Deciphering the relative importance of fluvial and tidal processes in the fluvial–marine transition

CHAPTER

Deciphering the relative importance of fluvial and tidal processes in the fluvial–marine transition

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R.W. Dalrymple*, C.E. Kurcinka*, B.V.J. Jablonski†, A.A. Ichaso{, D.A. Mackay} *Department of Geological Sciences and Geological Engineering, Queen’s University, Kingston, Ontario, Canada † Statoil Canada Ltd., Calgary, Alberta, Canada { Shell E&P, Houston, Texas, USA } Serinus Energy, Calgary, Alberta, Canada

1.1 INTRODUCTION The interaction of tidal and river currents is a fundamental attribute of all rivers that empty into a marine basin, provided the basin is not tideless (as is essentially the case in the Mediterranean Sea and much of the Arctic Ocean; Candela, 1991; Kowalik and Proshutinski, 1994). The length of the zone of interaction is a complex function of several variables, including the coastal-plain gradient, the tidal range at the coast, and the fluvial discharge, with depth of the river playing a secondary role. A decrease in gradient causes an increase in the distance landward of the coast that tidal action can be expected, all else being equal: for example, a halving of the slope doubles the extent of tidal penetration (Fig. 1.1). An increase in the tidal range also causes tidal action to penetrate farther inland, whereas an increase in fluvial discharge decreases the tidal penetration, again all else being equal (Dyer, 1997; Nichols and Biggs, 1985). Water depth, together with channel width, also influences the upstream penetration of the tidal wave through their influence on frictional attenuation of the tide. In many systems, the landward decrease in the cross-sectional area of the channel initially causes the tidal range (and tidal–current speeds) to increase landward (i.e., a “hypersynchronous” situation; Salomon and Allen, 1983) to a location referred to as the “tidal maximum” (Dalrymple and Choi, 2007). Beyond this, friction causes the tidal range to decrease to zero at the tidal limit (Godin, 1999). Because friction is higher in shallow water, shallow systems such as braided rivers are likely to have shorter tidal-penetration distances than rivers that are deep. The end result of these various factors is that large rivers, which generally flow over low-gradient coastal plains, tend to have longer tidal-penetration distances than Developments in Sedimentology, Volume 68, ISSN 0070-4571, http://dx.doi.org/10.1016/B978-0-444-63529-7.00002-X © 2015 Elsevier B.V. All rights reserved.

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FIGURE 1.1 Tidal penetration as a function of coastal-zone gradient. In both panels, the tidal range at the mouth is the same and the only difference is slope: (A) has approximately twice the slope of (B) and half the distance of tidal penetration. In this figure, the tidal range is shown as increasing inward from the river mouth to the “tidal maximum” (T.M.; Dalrymple and Choi, 2007), beyond which the tidal range decreases toward the tidal limit because of frictional dissipation. Systems that show such an initial increase of the tidal range are termed “hypersynchronous” (Salomon and Allen, 1983). These diagrams do not take into account the influence of differences in fluvial discharge: the effect of an increase in river flow is to displace the tidal limit in a seaward direction.

small rivers with steeper slopes. [Larger rivers also tend to have stronger tidal currents at their mouth because of the larger tidal prism (i.e., the tidal water flux past a point on each half tidal cycle) that is caused by the longer distance of tidal penetration (cf. Dalrymple, 2010).] The longest tidal-penetration distance documented is in the Amazon River, in which a tidal influence is detectable up to 800 km landward of the coast. However, even small to medium rivers have tidal-penetration distances of several tens to more than a hundred kilometers in low-gradient, coastal-plain settings (Dyer, 1997; A.A. Ichaso, unpublished data; Nichols and Biggs, 1985; Van den Berg et al., 2007). It follows, therefore, that a significant fraction of the coastal-zone deposits in ancient sedimentary basins should have accumulated in the tidal–fluvial transition zone. If we are to undertake sophisticated interpretations of these deposits, it is necessary to be able to determine where a particular deposit formed in the tidal–fluvial transition. In simplistic terms as noted by Dalrymple et al. (1992, 2012) and Dalrymple and Choi (2007), the transition zone can be thought of as displaying a gradient in the relative importance of river and tidal currents (Fig. 1.2). Unfortunately, a robust method of positioning a deposit with respect to these gradients does not currently exist. To date, our best tool for determining the relative position of deposits in the fluvial–marine transition has been the trace-fossil assemblage, which has been calibrated against salinity with reasonable precision (M. Gingras, 2007, personal communication; cf. Gingras

1.2 Process framework for the fluvial–tidal transition

FIGURE 1.2 Simplified representation of the interaction of tidal and fluvial processes in the fluvial–marine transition zone, based on Dalrymple et al. (1992) and Dalrymple and Choi (2007). This diagram portrays only the time-averaged intensity of the two processes and ignores temporal variations in their relative strengths, a factor that is central to the approach advocated in this chapter.

and MacEachern, 2012; Gingras et al., 2012a,b; MacEachern and Gingras, 2007). A comparable tool using physical sedimentary structures has not yet been developed. The objective here is, therefore, to propose an approach to determining the relative importance of fluvial and tidal currents in the formation of a deposit, which is based on an understanding of the dynamics of river floods and their interaction with tides in the transition zone. The application of the concept is then illustrated with a series of examples that we believe to show nearly the full range of possible ratios of tidal power to fluvial power, and to span the full transition zone from fluvially dominated with a weak tidal signal to tidally dominated with a barely recognizable fluvial signal. In this, we build on the pioneering studies of Jones et al. (1993), Gingras et al. (2002), and van den Berg et al. (2007), who were among the first to document the role of river-discharge variations in the formation of sand–mud alternations in occurrences of inclined heterolithic stratification (IHS; Thomas et al., 1987) and to recognize the interaction of fluvial and tidal processes in its formation. We emphasize that the approach presented here is most applicable to the heterolithic strata that are common in the fluvial–marine transition and not to the sand-dominated deposits of the channel thalweg where the temporal variations in the strength of the tidal and fluvial currents are commonly cryptic (cf. Martinius and Gowland, 2011).

1.2 PROCESS FRAMEWORK FOR THE FLUVIAL–TIDAL TRANSITION Our proposed model for tidal–fluvial sedimentation is based on the recognition that the temporal variability of fluvial and tidal processes are usually very different. Although rivers generally flow continuously, most fluvial sedimentation occurs during

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the relatively short periods of river flood because this is when the physical energy is the greatest and the sediment-transport capacity of the river is the highest. (Note: Because the word “flood” is used in two different ways in this chapter, we never use “flood” alone and always say “river flood” or “flood tide” to avoid potential confusion.) The duration of river floods is highly variable, both within a single river system and between river systems. As a general rule, however, the duration of river floods is a function of the size of the drainage basin, although other factors such as the intensity of precipitation and the infiltration capacity of the catchment also influence the shape and duration of the river-flood hydrograph. Small headwater streams with small catchments tend to have short-lived floods (i.e., commonly only a few days in duration) that represent the response to discrete precipitation events (i.e., a single storm; Fig. 1.3A). By contrast, rivers fed by large catchments are relatively insensitive to individual precipitation events because they integrate the drainage from many smaller tributaries and the river flood wave becomes stretched and attenuated during transit through the tributary network (Carter and Godfrey, 1960; Milliman and Farnsworth, 2013; Smith and Ward, 1998). Instead, large rivers respond to longer-term, generally seasonal, variations in runoff (Fig. 1.3B) as a result of, for

FIGURE 1.3 Examples of flood hydrographs for a “weather-dominated” river (A; Huai Bang Sai at B. Nong Aek, Thailand) and a “climate-dominated” river (B; Mekong River at Phnom Pemh, Cambodia). In (A), the river discharge varies erratically because the river responds directly to each precipitation event, whereas in (B) the river responds primarily to seasonal variations in precipitation as a result of the south-Asian monsoon. Note that in (B), the flood hydrograph is much smoother than that in (A), but there are still smaller peaks and valleys, especially during the rising limb, presumably because of local precipitation events or short-duration variations in the intensity of the monsoon. Such irregularities are not considered in the conceptual development presented in this chapter. Note the vast difference in the discharge scale of the two rivers. X-axis in months. Based on figure 4.10 of Smith and Ward (1998); reproduced with permission from John Wiley & Sons Ltd. Original diagram reproduced from Volker (1983, figures 1 and 2) with permission of IAHS.

