paths tended to be curved – first curving away from the 312


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stream above the step or riffle and then curving back to the stream below the step or riffle. 313

Building from their observations, model analyses show that along an idealized straight channel 314

with homogeneous isotropic porous sediment, hyporheic flow paths around a change in the 315

longitudinal gradient will exploit the full 3-dimensional saturated volume along the channel, thus 316

extending both vertically beneath the streambed and horizontally through the streambanks and 317

near stream aquifer (Figure 1A and 1B). Real streams are substantially more complicated, 318

however, such that changes in hydraulic conductivity of the alluvium, bends in the channel, and 319

the spatial location of lateral groundwater inputs lead to the development of a complicated flow 320

net through the valley floor (e.g., Cardenas and Zlotnik, 2003). Despite these complexities, the 321

steepness of the hydraulic head gradient imposed by the change in the longitudinal gradient and 322

the saturated hydraulic conductivity control the amount of stream water exchanged with the 323

subsurface. 324

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Many factors can modify the effect of steps or riffles on HEF. For example, the height of the step 326

(or steepness of the riffle) determines the head gradient available to drive HEF so that a single 327

very large step has the potential to drive more HEF than if the same amount of elevational 328

change is spread over several smaller steps (Kasahara, 2000). Because of this, large wood can be 329

important in determining the amount of HEF in forest streams. Single logs tend to create 330

frequent, small obstructions that collect and store small amounts of sediment, forming pool-step 331

sequences in which the extent of the hyporheic zone tends to be small (Wondzell, 2006). 332

Although log jams are less common, they can create large obstructions storing sediment in 333

wedges several meters deep and 10 or more meters in length, and significantly widen constrained 334

stream channels. Consequently, log jams can form extensive hyporheic zones in steep, confined 335

mountain streams (Wondzell, 2006). 336

337

Large, channel-spanning logs can wedge into steep narrow channels, forcing the accumulation of 338

sediment in channels, converting bedrock reaches to alluvial reaches with a step-pool 339

morphology (Montgomery et al. 1996), thereby greatly enhancing HEF. Similarly, large wood 340

can force plane-bed channels into a pool-riffle morphology (Montgomery et al. 1996) which 341

should lead to more HEF than would be present in a comparable wood-free channel. Large wood 342

can have the opposite effect in channels that would have a free-formed pool-riffle morphology. 343


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In one documented case, accumulations of large wood tended to force a pool-riffle channel 344

toward a step-pool morphology (Wondzell et al. 2009). The channel adjusted to removal of all 345

large, in-stream wood by developing a better defined pool-riffle structure around meander bends, 346

leading to increased sediment storage. Continued channel adjustment over time following the 347

removal of large wood eventually led to substantial increases in HEF. 348

349

The size, spacing, and sequence of channel units (e.g., pools and riffles) along the stream 350

longitudinal profile can also affect HEF (Anderson et al. 2005; Gooseff et al. 2006). Anderson et 351

al. (2005) made detailed measurements of channel profiles and patterns of HEF, and showed that 352

channel unit size and spacing increased as did the length of channel characterized by 353

downwelling with increasing drainage area in a mountainous stream catchment. Gooseff et al. 354

(2006) built on these results, examining HEF using 2-D groundwater models of idealized 355

longitudinal profiles of mountain streams. Gooseff et al.’s (2006) modeling results confirmed 356

that both channel unit spacing and size were important in determining hyporheic exchange 357

patterns of upwelling and downwelling. Perhaps more surprising, however, was the observation 358

that the sequence of channel units also affected simulated HEF. Gooseff et al. (2006) compared 359

pairs of idealized stream reaches that varied only by the way the longitudinal gradient changed 360

over the pool-riffle sequence – i.e., the slope of the riffle was gradual on its upstream end and 361

steepest at its downstream end (described as a pool–riffle–step sequence) versus riffles that were 362

initially steep with the slope decreasing toward the downstream end (described as a pool–step–363

riffle sequences). Simulated downwelling lengths were substantially longer for pool–riffle–step 364

sequences than for pool–step–riffle sequences. 365

366

9.13.2.3.2. Meander bends and point bars 367

368

A variety of channel and valley floor morphologic features, in addition to changes in the 369

longitudinal gradient, create head gradients with the potential to drive HEF. These include 370

channel meander bends and associated point bars, back channels or floodplain spring brooks, and 371

islands set between main and secondary channels. In all these cases, differences in the 372

elevational head of surface water between two channels, between different points in a single 373

channel around a meander bend, or between points on opposite sides of an island create head 374


