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Home , Channel (geography), Santee River

Fluvial-tidal transition channel Channel geomorphology along the fluvial- tidal transition, Santee , USA

Raymond Torres† Department of and Ocean Sciences, University of , 701 Sumter Street, EWS617, Columbia, South Carolina 29208, USA

ABSTRACT INTRODUCTION Pittaluga et al., 2015). Since terrestrial runoff, tidal forcing, and storm surge vary with time, There exists a rich understanding of chan- Along the river continuum, a single channel the tidal effects on channel form vary along the nel forms and processes for with uni- can transition from fluvial to tidal dominance in channel (Wright et al., 1973; Dalrymple and directional flows, and for their estuarine flow and sedimentary processes, and in benthic Choi, 2007). For instance, river flow responses components with bidirectional flows. On the ecology (Dalrymple and Choi, 2007; Jablon- to tidal oscillations can vary temporally with lo- other hand, complementary insight on the ski and Dalrymple, 2016). In particular, many cal weather effects (

GSA Bulletin; November/December 2017; v. 129; no. 11/12; p. 1681–1691; doi: 10.1130/B31649.1; 4 figures; Data Repository item 2017186; published online 30 June 2017.

Geological Society of America Bulletin, v. 129, no. 11/12 1681 © 2017 The Authors. Gold Open Access: This paper is published under the terms of the CC-BY license.

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found systematic channel variations in 122 estuary of a coastal plain river (sensu Savenije flux over 20 yr time intervals to estimate the cross sections along an 18 km channel reach. et al., 2008). The detailed observations and pre-European, predam, and postdam conditions. However, the study site only experienced the ef- analyses presented here will help improve our The corresponding values are 2.24, 5.80, and fects of the higher tides and no reversals understanding of transition zone morphody- 0.81 Mt yr–1, respectively. At this time, it is not due to the placement of a downstream weir, and namics, and they will help to support the util- known if the channel has attained morphody- as a consequence, those findings provide lim- ity of generalized facies models that are taken namic equilibrium in response to dam effects. ited insight. Further, Gardner and Bohn (1980) to represent average conditions over much lon- However, a comparison of aerial images from and later Ensign et al. (2013) provided a con- ger time scales (e.g., Bokuniewicz, 1995; Blum 1939 and 2011 indicates that the ceptual view on how terrestrial and tidal chan- and Tornqvist, 2000; Cattaneo and Steel, 2003; and width have remained relatively stable over nel cross sections should differ by highlighting Phillips and Slattery, 2007; van den Berg et al., 72 yr, except for ~12 m of widening near the the characteristic change in channel properties 2007; Dalrymple and Choi, 2007; Jablonski rediversion where flows are returned to the due to the onset of tidal effects, with particular and Dalrymple, 2016). Also, given the dynamic system (Fig. 1). emphasis on width. Inokuchi (1989) and later nature of the transition channel, it is likely that This study is focused on an ~64 km reach Nittrouer et al. (2011b) examined particle size, both process and form respond at time scales between 30 km and 94 km from the mouth. channel cross sections, and long profiles of commensurate with land-use change, climate For clarity, hereafter “mouth” is taken to mean hundreds of kilometers of the lower Mississippi change, and sea-level rise (Florsheim et al., “estuary mouth,” where the estuary meets the River. Overall, inconsistencies in cali- 2008). Moreover, with the relative uniformity of ocean (sensu Savenije, 2012). The site was cho- ber and bed slope were interpreted as resulting the modern southeastern U.S. coastal plain land- sen because a U.S. Geological Survey (USGS) from coastal backwater effects (e.g., Fernandes scape (Hayes, 1994), it is likely that the channel gauging station at 59 km upstream of the mouth et al., 2016). Tidal distortion and tidal wave- features reported here may apply to many rivers shows that the reach has intermittent tidal- and length effects on channel properties were re- that to the southeastern U.S. Atlantic non-tidal-dominant flow conditions. The down- ported by Wright et al. (1973) and expressed as (Fig. 1), and perhaps to rivers that tra- limit of the study reach was set by the equilibrium between the tidal prism and equal verse coastal plains in general, landscapes that position where channels begin, by work per unit channel bed area for the channel occupy 5.7 × 106 km2 worldwide (Colquhoun, changes in land cover, and by the presence of of a funnel-shaped estuary. Phillips and Slattery 1968). Overall, this work attempts to fill a gap dikes along the channel. The upstream limit was (2008) analyzed long- and cross-channel river in knowledge of the geomorphic structure of the set by the position where <0.04 m tidal oscilla- profiles to assess the role of topography and fluvial-tidal transition channel, and it is driven tions were detected during low-flow conditions. antecedent in sediment bottlenecks by the hypothesis: Geomorphic discontinuities Hereafter, unless otherwise noted, all distances upstream of the modern fluvial-tidal transi- define the channel reach linking upstream flu- are reported as along channel, and relative to the tion, and their effects on the sediment budget vial-dominant and downstream tidal-dominant Highway 17A bridge (“Bridge” of Fig. 1) near and channel morphology further downstream. parts of the river continuum. Jamestown, South Carolina; distances upstream These effects were expressed as a systematic of the bridge are negative. This reference is used decline in channel slope and stream power. STUDY SITE because the bridge is easily identified in images, Together, these studies show that fluvial-tidal and it precludes arbitrary measures of distance transition systems have a range of dynamic at- The Santee River is one of 20 larger and many along a complex distributary system (Fig. 1). tributes, but few details are known about active smaller rivers that discharge to the Atlantic Ocean Discontinuous outcrops of the Santee For- channel dimensions and geomorphology. along an ~1400 km section of coast (Fig. 1). The mation (Siple, 1960) were observed along part A noteworthy point about these modern drainage area is ~38,000 km2, with headwaters of the riverbed and banks (Fig. 1), and they con- ­process-based studies is that most are from in the Blue Ridge and Provinces, and sist of indurated gray to buff massive limestone field sites with rivers that traverse a coastal it is one of the larger river systems in the south- interspersed with layers of weakly cemented plain, while the conceptual models published eastern (McCarney-Castle­ et al., lithic limestone, dipping <10° east. In particu- elsewhere to describe transition zone processes, 2010). The Santee forms at the of lar, three outcrops of 2–3 km in length occur geomorphology, and sedimentary facies were the Wateree and Congaree systems at ~230 km along the study reach but are limited to down- almost exclusively developed for submerged along channel from the coast, and it flows across stream of the bridge. The three occurrences of river valleys or funnel-shaped not the South Carolina Coastal Plain. Further down- bedrock banks are at 3 km to 7 km, 9 km to necessarily on the coastal plain (e.g., Dalrym- stream, the Santee bifurcates at the apex of the 11 km, and 19 km to 22 km (Fig. 1). Hence, in ple and Choi, 2007; Jablonski and Dalrymple, Santee , ~22 km inland, giving rise these bedrock-influenced reaches, the channel 2016). This distinction is important because to the North and South Santee distributary chan- adjustments to modern conditions have been the latter will have stronger upstream channel nels, and several smaller ones (Fig. 1). Kjerve limited to the alluvial north . Bedrock area convergence (e.g., Dalrymple and Choi, and Greer (1978) reported that during mean an- banks were not observed upstream of the bridge 2007), while the low-gradient coastal plain riv- nual discharge, 73% of the flow occurs through despite a determined search. Together, these ers likely will have more subtle convergence. the North Santee. observations indicate that the lower study reach The ways in which these differences translate The Santee was dammed in 1942 at ~150 km between 3 km to 22 km can be characterized as to processes, forms, and facies are not well un- from the coast, along channel, and flows were a mixed alluvial-bedrock river (Howard, 1998; derstood, and few data on channel form exist to redirected back to the Santee at 77 km (Fig. 1; Turkowski et al., 2008), while further upstream, make a robust comparison. Kjerve and Greer, 1978). Despite the presence it is alluvial (Fig. 1). The purpose of this study was to investigate of the dam, sediment flux to the coast is esti- Tree fall, undermined or leaning trees, and in- the modern geomorphic structure of a channel mated to persist at 0.86 Mt yr–1 (Milliman and tact trees in the channel indicate that parts of the reach that links fluvial-dominated and tidal- Farnsworth, 2005). Also, McCarney-Castle­ et system are actively shifting or widening, or both, dominated river segments in the upper alluvial al. (2010) conducted simulations of sediment particularly at about -25 km (Fig. 1), where the