1.2 Process framework for the fluvial–tidal transition

example, monsoonal variations in rainfall or changes in snow melt as a function of temperature changes. Most such rivers have a single major flood each year that lasts for several weeks or months. Small rivers tend to be “flashy,” but can also be termed “weather-dominated,” whereas larger rivers can be called “climate-dominated” because their discharge regime reflects the general climate of the drainage basin. The rivers that are primarily responsible for the construction of coastal plains are likely to be of the latter type, because medium to large rivers are the main sources of sediment, especially in low-gradient sedimentary basins, because it is the larger rivers that have their headwaters in the mountainous regions that are the primary source of sediment. (See Syvitski and Milliman (2007) and references therein for a review of sediment delivery as a function of drainage-basin size, topographic relief, and bedrock lithology. See also the recent work on estimating the size of drainage basins in coastalplain succession (Bhattacharya and Tye, 2004; Davidson and North, 2009)). Regardless of the nature of the discharge regime and the duration of an individual river flood, fluvial sediment deposition, whether on point bars or deltaic mouth bars, is episodic to at least some degree. Flow discharge and current speed commonly increase relatively quickly from their preflood levels, reaching a peak value before decreasing more slowly to low levels (commonly something approaching “base flow”; Rodda, 1969; Smith and Ward, 1998). Because of the power–function relationship between current speed and sediment transport, it is common for most, if not all, fluvial sedimentation to occur during the river flood, with relatively less or perhaps no accumulation during the intervening, low-flow (i.e., interflood) period. Given that river-flood sedimentation is almost certainly the predominant style of sedimentation in most rivers, it is surprising that so little emphasis has been placed on documenting the nature and variability of river-flood deposits. The only systematic review of the deposits of individual river floods known to the authors is that provided in figure 4.54 of Bridge (2003) (Fig. 1.4); most textbooks, by comparison, have almost no discussion of the nature of the deposits of an individual flood (e.g., Miall, 1996). The rapid rise of current speeds at the onset of the river flood (Fig. 1.3) typically leads to erosion of the bed. Deposition begins near peak discharge when flow conditions are approximately steady (Bridge, 2003). The grain size and sedimentary structures of the deposits formed at this time reflect the grain size that is available and the local combination of current speed and water depth. The upper part of the river-flood deposit accumulates during the waning stage of the flood (i.e., the falling limb of the flood hydrograph; Fig. 1.3). (Although Fig. 1.4 has been annotated to show the general, proximal–distal trends that might be expected in a fluvial system, the changes can also occur vertically on a single point bar, because of the tendency for sand grain size to decrease upward.) The interflood deposits, if any accumulate, are commonly not preserved in channel-axis locations, and successive river-flood beds are amalgamated because of erosion by the next river flood; indeed, there is likely to be the preferential removal of thin flood beds formed by smaller river floods, and preferential preservation of the thick deposits formed by the most energetic floods (Sambrook-Smith et al., 2010; see Thorne et al. (1991) for a comparable situation with storm-event beds on shorefaces). In more sheltered locations such as higher on point bars, in counter-point-bar areas (Smith et al., 2009) or the more distal parts

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FIGURE 1.4 Variation in the character of river-flood deposits as a function of grain size, which is in turn a function of proximal–distal location, position on a point bar, etc. No absolute thickness or temporal scale is implied. Based on figure 4.54 of Bridge (2003). In the tidal–fluvial transition, the muddy interflood deposits that cap each succession can contain evidence of tidal action as well as marine bioturbation because of the upstream migration of the tidal and salt-water limits. Reproduced with permission from John Wiley & Sons Ltd.

1.2 Process framework for the fluvial–tidal transition

of mouth bars, the interflood deposits are likely to be muddy, especially in the coastal plain, where IHS is commonly developed (Johnson and Dashtgard, 2014; Jones et al., 1993; Sisulak and Dashtgard, 2012; Thomas et al., 1987). It is these heterolithic deposits that are the focus of the analysis here. Compared to the episodic nature of river floods, tides operate continuously, albeit with regular variations in tidal range and peak current speed as a result of the neap–spring and other longer-period cycles (Allen, 1997; Kvale, 2006). There is potentially, therefore, a sedimentation event every half tidal cycle (i.e., approximately every 6 h (semidiurnal tides) or 12 h (diurnal tides)). Because of the asymmetry that develops in the tidal wave as it propagates into shallow water (Allen, 1997; Dalrymple, 2010), the flood tide is usually of shorter duration than the ebb tide, with correspondingly higher current speeds during the flood, leading to the widespread development of a net landward transport of both bed-load and suspended-load sediment (i.e., a flood-tide dominance; Dalrymple, 2010; Dyer, 1997; Yu et al., 2014). As noted in Section 1.1, many estuaries (sensu Dalrymple et al., 1992) and deltas are hypersynchronous (Salomon and Allen, 1983), such that the zone with the largest tidal range and maximum tidal influence are located some distance inland of the coast (Fig. 1.1). Beyond this point, tidal influence decreases and the relative importance of river currents increases when conditions are averaged over a time span of many years. Although tides theoretically operate continuously, their intensity is modulated by river discharge: just as the salt wedge is displaced seaward by increased fluvial discharge, so too is the tidal limit (Fig. 1.5; Allen et al., 1980; Kravatsova et al., 2009; Sisulak and Dashtgard, 2012; Uncles et al., 2006). This means that there is generally a reciprocal relationship at any point in the transition zone between the intensity of river currents and the strength of the tidal currents (Figs. 1.6–1.9), such that times of maximum fluvial influence are the times of minimal tidal influence, and vice versa. This relationship can cause sedimentation in many parts of the fluvial–tidal transition to alternate between fluvially dominated during river floods and tidally dominated during the interflood periods (van den Berg et al., 2007). At the very least, the relative intensity of the two processes varies on the period of the changes in fluvial discharge (Johnson and Dashtgard, 2014; Sisulak and Dashtgard, 2012), regardless of how long the river flood lasts. The variation in river discharge also changes the location and intensity of the turbidity maximum, the zone of the highest suspended-sediment concentrations in the fluvial–marine transition zone (Dyer, 1997). At times of increased river discharge, the turbidity maximum is shifted seaward and commonly it increases in overall turbidity because of the net addition of fine-grained sediment by the river (Castaing and Allen, 1981; Doxaran et al., 2009; Lesourd et al., 2003; Uncles et al., 2006). Therefore, areas seaward of the low-flow/interflood position of the turbidity maximum will experience increased turbidity during river floods, whereas areas farther landward might see higher suspended-sediment concentrations during times of low river flow (van den Berg et al., 2007). Of course, changes in river discharge will also cause the salinity gradient to migrate up-river as discharge decreases, and in a seaward direction as discharge increases (Fig. 1.5). Such changes should be reflected in the

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FIGURE 1.5 Schematic illustration of the control that variations in river discharge have on the inland penetration of salt water and tides. Diagram inspired by data from the Irrawaddy River (Kravatsova et al., 2009). The four small inset panels at the top show the locations of Figs. 1.6–1.9, positioned relative to the water-level profiles in the main diagram. In these panels, the black line is the river-discharge hydrograph and the blue line shows the peak speeds of the tidal currents. This example has a large tidal range at its mouth, thereby allowing illustration of the full spectrum of depositional conditions from fluvially dominated near the limit of tidal penetration to tidally dominated near the coast. Systems with a smaller tidal range at their mouth will not experience depositional conditions like those shown in insets Figs. 1.8 and 1.9 and conditions such as those shown in inset Figs. 1.6 or 1.7 could occur at the mouth of the system, as is seen in case studies 1 and 3.

ichnological character of the deposits, with the most diverse trace-fossil assemblages occurring in the deposits that accumulated between river floods (cf. Gingras et al., 2002), when the salinity was the highest and the current energy and sedimentation stresses were least. The temporal variations in depositional conditions over a single river flood at any given point, which are a direct result of the changes in fluvial discharge, provide the basis for recognizing the deposits of the tidal–fluvial transition, and for determining the relative strengths of river and tidal processes. We also suggest that this approach should allow deposits to be positioned, in relative terms, within the tidal–fluvial transition. The anticipated changes in depositional conditions over the duration of a single river flood of moderate duration (i.e., several weeks, which is typical of a medium-sized river; cf. Fig. 1.3B), in situations with different relative intensities