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gradients that drive HEF. For example, head gradients through the point bar in a meander bend 375

are steeper than the longitudinal gradient of the stream channel around the point bar (Peterson 376

and Sickbert, 2006) so that stream water infiltrates the upper end of the point bar and is returned 377

to the channel at the lower end of the point bar (Figure 6A; Vervier and Naiman, 1992). More 378

generally, these exchange flows occur across the full length of meander bends and are influenced 379

by both the change in stream water elevation around the meander bend and the plan-view shape 380

of the meander bend. Highly evolved meander bends support steep head gradients across the 381

mender neck because of the close proximity of the stream channels (Figure 6B; Boano et al. 382

2006; Revelli et al. 2008) so that HEF is dominantly located in the meander neck, with much 383

reduced HEF across the remainder of the meander where head gradients are much lower. In other 384

cases, meanders develop a characteristic pattern of alternating pools and riffles, with riffles 385

located at the thalweg cross-overs in the inflections between adjacent meanders and pools or low 386

gradient runs wrapping around the point bar (Figure 6C). This combination of channel 387

morphologic features can create complex HEF flow paths within meander bends. The residence 388

times of HEF traversing meander bends can be quite short where meanders are small and 389

saturated hydraulic conductivities are high (Pinay et al. 2009). Conversely, residence times of 390

HEF may be extremely long in meander bends of low gradient rivers with fine textured sediment 391

(Boano et al. 2006; Peterson and Sickbert, 2006). 392

393

9.13.2.3.3. Back channels and floodplain spring brooks 394

395

Channel planforms are often complex in wide floodplains, including a network of old or 396

abandoned channels. If the upstream ends of these channels are plugged with sediment and if the 397

downstream ends are sufficiently incised to intercept the water table and are connected back to 398

the river at their downstream ends, they will act as drains, imposing head gradients from the 399

stream to the old channel (Figure 7A; Wondzell and Swanson, 1996; Poole et al. 2006). These 400

channels are also known as floodplain spring brooks because water upwells into the channel, 401

forming a spring at its head. In addition to creating HEF, these channels will capture whatever 402

water is in the surficial aquifer of the floodplain, including down valley flows from upstream 403

locations, and lateral inputs of groundwater or hillslope water from the valley margin. However, 404

because lateral inputs tend to be small and spatially isolated (Jencso et al. 2009; and as discussed 405


15

above), floodplain spring brooks will most often be fed by HEF (Wondzell and Swanson, 1996; 406

Jones et al. 2007). 407

408

Abandoned channels can also be plugged at their downstream ends and open to the river at their 409

upstream ends. In this case, stream water can flow into the abandoned channel, infiltrate the 410

channel bed and raise the water table in the middle of the floodplain, thereby creating head 411

gradients and driving HEF from the abandoned channel back to the main stream channel (Figure 412

7B). More complex situations arise when the longitudinal gradients in either the back channel or 413

mainstem channel are interrupted by steeper riffles or steps. Figure 7C shows the interactions 414

between a back channel and riffle. Above the riffle, water in the main channel is higher than the 415

back channel so water flows towards the spring brook. Downstream of the riffle, the main 416

channel is lower than the back channel so that the back channel loses water over its downstream 417

extent, eventually going dry before reaching the main channel. 418

419

The channel planform features that drive HEF can occur over a range of spatial scales, and their 420

influence may change through time as the stage height of water in the main channel changes. For 421

example, a small gravel bar may have low points along the stream bank. At high stage, the entire 422

gravel bar may be submerged. As stage decreases the center of the bar may become exposed, 423

creating a secondary channel along the bank. As stage decreases further, flow may become 424

discontinuous through the secondary channel such that it functions as a drain if it is plugged at 425

the upstream end, or functions as a conduit allowing stream water to infiltrate the surface of the 426

gravel bar if it is plugged at its downstream end. Old channels in large floodplains may act 427

similarly, with continuous flow along their full length during floods, but becoming disconnected 428

at intermediate to low stage, or even dry completely during periods of minimum discharge. In 429

large floodplain reaches, these channels can be 100’s of meters to kilometers in length, extending 430

nearly the full length of the stream reach (Poole et al. 2006; Arrigoni et al. 2008). 431

432

9.13.2.3.4. Secondary channels and islands 433

434

Islands present a special case of back channels in which the channel is continuously connected to 435

the main channel over its full length. Hyporheic hydrology of islands has not been extensively 436