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83°W 82°W 81°W 80°W 79°W 78°W 77°W 76°W 75°W

38°N James

37°N

36°N

35°N

34°N Ocean

33°N Atlantic

Fluvial-tidal 32°N Altamaha Coastal Plain N Tr ansition 31°N

0 50 100 200 km X X

North Santee

Bridge South

Bedrock outcrop in channel N Atlantic Distance indicator, 10 km increments from bridge Ocean X Observation sites of Yankovsky et al. (2012) 10 km

Figure 1. The of the lower Santee River; note that fluctuations in river width are representative but not to scale. The white dashed segments on the main stem indicate observed bedrock exposures, and data from these reaches are shown with gray symbols in the following figures. Tick lines across the main channel are in 10 km increments of river distances relative to the bridge-river intersection (lat. 33.305°N, long. 79.678°W); upstream values are reported as negative distance. The geomorphic fluvial-tidal transition reach is centered at ~68 km from the mouth, or between -17 km and -4 km (see also Fig. 2B). The U.S. Geological Survey (USGS) gauging station (#02171700) is on a bridge pier. The X symbols depict the instrument placements of Yankovsky et al. (2012) at -4.5 and 1.6 km. Inset: The Santee River in the context of the southeastern coastal plain, with arrow indicating mouth location, and perimeter divisions given in degrees. The Santee, Altamaha, and James Rivers are identified accordingly. Dark-gray area represents the combined Piedmont and Appalachian Provinces; light-gray depicts the coastal plain; the boundary between them is the fall line.

canal delivers water to the main channel. Bed rounding , and their net contributions 0.54 m, MSL -0.05 m, and MLW at -0.63 m sediment from 42 grab samples between -12 km to flow are assumed to be negligible. (NAVD88). Tidal oscillations at the USGS sta- and 8 km (Fig. 1) consisted of nearly pure quartz The USGS maintains the Jamestown gauging tion vary with discharge but range from zero at

with a group average D50 of 0.67 mm. The station at the bridge (#02171700, Fig. 1) with a high flows to 1.3 m under low flows. Mean annual salinity within 3 km of the bridge was zero dur- pressure transducer and velocimeter that record discharge from 1987 to 2005 was ~311 m3 s–1. ing low river discharge (Yankovsky et al., 2012), data at 0.25 h intervals; all references to dis- For discharges of ~300 m3 s–1 the semidiur- and Hockensmith (2004) reported that the zero charge are from this station. The corresponding nal, mixed tide range varied but was typically isoline is limited to the distributary channels, stage datum is at 0.33 m, with mean high water <0.6 m, with highs and lows more strongly de- within 20 km of the mouth. Therefore, this study (MHW), mean sea level (MSL), and mean low veloped on tides. Yankovsky et al. (2012) reach is part of a tidal freshwater system (e.g., water (MLW) at 0.50 m, 0.31 m, and 0.16 m, reported that a typical flood tide lasts ~2 h, while Odum, 1988). There are seven along respectively (relative to North American Ver- the ebb tide lasts ~6 h, and high-tide velocities the study reach, all with mouth widths <6 m and tical Datum of 1988 [NAVD88]). Tidal range appear quasi-steady for 5–6 h, coincident with their respective drainage areas limited to the sur- at the estuary mouth is ~1.16 m, with MHW the duration of the ebb tide. The flood tide or