1.2 Process framework for the fluvial–tidal transition

FIGURE 1.6 Schematic representation of the temporal variation in the speed of the river current (heavy black line) and of the tidal currents (blue (gray in the print version) line) over the course of a river flood, at a site that is strongly fluvially dominated. The river flood is shown as lasting approximately 1 month, a value that is not uncommon for medium-sized rivers close to the coast (cf. Fig. 1.3B). Because the tidal wave typically becomes asymmetric in shallow water, flood–tidal currents are shown as being stronger than those of the ebb, and the flood–tidal currents are plotted above the zero line, so that their magnitude can be compared more easily with the river currents in order to evaluate the process dominance. For simplicity of representation, the neap–spring variation in tidal–current speeds is not shown, but would be present in a real example. In this example, the tidal currents are weak, and the river-flood currents dominate sedimentation except during the low-flow period (see top of diagram), when the flood–tidal currents are faster than the river current. Tidal currents go to zero during the river flood because the tidal limit is pushed seaward of this location. Conditions such as this would occur near the position of the tidal limit (cf. Fig. 1.5), regardless of its absolute distance from the shoreline. The depositional conditions shown here are interpreted to be similar to those that formed by case studies 1 and 2 (Figs. 1.15 and 1.16). Depositional conditions for case study 3 (Fig. 1.17) are believed to be intermediate between those shown here and illustrated in Fig. 1.7. See text for additional discussion.

of the river-flood and tidal currents, are shown in Figs. 1.6–1.9, and the anticipated relative positions of these figures in the fluvial–marine transition are shown in Fig. 1.5. In these schematic representations, the flood peak is shown as becoming more subdued as it passes downstream (i.e., in the progression from Fig. 1.6 to Fig. 1.9), leading to a seaward decrease in total fluvial energy as shown in Fig. 1.2. At the same time, the tidal energy increases in a seaward direction. In

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FIGURE 1.7 Schematic representation of the depositional conditions at a site that has stronger tidal currents (blue (dark gray in the print version) line) and a less pronounced river-flood peak (black line) than those shown in Fig. 1.6. In this case, tidal currents still go to zero during the peak of the river flood, but begin to be expressed during the waning stage of the river flood, and are strongly expressed during the interflood period. Sedimentation is fluvially dominated during the river flood, but tidally dominated during the interflood period (see top of diagram). (See caption of Fig. 1.6 for information about the representation of the tidal currents, and Fig. 1.5 for relative position of this case in the fluvial–marine transition.) The depositional conditions shown here are believed to be similar to those that formed in case study 4 (Fig. 1.18). See text for additional discussion.

Figs. 1.6–1.9, the tidal energy decreases during the river flood, because of the seaward displacement of the tidal limit (Fig. 1.5), although the suppression of the tidal signal during river floods is thought to be less near the river mouth (Fig. 1.9), because this area is subject to the externally imposed tide in the depositional basin. In the following section, we describe a series of case studies that illustrate these concepts and show the ways in which these variations in the relative intensity of the tidal and fluvial currents are recorded in the deposits.

1.3 SETTING OF THE CASE STUDIES USED IN THIS CHAPTER No systematic study exists of a single depositional system, modern or ancient, that can be used to test the utility of the concepts outlined above. Instead, a series of seven individual case studies have been selected from five different Mesozoic successions in which the depositional environments are well documented, to demonstrate the

1.3 Setting of the case studies used in this chapter

FIGURE 1.8 Schematic representation of the depositional conditions at a site that has stronger tidal currents (blue (dark gray in the print version) line) and a less pronounced river-flood peak (black line) than those shown in Fig. 1.7. In this case, the tidal currents are strong enough, relative to the strength of the river-flood currents, that they exist throughout the entire river flood, although they decrease slightly in strength during the river-flood peak: sedimentation is only fluvially dominated for a relatively brief period at the peak of the river flood (see top of diagram). Such a situation would occur a significant distance seaward of the tidal limit because the tidal limit did not move seaward of this location during the river-flood peak. (See caption of Fig. 1.6 for information about the representation of the tidal currents and Fig. 1.5 for relative position of this case in the fluvial–marine transition.) The depositional conditions shown here are believed to be similar to those that formed in case studies 5 and 6 (Figs. 1.19 and 1.20). See text for additional discussion.

application of the approach advocated here. The stratigraphic and environmental setting of each of the successions is presented in this section, followed in the next section by the description and interpretation of the deposits at the scale of individual beds. The case studies are presented in order from the example with the least tidal influence to the one with the greatest tidal influence (i.e., from the most fluvially dominated example to the most tidally dominated case). Note that it is not possible in the context of this overview to document fully the environmental interpretations of each succession; interested readers are encouraged to examine the work cited in each subsection.

1.3.1 LAJAS FORMATION, NEUQUE´N BASIN, ARGENTINA The Lajas Formation (Bajocian to Bathonian, Middle Jurassic) in the Neuque´n Basin, west-central Argentina (Fig. 1.10A), comprises the coastal, deltaic deposits in a progradational basin-fill succession (Howell et al., 2005; Legarreta and Uliana, 1996;

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FIGURE 1.9 Schematic representation of the depositional conditions at a site that has stronger tidal currents (blue (dark gray in the print version) line) and a less pronounced river-flood peak (black line) than those shown in Fig. 1.8. In this case, the tidal currents exceed the river currents through the entire river flood, such that sedimentation is tidally dominated at all times (see top of diagram). This would occur if the river flood was very subdued or the strength of the tidal currents was high. (See caption of Fig. 1.6 for information about the representation of the tidal currents and Fig. 1.5 for relative position of this case in the fluvial–marine transition.) The depositional conditions shown here are believed to be similar to those that formed case study 7 (Figs. 1.20 and 1.21). See text for additional discussion.

McIlroy et al., 1999, 2005). The succession begins with the deep-water to slope mudstones and turbidities of the Los Molles Formation, which are gradationally overlain by the Lajas Formation deltaic sediments that are in turn gradationally overlain by the fluvial deposits of the Challaco Formation (Fig. 1.10B). The entire Lajas Formation, which is 500–800 m thick, is heterolithic at a wide range of scales. At the largest scale, the succession consists of a vertical alternation of prodeltaic mudstones and mouth bar and associated distributary-channel sandstones, especially in the lower half of the Lajas, which is the portion from which the examples discussed here come (Fig. 1.10C). Coastal-plain mudstones of various types, including well-drained paleosols and poorly drained floodplain and interdistributary-bay deposits, are present at various levels in the middle and upper part of the Lajas. Amalgamated fluvialchannel sandstones are present at various levels, but become more abundant in the more proximal upper part of the Lajas (not shown in Fig. 1.10C). Previous work on the Lajas interpreted it as having formed in a tide-dominated deltaic environment (McIlroy et al., 1999, 2005). Recent, more detailed sedimentological

1.3 Setting of the case studies used in this chapter

FIGURE 1.10 Location map (A), general stratigraphic setting (B), and detailed stratigraphic succession of the lower Lajas Formation (Jurassic) in the Neuque´n Basin, west-central Argentina. The red (black in the print version) stars in (C) give the locations of the examples used in case studies 1 and 3. (A) Modified after Howell et al. (2005), (B) modified from Howell et al. (2005) and McIlroy et al. (2005), and (C) from Kurcinka (2014).

investigations of the lower and middle Lajas (Gugliotta et al., 2015a,b; Kurcinka, 2014) have determined, however, that the degree of tidal influence is not as great as previously believed. The lowest Lajas, below the lowest sequence boundary (Fig. 1.10C) is interpreted as consisting of river-dominated mouth bars that are separated by fine-grained prodeltaic deposits. Above the sequence boundary, the increased abundance of cyclic fine-grained drapes in the mouth-bar deposits indicates that the

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amount of tidal influence is greater than in the section below the sequence boundary, but the deltaic deposits are nevertheless interpreted as river dominated.