16

studied. However, we expect that the surface water elevations in channels bounding the island 437

create boundary conditions for total head and control HEF through islands as is generally 438

indicated by the available literature (Dent et al. 2007; Francis et al., in press). If channels along 439

both sides of the island are parallel and symmetric with constant longitudinal gradient, then flow 440

through the island will parallel the channels and the head gradient driving flow will equal the 441

overall longitudinal gradient of the stream reach (Figure 8A). If riffles are present in the 442

channels, the head gradient through the island adjacent to the riffles can be much steeper than the 443

reach averaged longitudinal gradient (Figure 8B). Also, if riffles are displaced along the primary 444

and secondary channels surrounding an elongated island such that a riffle is located near the head 445

of the island in one channel and near the tail of the island in the second channel, the resulting 446

head gradients would tend to drive flows laterally through the island, leading to very large cross-447

sectional areas experiencing HEF, and therefore large amounts of HEF, albeit, with shorter 448

length flow paths (Figure 8C). While islands may be uncommon in most channel types, they may 449

dominate HEF in braided and anastomosing stream reaches (Ward et al. 1999; Arscott et al. 450

2001). Given the complexities of potential sizes and shapes of islands and patterns in 451

longitudinal gradients in the bounding channels, the resulting flow nets, residence times, and 452

amounts of HEF are likely to vary widely. 453

454

9.13.2.3.5. Spatial heterogeneity in saturated hydraulic conductivity 455

456

Fluvial processes control the depositional environment on the streambed and across the 457

floodplain creating spatial heterogeneity in the texture of deposited and reworked sediment 458

across a range of scales, from the surface of the streambed to the entire floodplain. Because 459

sediment texture is closely related to saturated hydraulic conductivity (K), these processes can 460

substantially influence HEF. However, because of the difficulties in quantifying these patterns at 461

the scales at which they influence HEF, they have been relatively little studied. At fine scales, 462

streambed roughness can control the depositional environment across the streambed (Buffington 463

and Montgomery, 1999), which lead to spatial patterns in the distribution of K within the 464

streambed (Genereaux et al. 2008), which in turn can influence both the location and amount of 465

HEF. HEF will be restricted where the streambed is clogged with fine sediment and 466

preferentially located in zones with higher K. Experiments in flumes have also shown that HEF 467


17

can also influence patterns of fine-sediment deposition, with fine sediment preferentially 468

deposited in downwelling zones (Packman and MacKay, 2003; Rehg et al. 2005) which may 469

explain differences in K between upwelling and downwelling zones observed in a steep 470

headwater stream (Scordo and Moore, 2009). 471

472

Spatially heterogeneous patterns in K influence HEF. For example, groundwater flow modeling 473

studies using homogeneous vs. heterogeneous K showed that spatial heterogeneity may add 474

substantial complexity to the spatial patterns of the hyporheic flow net (Woessner 2000). When 475

relatively high K regions are aligned parallel with head gradients they create preferential flow 476

pathways (Wagner and Bretschko, 2002) that can increase the total amount of HEF (Cardenas 477

and Zlotnik, 2003; Cardenas et al. 2004). Results from Cardenas et al. (2004) showed that 478

influence of heterogeneity in K was relatively greater in lower gradient streams and where head 479

gradients driving HEF were reduced. To our knowledge, the influence of fine-grained 480

heterogeneity has not been studied in steeper channels where hydrostatic processes dominate. 481

482

Fluvial processes also influence spatial patterns in K at the scale of the entire floodplain. 483

Especially important is the layering of stream and floodplain alluvium. Layering can create 484

strong vertical anisotropy (Chen, 2004), limiting vertical exchange and promoting lateral flows 485

through the streambed and floodplain (Packman et al. 2006; Marion et al. 2008). Overbank 486

deposition can also bury back channels creating “paleochannels” where coarse streambed 487

alluvium is buried under finer floodplain soils (Stanford and Ward, 1993; Stanford et al. 1994; 488

Poole et al. 2004). If these paleochannels intercept the water table, they will function as large 489

preferential-flow pathways that can route water the full length of a floodplain. In this regard they 490

function much like a subsurface version of back channels or floodplain spring brooks – either 491

acting as drains lowering the water table in the floodplain and imposing head gradients from the 492

stream to the paleochannel, or acting as distributaries, routing water into the floodplain and 493

imposing head gradients from the paleochannel to the stream. Locations of paleochannels are 494

sometimes evident from shallow depressions along the floodplain. In other cases, over-bank 495

deposition will have completely filled old channels so that there is no surficial indication on the 496

flat floodplain surface. The influence of paleochannels is difficult to discern because networks of 497