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upstream currents may reach 0.3 m s–1, but these peak values are short-lived. The ebb or down- 1000 5 (m)

stream currents can be much higher, but for d A 4 frequently occurring annual average discharges, 800 they are ~0.3 m s–1, comparable to the upstream 3 tidal velocities. 2 At this study reach, Yankovsky et al. (2012) 600

made simultaneous measurements of velocity 1 Average Depth, profile and stage at two locations (“X” symbols 400 at -4.6 km and 1.5 km in Fig. 1). They found Image Width (m) that tidal amplitude did not change between observations sites 6.1 km apart, despite having 200 the water depth and flow area decline by nearly 50% in the upstream direction. They also esti- 0 mated tidal energy dissipation and reported that -40 -30 -20 -10 010203040 the downstream location had a nondimensional 4 tidal velocity that was more than double that B from the shallower upstream site. Moreover, they characterized the flow regime as having a 0 substantial phase lag between free surface and velocity fluctuations. These analyses highlight the dominant effect of friction and dissipation -4 5

over tidal inertial effects, and the larger phase σ (m) 4 lag at the shallower upstream end, which indi-

cated higher tidal energy dissipation relative to Bed Elevation (m) 3 -8 the downstream site. Based on these findings, 2 the local flow conditions in proximity to the 1 bridge (Fig. 1) were characterized as strongly -12 0

dissipative and weakly convergent (after Lan- Standard Deviation, zoni and Seminara, 1998). -40 -30 -20 -10 010203040

METHODS 8 C

Bank-top width was estimated by two meth- ods, from a river survey and from the February 4 2013 image of Google Earth reflecting high-flow conditions. For the latter, the canopy-to-canopy width was measured at 1 km intervals over an 0 86 km reach from the mouth to 34 km upstream of the bridge (Fig. 1). Width in places without tree canopy was measured directly from the im- -4

age; measurements for canopy-covered banks /Water Surface Elevation (m) assumed the bank extended 5 m landward. For the bifurcating North and South Santee Rivers, -40 -30 -20 -10 010203040 the channel width was taken as the sum of both Along channel Distance (km) widths at the corresponding distance (e.g., after Pethick, 1992; Davies and Woodroffe, 2010). Figure 2 (Continued on facing page). Along-channel variations in channel geomorphic prop- Width was also measured directly at ~1 km erties. The bedrock-influenced measurements are shown with gray symbols, and the flu- intervals along the study reach, at each cross vial-tidal transition zone extends from -4 km to -17 km. (A) Channel width (circles) from section. A Trimble R8 global positioning sys- satellite imagery showing typical exponential decline. Inset (plus signs) shows average chan- tem (GPS) system and a Seafloor SonarMite nel depth. (B) Channel bed elevation point and linear curve fit (line) along the river center (Hydrolite-TM) echo sounder were installed line. Note the discontinuous but extensive interpreted as bedrock highs. Inset is on a 5.3-m-long vessel to acquire bathymetry standard deviation (line) of the bed elevations over a 5 km window. Note downstream values at cross-section sites and along the channel. All are generally higher than upstream. (C) Along-channel thalweg elevation (filled circles). data were collected at 1 m spacing with a vessel Above the thalweg trend line are the corresponding free surface elevations (open circles) speed <2 m s–1. The base station for the real-time taken during a flood event, where both data sets are referenced to NAVD88, and both have kinematic survey was part of the South Carolina linear trend line fits. Gray symbols correspond to locations of bedrock channel. GPS Virtual Reference Station (VRS) connected through cellular internet connection (Lapine and Wellslager, 2007). The GPS system gives an ac-

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curacy of 10 mm ± 1 ppm root mean square er- USGS stage was relatively constant. The free bank tops. All cross-section metrics for bank- ror (RMSE) horizontally and 20 mm ± 1 ppm surface profile data are presented as water sur- top conditions were assessed with WinXSPro RMSE vertically for static surveys. The - face elevations averaged over a 20 m reach (Hardy et al., 2005). ings for bathymetry ran at 6 Hz with a frequency (s < 0.014 m) upstream and downstream of each Cross-section area-stage relations were of 235 kHz, and accuracy was <10 mm (or 0.1% cross-section profile location. Cross-section further evaluated by converting area and cor- of depth), with a measurement capacity ranging surveys took place on 11–12 March, 31 August, responding stage to a proportion and plotting from 0.3 m to 75 m. The VRS-GPS and echo 17 November, and 20–21 December in 2013. all curves as values ranging from 0.0 to 1.0. sounder were integrated with a Trimble survey Among these days, river stage varied by 2.2 m, The concave-up curves with stage as the inde- controller TSC2 via wireless connection. A Wil- but 41 of the 64 cross sections were surveyed at pendent variable were well represented with a son Electronic Sleek Wireless Signal Booster or about bank-top stage. In most cases, the verti- power function: A = aSb, where A = cross sec- was used to augment the requisite internet con- cal banks allowed the echo sounder transducer tion area (m2), S = stage (m), a = 1, and b is nection to the VRS. All field survey data were to approach to within 1–2 m. When the banks the fitting parameter. This facet of the analysis referenced to Universal Transverse Mercator 17 were more gradual, the hydrographic transect provides a metric for comparing channel shapes North and to NAVD88. ended at 0.5 m water depth, and cross-section via a “shape factor,” b. The b values ranged Multiple hydrographic surveys were con- bathymetry was augmented with channel bank from 1.0 to 4.0, corresponding to rectangular ducted during 2013; the centerline survey was surveys acquired with a 5 m stadia rod and to “wide V” shapes, respectively. The along- conducted during flood conditions when the handheld leveling scope on the dry bank to the channel survey data were used to assess river sinuosity (S) expressed as the ratio of chan- nel distance to straight-line distance between 500 points. S was computed over 1-km-long reaches D centered about each cross section, or about ten 400 times the average river width (Leopold and Wol- ) 2 man, 1957). (m

A 300 RESULTS

200 3 Channel Planform S

Bankfull Area, 2 The 64-km-long study reach has a cor- 100 responding length of 45 km, giving a 1 Sinuosity, sinuosity of 1.42, characteristic of a - 0 ing system. In finer detail, however, the channel has a mix of straight and meandering reaches, -40 -30 -20 -10 010203040 and in the meandering segments, there is a mix 4 of irregular and regular patterns, with the latter b

V-shaped being more prevalent downstream. The larg- 6 3 est are centered at 24 km, near the E confluence of the (Fig. 1). Local 5 2 kilometer-scale sinuosity, S, varies from 1.0 to 3.8, but 78% of the values are less than 1.2, e.g., (m) Shape Factor - R from -16 km to 23 km (Fig. 2D).1 Therefore, 4 1 Rectangular the middle part of the reach is mostly straight, but beyond that, it is meandering. Given the lim- 3 ited extent of bedrock outcrops, the meandering part of the channel system can be characterized as unconfined. Finally, it is noteworthy that the Hydraulic Radius, 2 highest downstream S values coincide with the downstream end of the bedrock outcrops, be- 1 tween 23 km to 27 km (Figs. 1 and 2D).