1.3.2 MCMURRAY FORMATION, NORTHERN ALBERTA The McMurray Formation (Neocomian-Aptian, Early Cretaceous) is widely known for its enormous heavy-oil reserves. Sedimentation occurred in a series of north– south-oriented valleys, cut by rivers of continental scale (Blum and Pecha, 2014; Mossop and Flach, 1983) that flowed northward along the eastern side of the Rocky Mountain foreland basin. The outcrop from which the example used here comes occurs along the Steepbank River, which is a tributary that joins the Athabasca River 30 km north of Fort McMurray, Alberta (Fig. 1.11A and B). The sampled succession, which comprises the informally named “middle McMurray” in that area, consists of a single, upward-fining channel-point-bar succession that is 30–35 m thick (Fig. 1.11C; Jablonski, 2012; Mossop and Flach, 1983; Musial et al., 2011). The lower 5–10 m of the succession is composed of cross-bedded sandstone that accumulated in a channel thalweg to lower point-bar setting. These sandstones are overlain by and interfinger with a 25–30 m thick succession of interbedded sandstones and mudstones that forms a single, uninterrupted set of IHS, with dips of 5°–10°, that accumulated on an actively migrating point bar (cf. Jablonski, 2012; Jablonski and Dalrymple, 2015; Smith et al., 2009). The presence of a low-diversity, brackishwater trace-fossil assemblage in the IHS has led to the interpretation of these deposits as “estuarine” in character (Gingras et al., 2012a,b; Langenberg et al., 2002; Lettley et al., 2005; Musial et al., 2011; Ranger and Pemberton, 1992; Ranger et al., 2008), but it is not known whether the succession is regressive (i.e., part of a delta) or transgressive (i.e., part of an estuary, sensu Dalrymple et al., 1992).

1.3.3 NESLEN FORMATION, BOOK CLIFFS, UTAH The Neslen Formation (Late Campanian, Late Cretaceous) crops out in the Western part of the Book Cliffs of Utah (Fig. 1.12A and B). It overlies the better-known Sego Sandstone (Fig. 1.12C), which has been interpreted as a series of tide-dominated estuaries and/or deltas, that may or may not sit within incised valleys (Aschoff and Steel, 2011; Kirschbaum and Hettinger, 2004; Van Wagoner, 1991; Willis and Gabel, 2003). The Neslen Formation, which is approximately 30–50 m thick, is much more mudstone-rich than the Sego, and is dominated by coastal-plain deposits (Willis, 2000) that are interpreted to consist of tidal flats, marshes, and carbonaceous floodplain mudstones that represent waterlogged paleosols. Sandstone intervals are interpreted by Willis (2000) as small tidal creeks that drained across the tidal flats. More recent work suggests that at least some of the floodplain mudstones accumulated in interdistributary-bay environments, whereas at least some of the sandstones accumulated in channels with a significant fluvial influence. Some of the sandstone bodies are also interpreted as delta mouth bars (Olariu et al., 2015).

1.3 Setting of the case studies used in this chapter

FIGURE 1.11 Location maps (A and B) and schematic stratigraphic succession (C) for the McMurray Formation in the area north of Fort McMurray, Alberta. The entire McMurray Formation in the study area is ca. 70–75 m thick. The example used in this study comes from the inclined heterolithic stratification (IHS) portion of the “middle” McMurray, which is ca. 25 m thick in total; the red star at the edge of the section (C) gives the approximate level from which the example comes. (C) After Mossop and Flach (1983).

1.3.4 TILJE FORMATION, OFFSHORE NORWAY The Tilje Formation (Pliensbachian, Early Jurassic) occurs in the subsurface of the Halten Terrace, which lies 100–200 km northwest (seaward) of Trondheim, Norway, beneath the Norwegian continental shelf (Fig. 1.13A and B). Deposition occurred in an overall transgressive setting in which accommodation was generated by crustal stretching and thermal relaxation associated with extension during the early stages of the opening of the North Atlantic Ocean (Dore´, 1991). The Tilje is underlain by the

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FIGURE 1.12 General (A) and detailed (B) location maps of the study area in the Neslen Formation, central Utah. The study area is located in the Book Cliffs, approximately 30 km east of Green River. (The yellow line traces the general trend of the Book Cliffs). (C) Stratigraphic position of the Neslen Formation, which overlies the well-known Sego Sandstone. The red star indicates the approximate stratigraphic position of case studies 4 and 5. (B) Map data courtesy of Google Earth#.

˚ re Formation and is overlain by the shelf mudstones of alluvial-plain deposits of the A the Ror Formation (Fig. 1.13C). The Tilje Formation itself is 150–200 m thick, and, like the Lajas Formation (see above), consists of tabular units that are composed of alternating prodeltaic mudstones and deltaic sandstones and tidal–fluvial-channel deposits (Ichaso and Dalrymple, 2014; Ichaso Demianiuk, 2012; Martinius et al., 2001). The previous sedimentological interpretation indicated that the Tilje accumulated in a succession of tide-dominated estuaries and deltas (Martinius et al., 2001), but more recent work (Ichaso and Dalrymple, 2014; Ichaso Demianiuk, 2012) proposes a more mixed-energy coastal setting, intermediate between the river- and tidedominated end members. Sediment input was generally from the north and northeast, and the succession accumulated in a southward-opening embayment, in which tidal action was accentuated, whereas the intensity of wave action was reduced by the sheltered setting. Fluid-mud deposits are abundant in the fluvial–tidal mouth bars and tidally influenced distributary channels (Ichaso and Dalrymple, 2009).

1.3 Setting of the case studies used in this chapter

FIGURE 1.13 General (A) and local (B) location maps of the study area in the Tilje Formation, central Norwegian continental shelf. The core from which the illustrated deposit (Fig. 1.20) comes was taken in the Smørbukk Field (B). (C) Generalized stratigraphic section for the Tilje Formation in the Smørbukk Field, showing the progression of depositional environments. “P.Z.” refers to the “production zones” used internally by the field operator (Statoil). The red (black in the print version) star in (C) shows the stratigraphic level from which the studied example comes. (C) From Ichaso Demianiuk (2012).

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1.3.5 BLUESKY FORMATION, PEACE RIVER AREA, ALBERTA The Bluesky Formation (Aptian-Albian, Early Cretaceous; 0–30 m thick) of westcentral Alberta is broadly coeval with the McMurray Formation (see above) and forms the middle part of the initial transgression of the Western Canadian foreland basin in the Early Cretaceous (Cant and Abrahamson, 1996). In areas where the subCretaceous unconformity is topographically low, the Bluesky is underlain by fluvially dominated coastal-plain deposits of the Gething Formation (Fig. 1.14; Cant and Abrahamson, 1996; Leckie and Smith, 1992; Mackay, 2014), but in areas where the sub-Cretaceous unconformity is high, the Bluesky rests directly on Mississippian

FIGURE 1.14 Location maps (A and B) and stratigraphy (C) of the Bluesky example used herein, which comes from the Peace River area, north-central Alberta, Canada. The succession there is temporally equivalent with the better-known McMurray Formation in the Fort McMurray area. The example used here comes from the lower part of the Bluesky (red (black in the print version) star in (C)), a heterolithic unit that is interpreted as being of regressive, deltaic origin (Mackay, 2014).

1.4 Description and interpretation of the case studies

carbonates. Everywhere, the Bluesky is overlain by the marine mudstones of the Wilrich Member of the Spirit River Formation. The first detailed sedimentological and ichnological study of the Bluesky recognized its brackish-water character and interpreted it as a wave-dominated estuarine succession (Hubbard et al., 1999, 2002). A recent re-evaluation of the Bluesky instead interprets it as having formed in a tide-dominated setting (Mackay, 2014; Mackay and Dalrymple, 2011), based on the pervasive presence of tidal sedimentary structures including cyclic tidal rhythmites, an abundance of thick (>2 cm) fluid-mud layers, a predominance of crossbedded sandstone, and an extremely low level of bioturbation (bioturbation index—BI 0–1), and by the absence of wave-generated structures. Furthermore, the clear tripartite facies organization that characterizes wave-dominated estuaries (Dalrymple et al., 1992) is not evident. Mackay (2014) has subdivided the Bluesky into lower and upper parts (Fig. 1.14C). The lower portion, from which the example described below comes, is sandstone dominated, but it is pervasively heterolithic with abundant thick mudstone layers that are interpreted as fluid-mud deposits (Mackay and Dalrymple, 2011); it is interpreted to represent a deltaic deposit because of its seaward transition into upward-coarsening delta mouth-bar successions. By contrast, the upper part of the Bluesky, which is composed almost entirely of sandstone, passes seaward into subtly upward-coarsening successions consisting of pervasively cross-bedded sandstone that are believed to represent tidal sand-ridge deposits. The overall back-stepping nature of the upper Bluesky indicates that it accumulated in a tide-dominated estuarine environment that is transitional into the overlying marine mudstones of the Wilrich Member. Both units contain abundant erosively based channel-bar successions that fine upward.