18

widely spaced wells are unlikely to find and trace the location of these features along the length 498

of the floodplain. As a consequence, their influence on HEF has not been widely studied. 499

500

9.13.2.4. The bedform scale – hydrodynamic processes 501

502

Channel hydraulics, and the spatial and temporal distribution of velocity (kinetic energy) across 503

streambeds are significantly influenced by the form of the channel and the bedforms that occur in 504

channels. The continuous feedback between pressure distribution and shear stress across the bed 505

surface and the potential to erode the bed will cause turn-over exchange to occur during times of 506

high flows. During lower flows, when bed sediment is relatively stable, bedforms cause some 507

level of form drag on the flows, inducing pressure distributions across the bedforms, thereby 508

driving HEF at a scale smaller than the bedform (Figure 9). The size of the bedform is set by 509

both the energy regime of the reach and the material that makes up the reach, and the form drag 510

induced on the water column by the bedform is of course partly controlled by its size. Thus, the 511

scale of HEF flowpaths induced by hydrodynamic exchange across the bedforms will scale in 512

part with the size of bedforms present (Cardenas et al. 2004). Finally, the heterogeneity of the 513

bed material that makes up the bedforms will have a distinct control on the flux rate and actual 514

flowpaths through and around the bedforms (Sawyer et al. 2009b). 515

516

In sand bed streams, hydrodynamic HEF has been extensively studied both theoretically and 517

empirically. Typical bed forms in sand bed streams are dunes and ripples, which have a fairly 518

predictable geometry and spacing, based on bed sediment composition and flow rate. 519

Thibodeaux and Boyle (1987) pioneered investigations of the hydrodynamic pressure 520

distribution across dunes, noting the penetration of channel water into the porous bed forms. 521

Further development of a ‘pumping exchange’ model by Elliot and Brooks (1997a,b) expanded 522

the ability to predict HEF and associated solute dynamics in channel-bed systems. Whereas most 523

studies of hydrodynamic exchange processes were generally carried out in or applied to flume 524

studies, there has been at least one application of incorporating the pumping exchange model to 525

tracer transport in field studies. Salehin et al. (2003) studied the transport of tracer along several 526

km of Sava Brook in Sweden and successfully applied a solute transport model to the observed 527

data to explain long time residence time distributions using the pumping exchange model theory. 528


19

The predictability of dune and ripple sizing and spacing makes the pumping exchange model a 529

useful tool to explore HEF in sand bed streams and rivers. 530

531

In gravel bed streams, bed form types may be generally predictable (i.e., Montgomery and 532

Buffington, 1997; Wohl and Merritt, 2008; Chin, 2002), but the exact geometry and spacing of 533

bed forms is less predictable, particularly at a scale that will directly influence head distributions 534

across and along the channel. Hence, the velocity distribution in the channel and around the bed 535

form, which contributes to hydrodynamic exchange, is also unpredictable. Tonina and 536

Buffington (2007) conducted careful studies of total pressure distribution across streambeds in 537

flumes that had ‘realistic’ geometry of a pool-riffle sequence in a gravel bed channel. Their 538

results indicated that total head distribution (i.e., incorporating velocity head in addition to 539

hydrostatic head) was important to exchange at focused points in the channel where high velocity 540

occurred. Further, they confirmed that in general, there was little or no contribution of velocity 541

head to parts of the bed that were overlain by deeper, slower flow, and therefore a hydrostatic 542

representation of exchange will likely be more applicable in these locations. 543

544

Regardless of the predictability of bed form geometry and spacing, the associated hydrodynamic 545

HEF may induce only limited lengths of exchange in the subsurface because much of the 546

exchange dynamics are expected to be vertical rather than lateral. Exchange lateral to the channel 547

is more likely to be driven by hydrostatic gradients set up across meander bends or bars (as 548

described above). Hydrodynamic HEF will contribute to, but be only one component of, total 549

HEF in natural channels, and its importance will be dictated by both channel hydraulics and, if 550

present, competing hydrostatic factors that can create steeper head gradients. 551