-40 -30 -20 -10 010203040 Channel Width and Depth Along Channel Distance (km) Figure 2 (Continued). (D) Cross-section area of flow, A, along channel (circles). Note The Santee system exhibits a downstream the transition reach characterized by a trough in values between end points of locally exponential increase in width to the mouth. For high area. The inset highlights channel sinuosity, S (plus signs). (E) Hydraulic radius, R (circles). Note the upstream peak at -17 km, taken as supporting evidence of the upstream limit of the fluvial-tidal transition. Inset depicts along-channel variations 1GSA Data Repository item 2017186, a of the in channel shape, expressed as the shape factor, b (plus signs; see Methods). Note the data presented in Figure 2, is available at http://www transition to more rectangular-shaped channels at -17 km. Bedrock is shown with .geosociety.org/datarepository/2017 or by request to gray symbols. [email protected].

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instance, the imagery data (Fig. 2A) show that within a narrow range of ~2 m. Therefore, the fects of the upstream channel deeps (Fig. 2B). at -3 km, the channel width, w, is 93 m, but overall downstream trend in d when account- From -4 km, the values increase, and this trend at the mouth, 62 km downstream, it is 825 m. ing for bedrock and river-bend effects reveals a continues to the end of the survey. In particular, The width convergence length, the limit of ex- 22 km section of increasing average depth fol- at 3 km to 4 km, values increase sharply from ponentially decreasing values, is estimated to lowed by a decline over a distance of 17 km and 0.5 m to 1.6 m, likely related to bedrock influ- be at least 94 km (after Savenije, 2012; based then a 25-km-long relatively uniform section ence. Thereafter, the values tend to increase on data from Yankovsky et al., 2012). Since (Fig. 2A). downstream to a maximum of 1.7 m, albeit with the overall study reach extends from 30 km to minor interruptions to the trend associated with 94 km inland, it follows that this study reach is Long Profile the bedrock plateaus. These observations dem- in the zone of channel convergence. Upstream onstrate that the downstream riverbed is more convergence in width adheres to the expres- The center-line riverbed elevations are irregular relative to the upstream. sion y = 6.53 exp(0.092x) + 90, with r2 = 0.94 highly variable, ranging from -9.4 m to 3.1 m The along-channel thalweg elevations show a (Fig. 2A). These data and curve fit provide a (NAVD88). Along the profile, there are several typical declining trend with a slope of 1.0 × 10-4 broader along-channel context in which to view local “deeps” taken as kilometer-scale channel (r2 = 0.78; Fig. 2C), i.e., a factor of ~10 greater the overall study. Thus, the study site is mainly reaches with a 3 m to 6 m departure below the relative to the center-line long profile. Patterns in along a channel reach of gradually declining surrounding riverbed (Fig. 2B). The long profile thalweg elevation show two breaks in slope: one width in the upstream direction, especially up- shows a quasi-linear decline in elevation down- corresponding to the upstream relatively uni- stream of the bridge. stream, with an overall slope of 5.7 × 10-5 and form decline to a -lying segment at- 6 km, There is a 25-km-long section, from -2 km r2 = 0.48. The declining trend, however, is inter- and the second transitioning to a highly variable to 23 km, with noteworthy departures in width rupted at 3 km to 4 km, where the downstream downstream decline at 3 km (Fig. 2C). Over the from the overall exponential trend (Fig. 2A). reach is offset by a 1.3-m-high step (Fig. 2B). upstream reach, from -34 km to -6 km, there is The locations and values of these departures The upstream and downstream reaches relative an ~4 m decline, giving a mean slope of 1.5 × are comparable in both the image and surveyed to 3 km have similar trend line fits with slopes 10-4 and r2 = 0.64. This condition is followed values. In particular, both sets of data have two of 7.4 × 10-5 and 9.2 × 10-5, respectively, and by an 8-km-long and relatively flat reach with distinct reaches of locally high values, one at corresponding r2 of 0.47 and 0.24, respectively. elevations between -2.3 m and -1.4 m. Further -2 km to 13 km and the other at 17 km to 23 km Further, along the upstream reach, the fluctua- downstream, the data exhibit strong fluctuations (Fig. 2A). The maximum departure from the tions in bed elevation are typically less than 2 m, that include an 8-km-long deep up to 5 m below trend in Figure 2A for the upstream reach occurs although channel deeps of up to 6 m occur (see the background elevation (centered at 16 km; at 5 km with a corresponding width of 213 m, -17 km to -13 km; Fig. 2B). In contrast, channel Figs. 2B and 2C). The highly variable but de- or ~130 m wider than expected from the expo- deeps in the downstream reach are longer, and clining elevation trend from 4 km to 30 km has nential fit. The downstream local maximum is at they occur more frequently. For instance, there a slope of 8.2 × 10-5, with relatively poor linear 20 km, and it is 75 m greater than background are ten deeps with more than 5 m of relief in the correlation coefficient, r2 = 0.19. As detected width. However, along the 4 km reach between downstream, whereas upstream, there are two. in the center-line profile (Fig. 2B), most of the these wider sections, width returns to the back- Among the types of variability in the profile higher thalweg elevations in the downstream ground trend (Fig. 2A). On the other hand, up- of riverbed elevation are several “plateaus” or reach coincide with bedrock outcrops. stream of these highs, there are relatively nomi- positive-relief structures that have relatively The corresponding water surface profile has nal fluctuations. Field observations indicate that smaller fluctuations along them, and they are two breaks in slope at -2 km and 4 km, and the aberrant highs in width coincide with the ~1 m to 3 m higher than the surrounding river- perhaps a third along the lower 4–6 km of the presence of the discontinuous bedrock outcrops bed (Fig. 2B). For instance, there are plateaus survey (Fig. 2C). The upstream reach from (see Fig. 2A), and both coincide with relatively from 3 km to 7 km, 9 km to 11 km, and at 19 km -30 km to -2 km has a linear trend line ex- straight parts of the channel (Fig. 2D). to 22 km. Comparable features are not pression of 3.39–1.72 × 10-4 (r2 = 0.99), while Average depth, d, is taken as the quotient apparent upstream, although a positive-relief downstream reach, from 4 km to 22 km has of bank-top cross-section area of flow, A, and structure at -24 km to -22 km resembles a pla- a trend of 3.84–6.02 × 10-5 (r2 = 0.98). This surveyed channel width. Values range from teau, but the top of this feature has an irregular break in slope is coincident with the transi- 0.8 m to 4.5 m, with a mean of 2.62 ± 0.85 m. trend (Fig. 2B). The downstream plateaus coin- tion to bedrock-influenced­ channel. Overall, Along-channel trends in d reveal two patterns of cide with the locations of bedrock outcrops with these profile conditions are characteristic of a variability (Fig. 2A). From -34 km to 3 km, d 2–4 km length (Figs. 2A and 2B), and together mild to milder slope transition that typically exhibits a weak parabolic-type response with a they illustrate the effects of a spatially variable gives rise to an M1-type free surface profile peak at -17 km. However, a pronounced depar- occurrence of resistant bedrock on riverbed (e.g., Dingman, 2009), most likely because the ture from this trend occurs with an abrupt but elevation. bedrock outcrops produce a backwater effect. persistent shoaling from -14 km to -11 km. The standard deviations of bed elevation Following the concepts of Parker (2004), the Further downstream, the variability has no clear based on a 5 km running mean range from backwater length L (the upstream distance of pattern and ranges from 1.2 m to 4.8 m, and 0.5 m to 1.7 m (Fig. 2B), and the along-channel channel flows affected by decreased free sur- the lower values correspond to the presence of trend has four distinct attributes. First, in the face slope) can be estimated as L = Dy/S, where bedrock (Fig. 2A). The higher values occur with upstream, there is a net decline from an initial Dy is the induced change in channel flow depth higher sinuosity S, but not exclusively (Fig. 2D) value of 0.7 m at -34 km to the minimum of at the transition, and S is the upstream riverbed as with Nittrouer et al. (2011b). Excluding the 0.5 m at -4 km. On this overall decline, there slope. On the other hand, this estimate indi- bedrock reaches and meander effects on depth is an extensive “high” of 0.5 m amplitude cates that L can be computed as the intersection at the downstream and more tide-influenced part between -23 km to -8 km, e.g., centered at of the upstream and downstream free surface of the reach, d is relatively uniform and remains -17 km. This feature corresponds to the ef- trend lines. It follows that L = 7.5 km, hence,