1.4 DESCRIPTION AND INTERPRETATION OF THE CASE STUDIES 1.4.1 CASE STUDY 1: LOWER LAJAS FORMATION

The lowermost Lajas Formation in the Los Molles area consists of three progradational episodes that are represented by upward-coarsening sandstone tongues, each ca. 10 m thick, composed of mouth-bar deposits, separated by finer grained, intensely bioturbated prodeltaic deposits (Fig. 1.10C; Kurcinka, 2014). Progradational clinoforms are clearly visible in outcrop (Fig. 1.15A), with dips that reach 16° (average
10°). Small channels (10–100 m wide; 1.5–4 m thick) that cut into the top of the mouth-bar deposits are interpreted to be terminal distributary channels. Internally, the clinoform packages are composed of sandstone beds that range in thickness from a few centimeters to 50 cm, with typical values in the range of 10–20 cm, that are separated by thinner intervals of finer-grained sediment (Fig. 1.15B). In proximal, up-dip locations, the bases of the beds is erosive, but the amount of erosion decreases distally, such that the finer interbeds are more prominent in distal locations. Many of the sandstone beds appear to be structureless

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FIGURE 1.15 Case study 1, lowermost Lajas Formation (Fig. 1.10). (A) Clinoform package in the lowermost part of the Lajas Formation (Fig. 1.10C), showing the seaward-dipping stratification that characterizes these river-dominated mouth-bar successions. Flat-lying, finer-grained prodeltaic deposits are locally exposed at the base of the cliff. Person for scale circled. The flood–interflood bedding shown in (B) comes from a laterally adjacent exposure within this sandstone tongue. (B) Typical example of the bedding within the sandstone tongues of the lowermost Lajas. Each sandstone bed has an erosive base. This bed contains faintly visible dune cross-stratification dipping to the left (i.e., in a seaward direction), which implies that deposition occurred by means of river-generated currents, presumably during a river flood. In more distal locations such as is shown here, the sandstone beds are separated by finer grained, intensely bioturbated material (recessive) that is interpreted as the sediment that accumulated during periods between river floods. See text for additional discussion.

internally; some of this may be due to the character of outcrop weathering and the uniform nature of the grain size (lower medium to upper fine-grained sand), but most is attributed to rapid deposition by density flows (Kurcinka, 2014). Locally, dunescale cross-stratification is evident, with paleocurrents uniformly in the seaward (i.e., northwesterly) direction. Current-ripple bedding, if present, cannot be seen due to the lack of textural differentiation and the poor quality of the outcrop exposure. Bioturbation is uncommon to rare within these beds, and the majority of the burrows in the sandstone beds extend down from the overlying, finer-grained

1.4 Description and interpretation of the case studies

deposits. These finer-grained deposits, which consist of fine to very-fine sand and silt, are, by contrast, intensely bioturbated (Fig. 1.15B), with bioturbation indices of 5–6 (cf. MacEachern et al., 2010, figure 3). The diversity of burrow types is not high (Rosselia, Rhizocorellium, Palaeophycus, Planolites, and Skolithos), but the ichnofossil assemblage is among the most diverse in the succession and includes fully marine trace fossils. Any primary sedimentary structures have been obliterated and the physical processes responsible for sedimentation can no longer be deduced. Sedimentation in this case study is clearly episodic, with periods of rapid sedimentation under energetic conditions (i.e., the physically structured sandstone beds) alternating with periods of low energy when the bioturbated interbeds accumulated. The seaward-directed paleocurrents in the sandstone beds indicate that they were emplaced by river currents, whereas the moderately diverse ichnofossil assemblages in the intervening finer-grained deposits indicate that sedimentation rates had decreased and salinity levels had increased. Therefore, the sandstone beds are interpreted to represent river floods, whereas the finer-grained interbeds accumulated in the periods between river floods. The extent to which tidal currents may have been active during the interflood periods cannot be determined because of the intensity of the bioturbation, but tidal currents cannot have been very strong because the presence of an actively shifting substrate, which would have existed if the tidal currents had been capable of significant sediment transport, would have inhibited bioturbation. On the other hand, the nature of the bioturbation indicates that the salinity during the times of low river flow was at least moderate. Given that tidal action extends farther landward than salt-water intrusion, tidal currents must have existed, but their intensity was too weak to have left any preserved record. Sedimentation conditions were, therefore, similar to those shown in Fig. 1.6, with a strong overall river dominance, with the likelihood of weak tidal action during interflood periods. In Fig. 1.5, this set of conditions, as shown by the positioning of the Fig. 1.6 inset, lies close to the tidal limit, in an area that is fluvially dominated, possibly with water becoming fresh during river floods. This is at odds, however, with the fact that the Lajas example described here experienced brackish-water conditions during low river flow, as shown by the trace-fossil assemblage in the fine-grained interbeds (Fig. 1.15B). Reconciliation of this discrepancy requires that the low-flow discharge of the river feeding the mouth bars in the lower Lajas was lower than that implicit in Fig. 1.5, thereby allowing salt water to penetrate farther inland and to extend closer to the low-flow tidal limit than is shown in Fig. 1.5.

1.4.2 CASE STUDY 2: MCMURRAY FORMATION The IHS deposits that comprise a large part of point-bar sediments in the middle McMurray Formation (Fig. 1.11C) consist of interbedded fine- to medium-grained sandstones and finer beds consisting of fine to very fine sandstone and siltstone (Jablonski, 2012; Jablonski and Dalrymple, 2015). The fundamental building block of the IHS consists of an erosively based sandstone bed (10–30 cm thick) that passes upward gradationally into a muddier interval that is 0.5–10 cm thick (Fig. 1.16). The

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FIGURE 1.16 Case study 2, McMurray Formation (Fig. 1.11). (A) General view of bedding in the sandy IHS of the middle McMurray Formation, showing the alternation of unbiotubated, sandy river-flood deposits that contain climbing, current-ripple cross-lamination, with intensely bioturbated intervals that are interpreted as interflood deposits that formed when river flow was weak and the salinity higher. In this example, the mud that commonly characterizes interflood deposition has been eroded by the subsequent flood (erosive bed bases highlighted by red lines), leaving only the lower portions of the burrows that penetrate down into the riverflood sand. It is assumed that the thickness of the river-flood sandstone can be used as a proxy for the magnitude (peak speed and/or duration) of the river flood (Jablonski, 2012; Jablonski and Dalrymple, 2015). (B) Close-up of two muddy interflood deposits that are less thoroughly burrowed than many examples, showing the thinly laminated nature of these deposits. In the best exposed examples, regular thickening and thinning of the laminae suggest that the lamination records weak tidal action. See text for additional discussion.