552

9.13.2.4. The particle scale – turbulent diffusion 553

554

At the particle scale on streambeds, turbulent diffusion is significantly influenced by the size and 555

arrangement of surface sediment. Because turbulent diffusion is induced by the momentum 556

transfer between the water column and the porous media, HEF due to turbulent diffusion is a 557

function of the decreasing velocity profile within the surface layers of the porous media (Shimizu 558

et al. 1990). Thus, the distribution of sediment at the surface will greatly influence the potential 559


20

for energy and mass transfer within this zone. Turbulent diffusion HEF is prominent in gravel 560

bed streams where surface pores are more likely to accommodate such open exchanges of 561

momentum across the bed (Tonina and Buffington, 2009). Beds composed of sand particle sizes 562

and smaller provide too much resistance to the momentum exchange between the water column 563

and the bed. Hence, turbulent diffusion is more likely to be an important component of HEF in 564

low order, high-gradient streams (Figure 2B). Careful theoretical and empirical research on 565

turbulent diffusion has been conducted largely on planar beds (Shimizu et al. 1990; Habel et al. 566

2002). Therefore, in the complex bed topography of typical gravel channels, turbulent diffusion 567

will be a component of HEF, likely not the singular driver of HEF. 568

569

9.13.3. Discussion 570

571

9.13.3.1 Multiple features acting in concert 572

573

In the examples presented above (Figures 1, 3–9), we have mostly focused on single types of 574

channel morphologic features that drive or enable hydrostatic and hydrodynamic HEF. However, 575

these features never occur in isolation. Rather, a single stream reach will typically contain many 576

of the morphologic features described above. Interactions among these features are likely to be 577

important in determining the actual HEF in any given stream reach. In some cases, the effects of 578

multiple features could be additive and result in higher HEF than if they did not co-occur. For 579

example, cross-valley flow paths between main channels and floodplain spring brooks can be 580

accentuated by riffles (Figure 7C). However, interactive effects could also cancel, for example 581

where riffles at the inflection points of meander bends reduce head gradients through point bars 582

(Figure 6C). The interactions between different processes driving or enabling HEF is complex, 583

and to some degree, site specific, making it difficult to quantify the effects of these interactions. 584

Because of these difficulties, there are relatively few comparative studies that have examined 585

multiple processes concurrently, within natural stream channels and attempted to evaluate the net 586

effect of each process on the total HEF within stream reaches. 587

588

Sensitivity analyses with groundwater flow models calibrated to simulate HEF in a studied 589

stream reach provide one opportunity to examine the relative importance of channel morphologic 590


21

features on HEF where multiple features are present in a single reach. For example, Kasahara 591

and Wondzell (2003) examined a number of channel morphologic features among stream reaches 592

of different sizes in a mountainous stream network under conditions of summer baseflow 593

discharge. In all cases, the single strongest driver determining the amount of HEF occurring 594

within the simulated stream reaches was the change in longitudinal gradient over step-pool 595

sequences in the 2nd-order channel (Figure 10A) and pool-riffle sequences in the 5th-order 596

channel. The shape of the hyporheic flow net in the 5th-order stream, however, was strongly 597

controlled by the presence, location, and relative elevation difference between water in the main 598

channel and the back channels (Figure 10B). Similarly, Cardenas et al. (2004) examined 599

sediment heterogeneity, size of bedforms, and both longitudinal and lateral head gradients in a 600

low gradient, sand bed stream. They found that HEF was greater where beforms had higher 601

amplitude and were more closely spaced. Spatial heterogeniety in K increased HEF relative to 602

homogeneous simulations, as did inclusion of lateral head gradients, but the effect was small 603

relative to the effect of the size and spacing of bedforms. 604

605

Channel morphologic features can interact with changes in steam stage and lateral groundwater 606

inputs in ways that can substantially influence the amount of HEF over time, across seasons or 607

within a single storm event. Storey et al. (2003) examined HEF in a pool-riffle sequence at both 608

high- and low-baseflow discharge. At high stage, the stream tended to “drown” the riffle, 609

substantially reducing the change in the longitudinal gradient over the pool-riffle sequence and 610

thus reducing HEF. In contrast, at low stage, the water surface more closely followed the 611

streambed topography, thus creating steeper head gradients that supported more HEF. Storey et 612

al. (2003) also showed that lateral inputs during the wet season were sufficient to eliminate most 613

of the HEF through the riffle. Cardenas and Wilson (2006) showed that low rates of groundwater 614

discharge limited the extent of the HZ formed by the hydrodynamics of stream bedforms, and 615

that high rates of groundwater discharge could completely eliminate HEF. 616

617

We know of only one study comparing the relative influence of hydrostatic and hydrodynamic 618

effects. In a flume, Tonina and Buffington (2007) investigated the control of total head (i.e., 619

including dynamic head) in driving hyporheic exchange. Their results suggested that there are 620