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L extends from bedrock ­outcrops at 3.5 km to these segments, -16 km to 8 km, A is highly influenced values are removed, the downstream -4 km (Figs. 1 and 2C). Close examination variable with a range of 210 m2 to 555 m2. R values have no clear trend. Therefore, from of Figure 2C shows that this backwater reach In particular, from -16 km to -14 km, A de- upstream to downstream, R, or channel effi- coincides with a local shoaling or flattening of clines by ~50% to 207 m2 and remains in this ciency, initially increases and then maintains the channel bed in the corresponding thalweg lower range to -10 km (Fig. 2D). From -7 km to a quasi-steady range. Downstream of 10 km, data (Fig. 2C). Therefore, the shoaling condi- -1 km A gradually increases, and from -1 km to however, the full range of variability occurs. As tion induced by backwater effects from the free 1 km it nearly triples, reaching a peak of 554 m2. is the case with A, the lowest R values are asso- surface decline in slope is limited to a 7.5 km Further downstream, from 3 km to 30 km, there ciated with bedrock outcrops (Fig. 2E). reach, and this value is a more conservative is a net decline to 230 m2 (Fig. 2D). Collectively, Another means of assessing channel cross- (longer) estimate than the method of Samuels the data have a trend mimicking y = –x2, but of section properties is with the rating curves of (1989). Further, using the same Dy/S approach course the downstream or declining part of this A. In all cases, plots of A as the independent of Yarnell (1934) reveals that the backwater ef- apparent trend is not sustainable because of variable versus stage height relative to the local fect of the bridge (Fig. 1) is on the order of the effects of the tidal prism on area (e.g., on thalweg give a characteristic concave-up trend. 1.3 km. Here Dy is based on the nearest mean the coastal plain, A typically increases expo- The same is true for the nondimensional repre- depth values upstream and downstream of nentially downstream; after Friedrichs, 2006). sentation of the same data normalized by each the bridge (Fig. 2A). The net result is that the Overall, along-channel A values have a highly transect’s range of values. The power function bridge backwater effect is less extensive than variable but convex trend, with peak values oc- fit applied to the nondimensional data allowed the bedrock effect. On the other hand, the back- curring at -1 km to 2 km (Fig. 2D). the fitted exponent, or shape factor b, to vary, water effect of tides can be expected to extend In order to further characterize the varia- and this provides a more quantitative metric for further upstream (Sassi and Hoitink, 2013) and tions in A, a linear trend line was applied to channel shape. Here, b = 1.0 corresponds to a therefore exert a greater influence on channel the 18-km-long reach that is furthest upstream rectangular channel, and b = 4.0 represents a geomorphology upstream of the -4 km mark. (Fig. 2D) and in the alluvial part of the chan- wide V-shape. Overall, b ranges from 1.16 to In summary, the long profiles from the center nel, from -34 km to -16 km (r2 = 0.74). The 3.06, with a mean of 1.83 ± 0.45. line and thalweg (Figs. 2B C) reveal a typical trend line was then extrapolated from -16 km The trend in b (Fig. 2E) from -34 km to riverbed elevation decline toward the mouth, to +4 km. This line fit illustrates how the area of -23 km declines from 1.6 to 1.1, respectively. but with upstream and downstream reaches the upstream section of alluvial channel can be This is followed by an increase to 1.9 at -9 km, that have distinctly different attributes. In the expected to change toward the downstream, and but further downstream, the values remain upstream, the bed relief is typically smaller with fluvial-dominant conditions (Yankovsky et within a range of 1.5–2.1 to 8 km. Continuing and less variable. The downstream reach has al., 2012). Between -16 km and -1 km, A values downstream, the values have no particular pat- nominal slope and kilometer-scale “plateau” plot substantially lower than this trend line. This tern with distance, and the variability increases features that stand higher than the surround- result emphasizes the aberrant lower A channel but remains within a range of 1.4–3.1. Since the ing riverbed. It is inferred that the plateaus at reach. Moreover, this shoaling reach is bounded shape factor corresponds to nondimensional 3 km to 7 km, 9 km to 11 km, and 19 km to at the upstream and downstream ends by an channel properties, it follows that the down- 22 km result from resistant bedrock outcrops. ~50% decline in A that occurs over a <2 km stream effects of bedrock (Figs. 2A and 2E) do Hence, properties of the lower Santee River distance (Fig. 2D). Together, these observations not generate overall aberrant channel shapes. channel are influenced by the presence of re- reveal a channel reach of substantial and abrupt On the other hand, bedrock produces wider and sistant bedrock, while outside of these reaches, changes in channel properties. shallower reaches that are rectangular but with a the system is alluvial, and channel properties The hydraulic radius, R, varies from 1.28 m to slight V-shape (Figs. 1, 2A, and 2E). there can be expected to be more dynamic. 4.14 m, with a mean of 2.68 ± 0.59 m, but most Overall, in going along the channel from Likewise, the long profile of the free surface values are less than 3.3 m. The pattern of R with upstream to downstream, there were consis- has a particularly large break in slope that co- distance from upstream to downstream shows an tent patterns with some metrics, and in some incides with the bedrock outcrops. This transi- initial increasing trend, followed by high vari- reaches, and high variability in others (Figs. 2B tion is shown as a lower free surface slope over ability likely associated with bedrock outcrops and 2C). Typically, depth, area, and hydraulic the bedrock reach, and this creates conditions (Fig. 2E). From -34 km to -17 km, there is an radius initially increase toward -17 km. Mean- that favor generation of a backwater effect. The increase from 1.3 m to 3.3 m, the sixth highest while, the upstream b values start with slightly backwater length is 7.5 km, and it extends from value. This is followed by consecutively declin- varying lower values from -35 km to -17 km -3.5 km to 4 km (Figs. 1 and 2C). Therefore, ing values, reaching 2.2 m at -13 km, and a re- before undergoing a step-like increase where local channel metrics upstream of -4 km likely turn to 3.1 m at -9 km. From there, the values depth, area, and hydraulic radius decline. These arise from a combination of are 2.5 m to 3.3 m to the 3 km mark, but further trends reach a minimum and then recover, but and tidal backwater effects. downstream, R values range between 1.4 m and only area continues to increase substantially to a 4.2 m and lack a clear trend (Fig. 2E). peak at -1 km. Further downstream from 3 km Channel Cross Sections Exploring the details of R downstream of the to 30 km, all variables exhibit the greatest varia- bridge, however, sheds light on the bedrock ef- tions. In particular, bed elevation and area have a Bankfull cross-section area values, A, range from fects. For instance, the lower values of 1.2 m highly variable but declining trend, while widths 119 m2 to 554 m2, with a mean of 282 ± 82.7 m2. to 1.7 m coincide exactly with the presence of remain relatively low. Mean depth, hydraulic ra- From -34 km to -16 km, there is a net increas- bedrock outcrops (Fig. 2E). Moreover, the high dius and the shape factor fluctuate the most, and ing trend from 115 m2 to 324 m2, respectively downstream values of 3.0 m to 4.1 m are coinci- their respective locally lower values are associ- (Fig. 2D). On the other hand, there is a net de- dent with abrupt river widening or deepening, or ated with bedrock outcrops. cline at 8 km to 30 km, from 385 m2 to 210 m2, both (Figs. 1, 2A, 2B, and 2C). Taken together, Field observations, and the patterns of along- respectively. For the channel reach between these observations show that when the bedrock- channel width and profile elevation support