1.4 Description and interpretation of the case studies

sandstone beds locally contain dune cross-stratification, which is much more abundant in the lower part of the point-bar succession, but more commonly contain cosets of ripple cross-lamination (Fig. 1.16A) that climbs in some places. In the outcrops studied, all paleocurrent indicators are unidirectional in the downstream (northerly) direction. Cyclically organized mud/silt drapes are uncommon in the sandstone beds, but groupings of silt drapes do occur within some of the rippled intervals, suggesting regular retardation of the river flow. (Reversed paleocurrents and rhythmically spaced reactivation surfaces are not present in the cross-bedded sandstones that gradationally underlie the IHS deposits.) Bioturbation is generally absent from the sandstone beds, except for “top-down” penetration of the sand beds by organisms that colonized the intervening muddy deposits. Relatively uncommon burrows within the sandstone beds suggest that the depositional conditions during sand-bed deposition were infrequently able to sustain a benthic community; such within-sandstonebed bioturbation is most abundant in the thinnest of the sandstone beds. The intervening finer-grained beds are more intensely bioturbated than the sandstone beds, but BI values are generally only 2–4 (i.e., moderately bioturbated; MacEachern et al., 2010), with only some examples having BI values of 5–6 (intensely bioturbated). The trace-fossil assemblage has a very low diversity, consisting almost entirely of Cylindrichnus and Planolites, with only minor Gyrolithes. In those examples, where primary structures are still recognizable, the mudstone intervals are laminated at a millimeter scale (Fig. 1.16B). These laminae commonly show cyclic variations in thickness that are interpreted to be tidal rhythmites (cf. Coughenour et al., 2009; Kvale, 2012, and references therein) and clear double mud drapes are present locally (Jablonski, 2012; Jablonski and Dalrymple, 2015). The sandstone beds with their unidirectional paleocurrents are interpreted to be the product of river floods (Jablonski, 2012; Jablonski and Dalrymple, 2015), with rapid sedimentation being indicated by the presence of climbing-ripple lamination. It is believed that the thickness of the sandstone beds correlates with the magnitude of the river flood, thinner beds representing smaller river floods, and thicker beds representing larger floods (i.e., either higher current speeds and/or longer duration). The absence of cyclically organized silt drapes within most of the sandstone beds implies that tidal influence was not present during most river floods, but the rare presence of cyclic drapes suggests that tidal currents were present but perhaps only during brief intervals such as when spring tides (with their stronger currents; see also case study 4) coincided with lulls in the river-flood currents (cf. the discharge oscillations during the rising limb of Fig. 1.3B). The absence of reversed ripples suggests that the tides were only able to retard the river flow but not reverse it (cf. Martinius and Gowland, 2011). The water apparently became fully or nearly fresh during these river floods, thereby excluding the activity of marine organisms, although rapid sediment accumulation probably contributed to the low bioturbation indices. The occurrence of a small amount of bioturbation during the deposition of some sandstone beds, and especially the thinner ones that are attributed to smaller river floods (Fig. 1.16A), suggests that the weaker river-flood currents were not able to push the salinity-penetration limit seaward of the site of deposition. During the deposition

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CHAPTER 1 Physical record of fluvial–tidal interaction

of the muddy, interflood intervals, the water was at least somewhat brackish, but the salinity was presumably low, as indicated by the very low diversity of the trace-fossil assemblage. Tidal currents appear to have been present, but were not very strong, as indicated by the silty nature of the interflood deposits and the lack of currentgenerated ripples within them. Relative to the spectrum of depositional conditions shown in Figs. 1.6–1.9, the conditions of sedimentation appear to be most similar to those illustrated in Fig. 1.6, with tidal activity being expressed mainly during the times between river floods. Some of the thinner river-flood beds might have been formed under conditions intermediate between those in Figs. 1.6 and 1.7, with a low level of tidal activity occurring at times when the river-flood currents were weaker. Where this example lies within the fluvial–tidal transition relative to the deposits discussed in case study 1 is difficult to determine with certainty: there is more obvious evidence of tidal action in the McMurray (compare Figs. 1.15B and 1.16B), suggesting a more distal setting, but the more diverse trace-fossil assemblage in the lower Lajas suggests a higher salinity during the interflood periods, which in turn suggests that it formed farther seaward than the McMurray example. Such contradictory evidence can be explained by noting that the difference in fluvial discharge between flood peaks and low-flow periods can be very different between systems. Systems with larger river-discharge fluctuations (i.e., extreme flood peaks and/or very low base-flow discharge) will display a larger difference in salinity between river-flood and interflood times than will systems with a smaller difference in fluvial discharge between flood peak and base flow. Perhaps the Lajas, which was undoubtedly formed by smaller rivers than the McMurray, had larger discharge and salinity fluctuations (i.e., it was somewhat more flashy than the McMurray). Furthermore, the strength of the tidal signal is dependent on the tidal range and tidal prism. Thus, case study 1 from the Lajas might have formed in a setting with an overall small tidal range and/or small tidal prism, such that the strength of the tides was inherently lower than it was in the large river in which the McMurray accumulated. Such considerations indicate that the deposits formed within the fluvial–tidal transition are inherently diverse in their character because of the interaction of so many variables.

1.4.3 CASE STUDY 3: MIDDLE LAJAS FORMATION Like the lower Lajas, the middle Lajas consists of upward-coarsening successions that are of the order of 10 m thick (Figs. 1.10C and 1.17A). They begin with finegrained (silty to muddy) bioturbated deposits (BI 3–6) that are interpreted as prodeltaic in origin and pass upward into sandstone-dominated deposits that are interpreted to represent mouth bars. Small, terminal distributary channels are also present at the top of some successions. IHS caused by mouth-bar progradation is also visible, but is lower angle (i.e., <7°) than in the lower Lajas. In general, these mouth-bar deposits are more heterolithic in nature than those in the lower Lajas. Sedimentation within the mouth-bar deposits of the middle Lajas is clearly episodic and, like the two previous case studies, consists of sandstone beds that alternate with muddier intervals (Fig. 1.17B). As in the lower Lajas, the sandstone beds are

1.4 Description and interpretation of the case studies

FIGURE 1.17 Case study 3, middle Lajas Formation (Fig. 1.10). (A) Typical expression of the tideinfluenced, river-dominated mouth bars of the middle Lajas Formation, overlying muddy prodeltaic deposits in a gradational upward-coarsening succession. (B) Close-up view of a river-flood deposit, overlain by the sediment that is interpreted to have accumulated during the interflood period. The lamination in the interflood deposit is interpreted as tidal in origin because of the local presence of bipolar ripple cross-lamination. See text for additional discussion. Both images courtesy of Macello Gugliotta; see also Gugliotta et al. (2015b).

erosively based, typically consist of upper fine- to medium-grained sand, and commonly do not show many sedimentary structures, although some contain dune-scale cross-stratification, with local development of current ripples. Paleocurrents are overwhelmingly to the north–northwest (i.e., in a seaward direction), but rare landward-direction ripples are present in some of the sandstone beds. Mud and/or organic-debris drapes are rare within the sandstone beds but are present locally. As with the previous case studies, bioturbation is rare in the sandstone beds (BI 0–1).

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The muddy interbeds are more intensely bioturbated but have only moderate levels of bioturbation (BI 2–4). The ichnogenera present have a low diversity and are limited to Planolites and Ophiomorpha. Physical lamination is moderately to well developed and consists of interlaminated silt and fine- to very fine-grained sandstone; clay layers are rare but terrestrial organic detritus is moderately abundant. The lamination in the fine-grained layers locally shows regular thickening and thinning trends, suggesting that they are tidal rhythmites. Bidirectional current ripples are also present in some layers. As in the case studies discussed above, the sandstone layers are interpreted to be river-flood deposits because the grain size in these layers is coarser-grained than in the intervening deposits and the paleocurrents are predominantly seaward-directed. The low level of bioturbation is consistent with rapid sedimentation in a setting with low salinity. By contrast, the finer-grained interbeds indicate slower sedimentation under lower energy levels, and they are, therefore, interpreted as the interflood deposits. Evidence of tides, including rhythmic layering and bidirectional current ripples, is more pervasive in these interbeds than in either of the previous examples, indicating that the tidal currents were stronger than in the previous cases. Evidence of tidal–current action is also present in the river-flood deposits in the form of rare landward-directed current ripples and muddy drapes. As a result, the deposits in the middle Lajas are believed to represent conditions between those shown in Figs. 1.6 and 1.7, with some minor record of tidal–current action (perhaps forming during spring tides) within the deposits of the river floods. In general terms, the mouthbar deposits in which these beds occur are interpreted to be tidally influenced but river dominated (Fig. 1.8C; Gugliotta et al., 2015b; Kurcinka, 2014).