22

specific locations within channels where the velocity head can provide additional potential and 621

thereby influence the pattern of hyporheic exchange. 622

623

9.13.3.2 Change in processes driving HEF through the stream network 624

625

Hyporheic exchange will vary widely across the sequence of channel types found in stream 626

networks (Figure 2; Buffington and Tonina, 2009). Channel networks generally follow a pattern 627

of steep headwaters to low-gradient reaches downstream. In mountain stream networks in 628

particular, gradient changes are expected to be accompanied by channel morphology changes 629

resulting in a sequence of distinct channel morphologies (Figure 2A). Obviously, bedrock 630

reaches have negligible hyporheic zones (Gooseff et al. 2005; Wondzell, 2006). We are unaware 631

of any studies of HEF in colluvial and cascade channel morphologies, however the extremely 632

high longitudinal gradients of these channels likely result in high velocity underflow which has 633

been shown to restrict the extent of the hyporheic zone (Storey et al. 2003). Also, the relatively 634

disorganized structure of the bed sediment prevents development of stepped water surface 635

profiles so that hydrostatically driven exchange due to longitudinal changes in gradient will 636

likely be low. Turbulent diffusion is likely to be a primary driver of HEF (Figure 2B). 637

638

Free-formed step-pool channels occur at slightly lower gradients (Figure 2A). These channels 639

have well-organized structure with periodic spacing of both steps and pools (Chin, 2002; Wohl 640

and Merrit, 2008) that have been shown to be primary drivers of HEF (Figure 2B; Kasahara and 641

Wondzell, 2003). The addition of large wood can substantially increase sediment storage 642

(Nakamura and Swanson, 1993; Montgomery et al. 1996), the development of step-pool 643

structure, and the extent, amount and residence times of HEF in these stream reaches (Wondzell, 644

2006). Other hydrostatic factors tend to have less dominance on HEF; these reaches have low 645

sinuosity so meander bends are uncommon and steep longitudinal gradients limit the potential 646

for back channels to create lateral HEF flow paths. 647

648

We are unaware of any published studies examining HEF in plane-bed channels. However, we 649

expect HEF to be lower than in either step-pool or pool-riffle channels (Figure 2B). The 650

streambed tends to be smoothly graded in these channels as suggested by their name, and there is 651


23

low spatial heterogeneity in surface texture (Buffington and Montgomery, 1999). Pools are 652

widely spaced, and both steps and riffles are rare. While these channels occur as free-formed 653

morphologies, pool-riffle channels can be converted to plane-bed channels by land use practices 654

that increase sediment supply and through the direct removal of large wood, with concurrent 655

decreases in HEF. 656

657

Lower in the stream network, channels tend have lower longitudinal gradients (Figure 2A), and 658

even in mountainous areas, unconstrained stream reaches become increasingly common. Channel 659

planforms can be quite complex in these rivers and as a consequence, a wide array of channel 660

geomorphic features influences HEF. Braided and anastomosing channels may form where 661

sediment loads are high and stream banks are erodible; the complex of channels likely leads to 662

substantial HEF through islands. Meandering channels form under lower sediment loads and 663

where banks are more stable. Meandering channels typically have pool-riffle morphologies, 664

although complexes of secondary channels, back channels, and paleochannels are common, a 665

legacy of past floods, channel avulsions and overbank deposition. Because most HEF occurs 666

along short, near stream flow paths, riffles are the dominant feature determining the amount of 667

HEF (Kasahara and Wondzell, 2003). However, the shape of the hyporheic flow net and the 668

residence time distribution of HEF will be strongly influenced by the complex of channel 669

planforms. Finally, hydrodynamic processes are expected to dominate in streams with relatively 670

mobile streambeds characterized by dune-ripple bedforms. These streams have low longitudinal 671

gradients and therefore channel morphologic features tend not to create steep hydrostatic head 672

gradients (Figure 2B). 673

674

Other exchange processes are likely to be related to specific conditions. Turn-over exchange will 675

only occur when bed material is mobile – a characteristic feature of both anastomosing and dune-676

ripple channels. Transient exchange will only be appreciable during wet catchment conditions, 677

when channel stage is high and surrounding groundwater tables are comparatively low. 678