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the view that resistant bedrock affects channel a section of channel characterized by an abrupt ated with the observed rapid tidal energy dissi- properties at three discrete locations, 3 km to near doubling of A from 1 km to 4 km upstream pation, and the high rate of dissipation may cor- 7 km, 9 km to 11 km, and 19 km to 22 km. The of the bedrock outcrops, e.g., in the alluvial part respond to the onset of a fluvial-tidal transition effects of bedrock are clearly apparent in all but of the system (Figs. 1 and 2D). They reasoned in hydrodynamic processes. Hence, Yankovsky two metrics, sinuosity and shape factor. Typi- that the sharp increase in area is likely associ- et al. (2012) identified the­downstream limit of cally, there are shallower and wider channels in the bedrock reaches. Also, the along-channel bathymetry shows local highs or plateaus that 100,000 are consistent with the response in thalweg ele- A Potomac River vation, while channel area, depth, and hydraulic 80,000 radius are locally depressed. The lower part of the Santee River is a mixed bedrock-alluvium system, and the transition from alluvium to 60,000 resistant bedrock clearly has an influence on channel metrics. When channel metrics from bedrock segments are removed, the respective 40,000 trends with distance show reduced scatter. Bed- rock also produces a measureable decrease in 20,000 free surface slope, and this gives rise to a back- water effect from -3.5 km to 4.0 km. Finally, it should be noted that two of the 10 metrics 0 highlighted in Figure 2 were directly computed 175 150 125 100 75 50 25 0 from cross-section area of flow A, those being 50,000 the hydraulic radius R, and mean depth d, and B James River therefore can be expected to covary accord-