1.4.4 CASE STUDY 4: MIDDLE NESLEN FORMATION The middle part of the Neslen Formation contains several upward-fining sandstonedominated channel successions that are 5–8 m thick (Olariu et al., 2015). Inclined stratification is clearly evident in these sandstone bodies (Fig. 1.18A) and consists of centimeter- to decimeter-thick sandstone beds of medium to fine sand that alternate with several centimeters of finer-grained (silty) deposits that can be discontinuous because of erosion at the base of the subsequent sandstone bed (Fig. 1.18B). The sandstone beds can be structureless or can contain dune cross-stratification or ripple cross-lamination. Most of the paleocurrent indicators have a seaward-directed orientation but ripples with the opposite direction (i.e., landward) are present locally in the sandstone beds. In relatively uncommon instances, tidal rhythmites (cf. Coughenour et al., 2009; Kvale, 2012, and references therein) are present in the upper part of some sandstone beds (Fig. 1.18B). The first laminae within the rhythmite succession are thick and the laminae then become progressively thinner, before thickening again, with a final thinning trend into the finer-grained deposits that occur between the sandstone beds. Bioturbation is absent from the sandstone beds but it is present in the finer-grained interbeds. Bioturbation intensity in these interbeds is never high, however, with BI values in the range of 0–3. The only recognized trace fossil in

1.4 Description and interpretation of the case studies

FIGURE 1.18 Case study 4, middle Neslen Formation (Fig. 1.12). (A) General photo of outcrop showing laterally migrating tidal–fluvial channel. Red star shows approximate location of the outcrop photo in (B). (B) River-flood deposit consisting of an erosively based sand bed. Tidal rhythmites are absent from the lower part of this bed, but are present in the upper part of the bed (note the cyclic variation in lamina thickness) that is interpreted to represent the waning stage of the river flood. At the left edge of the photo, a remnant of the finer-grained interflood deposit has escaped erosion by the next river flood. See text for additional discussion.

the sediment is Paleophycus, with Teredolites in the woody clasts that occur near channel bases (Olariu et al., 2015). As in the preceding case studies, the sandstone beds are interpreted to be riverflood deposits because of their erosive bases, relatively coarse grain size, highenergy sedimentary structures, and predominantly seaward-directed paleocurrents, whereas the silty interbeds are interpreted to represent the interflood periods, when energy levels were less, sedimentation was slower, and the salinity increased to allow the higher degree of bioturbation. The presence of landward-directed current ripples within the river-flood beds implies that tidal currents were capable of reversing the river-flood currents, but the tidal rhythmites shown in Fig. 1.18B provide unequivocal evidence of the ability to the tidal currents to influence sedimentation, even

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during times when the river was in flood. The presence of the tidal rhythmites only in the upper part of the bed suggests that the tidal currents were not felt at this location during the peak of the river flood, and only became able to influence sedimentation during the waning stage, as river currents were becoming weaker. (That the tidal rhythmites accumulated during the river flood rather than after the flood is indicated by the fact that the tidal rhythmites require rapid sedimentation: the 10 cm of sediment within the rhythmite interval accumulated in only about 2 weeks, which requires a more rapid influx of sediment than appears to have been the case during the interflood periods.) It is significant that the first tidal laminae are thick and, therefore, represent spring–tide deposition: the strongest flood–tidal currents will occur at spring tide and these will be the ones that are able to counteract the slowly decreasing river-flood currents first. This pattern of tidal–fluvial interaction is identical to the situation shown in Fig. 1.7 and is interpreted as being relatively more distal, and certainly more strongly tidally influenced, than any of the preceding case studies.

1.4.5 CASE STUDY 5: MIDDLE NESLEN FORMATION At approximately the same level as the tidal–fluvial channels described above from the Neslen Formation, there occurs a 5 m thick, upward-coarsening sandstone body (Fig. 1.19A) that starts with bioturbated muddy deposits and passes upward gradually into IHS that becomes progressively more sandstone-dominated upward. Paleocurrent indicators are predominantly directed down the dip of the IHS, indicating that the body formed by forward accretion rather than lateral accretion. Consequently, this sandstone body is interpreted as a delta mouth bar (Olariu et al., 2015). The coarser-grained beds within the IHS are most commonly 20–30 cm thick, but reach up to 1 m in thickness. They consist of fine to very fine sand and are crossbedded in the upper part of the clinoform package, but are ripple cross-laminated in the lower part where the example described here occurs (Fig. 1.19). Unlike the previous examples, these sandstone beds contain prominent siltstone partings that drape over the formsets of the current ripples that cap each of the sandstone layers that comprise the larger sandstone bed (Fig. 1.19B). These thinner sandstone layers themselves show weakly developed cyclic variations in layer thickness, thick–thin couplets are present locally, and current ripples oriented both down and up the clinoform, although down-slope (i.e., seaward directed) orientations are predominant. The sandy beds that contain these layers alternate with silty intervals that are a few centimeters to a decimeter thick. These fine-grained beds are locally rippled, but more commonly contain no internal structures except for irregular lamination. Bioturbation is absent from the sandstone beds, and uncommon but present in the silty intervals (BI 1–2), with Paleophycus as the only identified trace fossil (Olariu et al., 2015). As for the preceding case studies, the sandstone beds are interpreted to represent river-flood deposits because of their coarser grain size and the predominant seaward paleocurrent direction, whereas the more bioturbated silty interbeds are interpreted to be the interflood deposits.

1.4 Description and interpretation of the case studies

FIGURE 1.19 Case study 5, middle Neslen Formation (Fig. 1.12). (A) General photo of outcrop showing progradational delta mouth bar. Red star shows approximate location of the outcrop photo in (B). (B) River-flood deposit consisting of an erosively based sand bed, sandwiched between finer-grained interflood deposits. The river-flood deposit contains numerous silty stringers that separate laterally continuous sand beds that display current-ripple formsets on their top surface; most of the ripples migrated to the left, in the down-slope, offshore direction. Because a cyclic variation in the thickness of the sandstone layers is weakly developed, each layer is interpreted to represent a single tidal cycle, with deposition occurring during the dominant (i.e., river current plus ebb–tidal current) half of the tide. The subordinate (flood) tide was generally not capable of depositing sand, although thin sandstone stringers are present in the center of some siltstone layers (one example is indicated by the arrow), generating a thick–thin alternation. The presence of such tidal layering throughout the entire river-flood bed, including during the lower portion of the bed that is interpreted to represent the peak of the river flood, suggests that the tidal signal is stronger than in any of the preceding cases. See text for additional discussion.

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The river-flood deposits in this example differ from those in the preceding case studies because of the presence of continuous mudstone drapes within the sandstone bed, the ordered thickness variations shown by the sandstone layers that comprise the river-flood deposit, and the local presence of bidirectional ripples (Fig. 1.19B). The drapes imply a cessation of river flow that allowed the suspended sediment to settle to the bed. The sandstone layers show a consistent seaward paleocurrent direction, but there are hints of reversed ripples as millimeter-thick sandstone lenses within some of the drapes, causing the thick–thin alternation described above. All of these features strongly suggest that tidal currents were capable of stopping and locally reversing the river-flood currents. It follows, therefore, that the tidal currents were stronger, relative to the river-flood currents, than in any of the preceding examples, because the tidal signal is present throughout the entire river-flood deposit. As a result, this example (Fig. 1.19B) is thought to represent depositional conditions similar to those illustrated in Fig. 1.8. In case study 1, tidal action presumably did not extend very far up the river, even during times of low river flow, because the evidence of tidal action in the mouth bars was negligible. By comparison, tidal action must have extended relatively farther inland in the example described here, because of the stronger tidal signal recorded in these mouth-bar deposits.

1.4.6 CASE STUDY 6: TILJE FORMATION The sediments described in this case study are representative of the bedding style that characterizes mouth-bar successions in the Tilje Formation (Figs. 1.13C and 1.20A; Ichaso and Dalrymple, 2009, 2014; Ichaso Demianiuk, 2012). Such successions are 2–4 m thick and pervasively heterolithic, but coarsen upward overall by means of an upward increase in the thickness of the sandstone laminae and/or beds, and an increase in the grain size of the sand, from fine and very fine sand at the base of each succession to fine and medium sand at the top. The sedimentary structures in the sandstone beds consist predominantly of current ripples that commonly show bipolar paleocurrent directions. Some wave-ripple and combined-flow lamination is also present and more distal examples contain decimeter-thick event beds that contain hummocky cross-stratification. Thick macroscopically structureless mudstone layers more than 1 cm thick are abundant; they have been interpreted as the deposits of fluid muds (Ichaso and Dalrymple, 2009; Reith, 2013). The trace-fossil diversity in these mouth-bar deposits is high, indeed it is the highest of the case studies examined here, including Skolithos, Arenicolites, Berguaeria, Planolites, Paleophycus, Cylindrichnus, Teichichnus, and Diplocraterion, representing a mixed SkolithosCruziana ichnofacies assemblage. Some mouth-bar successions are capped by thin (up to 1 m thick), erosively based sandstone units that contain cross-bedding formed by dunes with abundant fluid-mud layers in the dune bottomsets; such sandstone bodies are interpreted as the terminal distributary channels that fed the underlying mouth bar. Within the mouth-bar deposits, a decimeter-scale alternation is present between two different styles of sedimentation, both of which consist of interbedded sandstone