However, transient exchange may be a dominant form of HEF in regulated rivers where stage 679

fluctuates over daily cycles due to hydroelectric generation. Turbulent diffusion, on the other 680

hand, is likely to occur in gravel bed sections of the network, likely with the greatest potential 681


24

influence in either cascade or plane-bed sections of stream networks where stream velocities are 682

expected to be high. 683

684

9.13.4. Conclusion 685

686

Hyporheic exchange results from distinct processes, and the relations between those processes 687

and geomorphology are well understood from a mechanistic perspective. Thus, geomorphology 688

provides a critical framework to understand hyporheic processes and how they change with 689

location within a stream network, and over time in response to changes in stream discharge and 690

catchment wetness. To the degree that these geomorphic patterns are predictable, they provide 691

the foundation for hydrologists to make general predictions of the relative importance of the 692

hyporheic zone at the scale of entire catchments. Reach to reach variability is high in stream 693

networks, however, so understanding HEF at the reach scale continues to require detailed study 694

of specific stream reaches. These studies are difficult and current methodological approaches are 695

insufficient to fully examine the full suite of processes that account for patterns of HEF in any 696

specific stream reach. Consequently, hyporheic studies tend to focus on a single factor, or at 697

most a small subset of the factors driving HEF. Hyporheic researchers recognize that such 698

studies are incomplete. Detailed, holistic understanding of the importance of different processes 699

in driving HEF, how the relative importance of these processes changes with location in the 700

stream network, with the specific structure of any given stream reach, and with changes in 701

discharge and lateral groundwater inputs remains elusive. 702


25

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1019


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Figure Legends: 1020

1021

Figure 1. Idealized conceptual model of nested hyporheic flow paths as influenced by step-pool 1022

or pool-riffle sequences. A) Plan view showing arcuate HEF flow paths through the adjacent 1023

floodplain created by the change in the longitudinal gradient over the pool-riffle sequence where 1024

the amount of HEF is proportional to the head gradient. B) Longitudinal-section along the 1025

thalweg of the stream showing the vertical component of HEF flows through the streambed. 1026

1027

Figure 2. A) Hypothetical distribution of channel types along a stream profile in a mountainous 1028

stream catchment (redrawn from Montgomery and Buffington, 1997), and B) the corresponding 1029

relative contribution of turbulent diffusion and both hydrostatic or hydrodynamic processes to 1030

the total amount of HEF occurring within a stream reach. (Note that boundaries between channel 1031

types are often less distinct than shown here and that a range of conditions occurs within each 1032

category, thus the contribution of each process varies both among and within each channel type). 1033

1034

Figure 3. Idealized conceptual model of the influence of lateral inflows on hyporheic exchange 1035

flows. A) A high gradient stream where floodplain alluvium has relatively high saturated 1036

hydraulic conductivity under relatively dry conditions when lateral inputs are low and easily 1037

transported down valley via subsurface flow. Lateral inputs still reach the stream, but are 1038

diverted towards zones with hyporheic upwelling. B) A low gradient stream where floodplain 1039

alluvium has relatively low saturated hydraulic conductivity under relatively wet conditions 1040

when lateral inputs are sufficiently large to overwhelm down valley transport, causing lateral 1041

inputs to cross the valley toward the stream. Lateral inputs severely restrict hyporheic exchange 1042

flows. Legend follows Figure 1. 1043

1044

Figure 4. Idealized conceptual model of the influence changing stream stage on transient 1045

hyporheic exchange. A and B) A losing reach at low baseflow is converted to a gaining reach 1046

during a storm because precipitation recharge and lateral inputs of hillslope water increase water 1047

table elevations more than the corresponding increase in the stream stage. The original stream 1048

and water table position from 4A is shown in 4B for reference (light grey line). C & D) An 1049

example of a river where changes in stream stage result from snow melt, tidal influences, or dam 1050


37

releases far upstream. Increased stream stage causes stream water to flow into the adjacent 1051

aquifer creating a losing stream reach. Conversely, decreased stage leads to drainage of the 1052

aquifer creating a gaining reach. Alternating increases and decreases in stream stage leads to 1053

transient hyporheic exchange. The neutral condition (where stream stage is equal to the water 1054

table elevation) is shown for reference (black and white dashed line). Legend follows Figure 1. 1055