) 40,000

ingly. The other metrics, however, are indepen- 2 dent of A. 30,000 DISCUSSION

Tide-influenced rivers typically have cross- 20,000 section area of flow, A, that increases exponen- tially toward the coast (e.g., Nichols et al., 1991; Friedrichs, 2006; Savenije, 2012; Nittrouer et Cross Section Area (m 10,000 al., 2011b). With the Santee River, however, there is an ~22-km-long discontinuous bed- 0 rock reach in which A declines downstream (Fig. 2D). Of course, A cannot decrease indefi- 100 75 50 25 0 1,000 nitely because the tidal prism must be accom- modated with greater A toward the mouth. The C Altamaha River declining trend therefore is likely an artifact of 800 bedrock influence. For instance, the trend in A for bedrock sections alone shows a relatively uniform downstream decline, and it appears 600 that this trend is translated onto the juxtaposed alluvial reaches (Fig. 2D). Hence, bedrock can 400 strongly influence alluvial channel properties in low-gradient, seemingly transport-limited coastal plain channels, similar to the Missis- 200 sippi River, with segments of -resistant but unconsolidated sediment (Nittrouer et al., 2011a). These observations highlight caveats 0 when using simplified one-dimensional process 75 50 models to help establish morphodynamic equi- Distance from Estuary Mouth (km) librium of tide-influenced river systems (e.g., Pittaluga et al., 2015). Figure 3. Along-channel variations in cross-section area of flow for three coastal plain rivers in Despite the confounding effects of bedrock the southeastern United States: (A) Potomac River, (B) James River, and (C) Altamaha River. on channel properties, Yankovsky et al. (2012) Note each system contains a reach of aberrantly low area that persists for tens of kilometers detected a geomorphic response to the fluvial- (circled), before recovering to background values, and then followed by a gradual decline tidal transition zone. In particular, they identified further upstream. A and B are taken from Friedrichs (2006), and C is from a related study.

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a fluvial-tidal transition reach as an abrupt in- through A, and at the very least, one should not (Fig. 4B), as first reported for the Santee River crease in A. rely exclusively on trends in channel width as a by Yankovsky et al. (2012). Bedrock effects The upstream limit of the transition reach proxy for trends in A. In the case of the Santee were omitted from this conceptual framework. also is associated with a sharp change in A. River, channel width is not a reliable indicator Between the tide-influenced fluvial and the For instance, upstream of -1 km, the A values of the fluvial-tidal transition channel. fluvial-influenced tidal parts of the system, there remain relatively low to -16 km (Fig. 2D). At exists an abrupt geomorphic discontinuity in -17 km, area increases by ~60%, followed by Conceptual Model A (Fig. 4A). From the upstream extent of tidal steadily declining A values in the upstream. It is oscillations (the tidal limit) to downstream, the proposed that the upstream peak in A is associ- Based on the field observations, and in con- transition reach is identified by a local high in ated with the upstream limit of the fluvial-tidal sideration of the work of Ashley and Renwick flow area, A, followed by a sustained decrease transition zone. Further, the limits of the tran- (1983), Allen (1991), Gurnell (1997), van den that may persist on the order of tens of kilo- sition reach coincide with the notable disconti- Berg et al. (2007), and Sassi and Hoitink (2013), meters. Further downstream, A abruptly recov- nuities in flow depth, hydraulic radius, channel the following augmentation to the prevailing flu- ers and attains a local peak value, followed by shape, and sinuosity (Figs. 1 and 2). Taken to- vial-tidal facies model of Dalrymple and Choi a rapid decline, perhaps related to the onset of gether, these data indicate that the fluvial-tidal (2007), Dalrymple et al. (2012), and Jablonski rapid tidal energy dissipation, as proposed by transition can be readily identified as a channel et al. (2016) is proposed. Of course, the posi- Yankovsky et al. (2012). At the upstream lo- reach of substantially lower A with abrupt up- tions of the various attributes of the fluvial-tidal cal high in A there is the limit of current re- stream and downstream boundaries. It is note- transition vary with time, and so this revised versals, or the transition from unidirectional to worthy that widths at, and on either side of the conceptual view is presented as channel prop- bidirectional flows. Further downstream, A transition reach are relatively uniform (Fig. 2A). erties that develop in response to the recurring develops the expected exponentially increasing For the Santee River, it appears that the geo- channel-forming processes (Fig. 4A). The larg- trend to accommodate the tidal prism (Figs. 2D morphic fluvial-tidal transition reach is between est departure from Dalrymple and Choi (2007) and 4A), with the salinity limit taken as being -1 km and -17 km (Figs. 1 and 2C). However, is an extensive channel reach where the fluvial proximal to the limit of fluvial influence (Odum, the bedrock backwater length, L, has an upstream and tidal energies are of comparable magnitude 1988). The channel reach with lower A is taken limit at -4 km. Therefore, the transition reach has a well-defined upstream limit, but the downstream limit is more ambiguous due to the backwater ef- fects. The downstream extent of the fluvial-­tidal Fluvial Tide-influenced Fluvial Fluvial ≈ Tidal Fluvial-influenced Tidal Tidal transition reach may be in the vicinity of -1 km, A or it may extend much further downstream. It follows that, the shoaling transition channel reach extends from -17 km to at least -4 km (Figs. 1 and 2C), i.e., at least 13 km long. Salinity Similar along-channel shoaling has been re- Current Limit ported for other southeastern U.S. coastal plain Tidal Reversal Limit Limit

rivers. The pattern of interest is a decline in A Cross Section Area of Fl ow that persists on the order of tens of kilometers, and the recovery to background values (Figs. 2D and 3). The Potomac River shows such a trend between 75 km and 100 km from the river B Fluvial Tidal mouth, and the James River at 60 km to 75 km, 100% respectively (Figs. 3A and 3B; data taken from

Friedrichs, 2006). Also, the Altamaha River has e Energy a similar pattern at 49 km to 59 km (Fig. 3C).