1056

Figure 5. Idealized conceptual model of the influence of the change in saturated cross-sectional 1057

area of the floodplain on hyporheic exchange flows. A) The influence of change in valley 1058

constraint with downwelling at the head of an unconstrained reach and upwelling at the 1059

downstream end of the reach caused by the transition from narrow bedrock gorges to wide 1060

alluvial valley floors. B) The influence of variations in depth to bedrock forcing upwelling 1061

upstream of a bedrock sill and downwelling downstream, where the depth of alluvium again 1062

increases. Legend follows Figure 1. 1063

1064

Figure 6. Idealized conceptual model of the influence of meander bends on hyporheic exchange 1065

flow. A) Simple, low radius meander with HEF traversing the point bar and floodplain. B) High 1066

radius meander with incipient meander-cutoff, where the short distance across the neck leads to 1067

much higher head gradients and thus greater HEF through the neck than the remainder of the 1068

meander bar. C) Meander bend with riffles located at the inflections between adjacent meanders 1069

so that head gradients through the point bar are low and much of the HEF occurs around the 1070

riffles, driven by longitudinal changes in gradient. Legend follows Figure 1. 1071

1072

Figure 7. Idealized conceptual model of the influence of back channels on hyporheic exchange 1073

flows. A) A back channel is incised below the water table, acts as a drain, and creats head 1074

gradients from the main channel to the back channel. B) A back channel is plugged near its 1075

downstream end, conducts water onto the floodplain, raises the water table and creats head 1076

gradienets from the back channel to the main channel. C) Complex pattern of HEF caused by 1077

interactions between a riffle in the main channel and a back channel. Paleochannels (dashed 1078

lines) support preferential flow. Legend follows Figure 1. 1079

1080


38

Figure 8. Idealized conceptual model of the influence of mid-stream islands on hyporheic 1081

exchange flows. A) Parallel and smooth longitudinal gradients in the channels on both sides of 1082

the island create HEF flow paths that parallel stream flow. B) Riffles at the head of the island 1083

enhance head gradients leading to greater HEF. C) Offset riffles create strong cross-island head 1084

gradients and flow paths, resulting in more HEF but with shorter flow path lengths and residence 1085

times. Legend follows Figure 1. 1086

1087

Figure 9. Idealized longitudinal-section in the center of a straight stream channel with bedforms 1088

(triangular dunes) showing the interaction with stream flow that creates regions of low- and high-1089

pressure on the streambed which drive HEF. Non-hyporheic subsurface flows, known as 1090

underflow (dashed arrows), are present beneath the hyporheic zone. Legend follows Figure 1. 1091

1092

Figure 10. Examples of complex hyporheic flow paths resulting from interactions between 1093

channel morphologic features: A) a steep, 2nd-order step-pool channel with abundant large wood, 1094

and B) a moderate gradient, 5th-order pool-riffle channel with two major spring brooks. Note the 1095

difference in spatial scale between the two stream reaches. Letters indicate morphologic features 1096

driving HEF: S – steps; R – riffles; M – meander bends; B – back channels / spring brooks; I – 1097

islands; and T – a steep riffle at the mouth of a tributary. Equipotential intervals (dashed lines) 1098

are 0.2 m. Hyporheic flow paths (arrows) are hand drawn to indicate general direction of 1099

hyporheic flow through the valley floor. 1100

Pool - riffle - pool sequence

B

Pool

Riffle

Pool

A

Floodplain

ActiveChannel

WettedChannel

Riffle

Subsurfaceflow path

Hill-slope

Hollow

Colluvial-bedrock

Cascade

Step-pool Plane-

bed Pool-riffleDune-ripple

Hydrodynamiccontribution

Turbulentdiffusion

Longitudinal stream profile through stream network

Hydrostaticcontribution

Rela

tive

HE

F

A

B

A

Higher gradient, low lateral inputs

B

Lower gradient, high lateral inputs

A B Precipitation

Late

ralin

puts

C D

Stage Increase

Stage Increase

Stage Decrease

A

B

Bedrock sill

Unconstrained stream reachBedrockgorge

B

C

A

Pool / Run

B

C

A

B

C

A

B

C

Stream flow

Scale = 50 m

B

R

R

R

R

T

B

B

R

R

M

M

I

BedrockLog

Valley floor alluviumWetted stream channel

Hillslope or terraceBack channel

Equipotential (0.2 m) Hyporheic flow path

Scale = 20 m

A

S

S

S

S

SS

IB

S S

S

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