Hence, a 10–30 km channel reach with unchar- Relativ 0% acteristically lower cross-section area appears to Uplands Estuary Mouth be a recurring feature in some coastal plain riv- Along Channel Distance ers of the southeastern United States. It is pro- posed that this pattern of lower area is a type of Figure 4. Conceptual summary of the fluvial-tidal transition zone properties. Note the over- geomorphic discontinuity. all decline in cross-section area of flow in the upstream direction. Superimposed on this Lastly, variations in transition zone A are not trend is a trough of lower area values bound by abrupt increases. (A) The three parts of necessarily associated with similar or compa- the transition zone are the tide-influenced fluvial, fluvial≈ tidal, and fluvial-influenced tidal rable variations in channel width. Since con- zones. These three zones are bounded by the salinity limit in the downstream reach and vergence in A has such an important effect on the tidal oscillation limit in the upstream reach. The transition channel is shown as having tidal properties up river, including current di- aberrantly low cross-section area of flow in the fluvial-tidal transition channel. (B) Relative rection and magnitude (e.g., Langbein, 1963; energy along the channel highlights two main points. First, the occurrence of a long chan- Wright et al., 1973; Allen, 1991; Friedrichs and nel reach of comparable fluvial and tidal energies in the transition channel, and second, the Aubrey, 1988; Dalrymple and Choi, 2007), a abrupt (e.g., not gradual) transition from fluvial to tidal dominance. For the Santee River, more meaningful assessment of upstream flood the upstream and downstream transition boundaries are 1–2 km long, and the transition tide tidal prism convergence is best represented reach is at least 13 km.

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as the fluvial-tidal transition reach. This gen- discharges are comparable to upstream tidal with recurring and comparable upstream and eralized conceptual view highlights a geomor- discharges, and the transition reach may retain downstream, or fluvial and tidal, discharge or phic discontinuity in A, and it results from a conditions that favor intermittent and short- hydrodynamic conditions. simplification of the most salient and transfer- lived changes in ebb and flood tide dominance able features detected in a coastal plain river that may give rise to conditions favoring sedi- ACKNOWLEDGMENTS (Figs. 3 and 4A). ment accumulation and declining cross-section This work was supported by National Science In the absence of bedrock effects, coastal area. Moreover, the recurring comparable flu- Foundation EAR GLD award number 10-53299 plain riverbed shoaling can be expected to arise vial and tidal energies dictate that the limit of and National Aeronautics and Space Administration from backwater effects due to tides (Sassi and current reversals must reside within the transi- (NASA) award number EPSCOR NNX16AR02A. Kyungho Jeon acquired the field data as part of a the- Hoitink, 2013) and the river-to-ocean transition tion reach (Fig. 4). sis project, with support from Jeff Ollerhead and Matt (Fernandes et al., 2016). Santee River tidal oscil- Finally, the role of bedrock on the fluvial- Balint. Alexander Yankovsky, Miles O. Hayes, and lations in stage were <0.04 m at -35 km (Fig. 1). tidal transition cannot be overlooked, especially Allan James provided insightful commentary during Although tidal oscillations can be expected to since it seems that the presence of bedrock is the development this manuscript. Reviews by Dale Leckie and one anonymous reviewer helped improve decline asymptotically further upstream (e.g., a recognized feature of low-gradient coastal this contribution. Jay, 1991), for practical considerations, it is plain rivers (e.g., Nittrouer et al., 2011a). In the assumed that the tidal limit is at about -40 km case of the Santee River, it is proposed that the REFERENCES CITED (Fig. 1), or ~100 km from the mouth. Following bedrock outcrops contributed to the rapid dis- the approach of Parker (2004), the backwater sipation of tidal energy over a relatively short Allen, G.P., 1991, Sedimentary processes and facies in the Gironde Estuary: A recent model for macrotidal es- length L for the river-ocean transition is up to bedrock reach. If the bedrock outcrops were tuarine systems, in Smith, D.G., Reinson, E., Zaitlin, ~53 km, using a downstream channel depth of not present, the fluvial-tidal geomorphic dis- B.A., and Rahmani, R.A., eds., Clastic Tidal Sedimen- 3.23 m (where average depth is 2.41 ± 0.82 m), continuity, the tidal limit, and other tidal river tology: Canadian Society for Petroleum Geologists Memoir 16, p. 29–40. and a downstream free surface slope of 6.02 × features would have occurred further upstream. Ashley, G.M., and Renwick, W.H., 1983, Channel morphol- 10-5 (Figs. 2A and 2C). It follows that the ocean Hence, tide-influenced coastal plain river fea- ogy and processes at the riverine-estuarine transition, the Raritan River, New Jersey, in Collinson, J.D., and backwater effects may extend up to the middle tures and their respective positions along the Lewin, J., eds., Modern and Ancient Fluvial Systems: of the bedrock outcrops. These observations in- channel can be strongly influenced by bedrock International Association of Sedimentologists Special dicate that the abrupt changes in A that define conditions that promote the development of Publication 6, p. 207–218. Ashley, G.M., Renwick, W.H., and Haag, G.H., 1988, Channel the fluvial-tidal transition upstream of the bed- backwater effects, and dissipation of tidal en- form and process in bedrock and alluvial reaches of the Rar- rock outcrops most likely developed in response ergy. The same can be expected for managed or itan River, New Jersey: Geology, v. 16, p. 436–439, doi:10 to tidal backwater effects. stabilized tide-influenced channels. .1130/0091-7613(1988)016<0436:CFAPIB>2.3.CO;2. Ashworth, P.J., Best, J.L., and Parsons, D.R., eds., 2015, Another facet of the fluvial-tidal transition Fluvial-Tidal Sedimentology: New York, Elsevier, De- is that the fluvial and tidal energies of flow CONCLUSIONS velopments in Sedimentology 68, 656 p., doi:10.1016/ B978-0-444-63429-7.00002-X. frequently attain comparable magnitude. Fig- Blum, M.D., and Tornqvist, T.E., 2000, Fluvial re- ure 4B summarizes the distribution of relative The channel reach linking the fluvial- and sponses to climate and sea level change: A review flow energy, similar to the conceptual models tidal-dominant parts of the Santee River has and look forward: Sedimentology, v. 47, p. 2–48, doi:10.1046/j.1365-3091.2000.00008.x. presented by Dalrymple and Choi (2007), Dal- distinct geomorphic features that give rise to a Bokuniewicz, H., 1995, Sedimentary systems of coastal rymple et al. (2012), and Jablonski et al. (2016). fluvial-tidal geomorphic discontinuity. 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