1.4 Description and interpretation of the case studies

FIGURE 1.20 Case study 6, Tilje Formation (Fig. 1.13). (A) Upward-sanding, heterolithic mouth-bar succession from the lower part of the Tilje Formation (T3 level in Fig. 1.13C), that sits on an interpreted flooding surface (FS; sinuous red (white in the print version) line) and is overlain erosively (upper sinuous red (white in the print version) line) by a terminal distributarychannel deposit consisting of cross-bedded sandstones with abundant thick mudstone layers interpreted to be fluid-mud deposits (Ichaso and Dalrymple, 2009). The section enlarged in (B) is indicated by the red (white in the print version) rectangle. (B) Enlarged view of deposits interpreted to represent a river-flood deposit, overlain by an interflood deposit. The sandstone layers in the lower (river flood) interval are coarser grained than the interflood deposit (medium sand vs. fine sand), contain dune cross-stratification (as opposed to ripple crosslamination only in the interflood deposit), and are separated by abundant thick, fluid-mud layers, whereas the interflood deposits contain only thin mudstone laminae. The intensity of bioturbation is much lower (BI 0–1) in the river-flood deposits than in the interflood sediments (BI 3–5). See text for additional discussion.

and mudstone (Fig. 1.20B). The first contains the coarsest sandstone layers in the mouth-bar succession, with grain sizes ranging from fine to medium or even coarse sand, locally with granules. The sedimentary structures in the sandstone layers are commonly ripples, but dune cross-stratification is present in some of these intervals (Fig. 1.20B). The intervening mudstone layers commonly are thick and structureless, and hence are interpreted as fluid-mud deposits (Ichaso and Dalrymple, 2009; Reith,

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2013). The degree of bioturbation in these intervals is low, generally BI 0–1, and only rarely BI 2. In the second style of sedimentation, the heterolithic deposits are, on average, more thinly bedded. The sandstone layers are noticeably finer grained (generally fine sand) and thick fluid-mud layers are absent. In addition, the degree of bioturbation is higher with BI values of 3–5. All of these differences are attributed to differences in river discharge and the associated changes in salinity and the position and intensity of the turbidity maximum (Ichaso and Dalrymple, 2009, 2014). The first style of sedimentation (i.e., coarser sandstone with abundant fluid-mud layers) has been interpreted to represent deposition during river floods, which are responsible for the introduction of the coarser sediment, the formation of the higher-energy dune cross-bedding and the lowering of salinity levels, which, together with the more rapid sedimentation, are responsible for the lower level of bioturbation in these units. The presence of fluid-mud layers in the river-flood deposits is interpreted to reflect the seaward displacement of the turbidity maximum, from a location landward of the mouth bars during low-flow periods to a location directly over the mouth bars during rivers floods. It is also likely that the peak suspendedsediment concentrations were higher during river floods (cf. Doxaran et al., 2009), thereby increasing the potential for the formation of fluid muds. The second style of sedimentation (i.e., thinner, finer sandstone layers with no fluid-mud layers) is interpreted to represent accumulation in the interflood periods, when the river currents were weaker as indicated by the development of current ripples instead of dunes, and the sand grain size was correspondingly finer. Lower suspended-sediment concentrations because of the landward migration of the less-intense turbidity maximum resulted in the formation of the thinner mudstone layers. Compared to case study 5, the river-flood deposits in the Tilje Formation are not as obvious. There is no longer a discrete, sharp-based sandstone bed representing the time of peak river discharge; instead, there is a heterolithic interval that contains the coarsest sand together with the thickest mudstone layers. A tidal signal is present throughout the river-flood deposit, in the form of the mudstone drapes and local occurrences of bipolar paleocurrents. As a result, the Tilje depositional conditions are thought to be similar to those shown in Fig. 1.8.

1.4.7 CASE STUDY 7: BLUESKY FORMATION The lower part of the Bluesky Formation from which the example used here comes (Fig. 1.21) has recently been reinterpreted to consist of a tide-dominated deltaic succession (Mackay, 2014). Much of this unit consists of erosively based, uniform to subtly upward fining, sandstone-dominated successions that are interpreted to have formed within channels and on the flanks of the tidal point bars or elongate tidal bars that bordered these channels. The successions are 5–10 m thick and contain the preserved remnants of large to very large compound dunes (cf. Ashley, 1990). Individual compound-dune successions are typically 1–3 m thick and have an upward coarsening and thickening of the sandstone layers, together with an upward decrease in the abundance of mudstone layers. Fluid-mud layers most typically occur near the

1.4 Description and interpretation of the case studies

FIGURE 1.21 Case study 7, Bluesky Formation (Fig. 1.14). (A and B) Representative intervals from channel-bar successions in the lower, heterolithic portion of the Bluesky Formation (Fig. 1.14C), which is interpreted to represent a regressive delta complex. The tide-dominated nature of the environment is indicated by the abundance of inferred neap (N)–spring cycles. In both core intervals, river-flood deposits are thought to be represented by the coarsest sandstones (medium to coarse sand) that also contain the highest energy structures (XB ¼ dune cross-bedding) and the greatest abundance of thick, structureless mudstone layers (UM2; mudstone facies codes follow Mackay and Dalrymple, 2011), that are believed to represent deposition from high-density suspensions (i.e., fluid muds). By contrast, the interflood deposits are finer grained (fine-grained sand), dominantly current-rippled, and contain predominantly thin mudstone layers (UM1) that are believed to have accumulated by slow settling from dilute muddy suspensions. See text for additional discussion.

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base of compound-dune successions, in the bottomset to toeset deposits. The thickness of the compound-dune successions, and thus the size of the dunes, combined with the abundance of fluid-mud deposits within them, suggest that they likely formed near the base of the channels (Mackay, 2014; Mackay and Dalrymple, 2011). The tidally dominated nature of the deposits is indicated by the pervasive presence of cyclic variations in the abundance and spacing of mud drapes that are interpreted to represent neap–spring cyclicity (Fig. 1.21), as well as the abundance of fluid-mud deposits which appear to be most common in settings with strong tidal currents (see references in Dalrymple, 2010; Mackay and Dalrymple, 2011), and the low degree of bioturbation as a result of high energy levels and very high turbidity. Based on these depositional conditions and their paleo-geographic position within the lower Bluesky, the deposits shown in Fig. 1.21 are interpreted to come from channels cutting the delta-plain portion of the deltaic system (Mackay, 2014; cf. Dalrymple, 2010). Initial examination of these deposits suggests that there is no evidence of riverflood sedimentation, given that all of the deposits appear to tidal in origin, as indicated by the ubiquitous presence of a complex interlayering of sandstone and mudstone that shows evidence of tidal sediments (i.e., regular changes in sandstone-layer thickness, thick–thin alternations, etc.). However, if the larger environmental interpretation is correct, then there must have been river floods to supply sediment to the delta. Indeed, if the criteria for distinguishing river-flood from interflood deposits that were developed in case study 6 are applied to the Bluesky deposits, it is possible to interpret the presence of flood–interflood cyclicity in the cores (Fig. 1.21). On this basis, the intervals that are interpreted here as the riverflood deposits are characterized by the coarsest sand (medium to coarse sand) and the thickest sandstone beds, commonly with the development of dune cross-bedding, and by the greatest abundance of thick, fluid-mud layers (UM2 mudstones in Fig. 1.21; Mackay and Dalrymple, 2011) as a result of the elevated suspendedsediment concentrations that are associated with periods of higher fluvial discharge (Castaing and Allen, 1981; Doxaran et al., 2009; Lesourd et al., 2003; Uncles et al., 2006). The interflood layers, by contrast, have slightly finer sand grain sizes (fine to medium sand), a greater abundance of ripple cross-lamination and a scarcity of fluidmud (UM2) layers (Fig. 1.21). Instead, the mudstone layers are almost invariably less than 1–3 mm thick, which implies that their deposition during tidal slack-water periods was by slow settling of individual grains (or flocs) from low-density suspensions (UM 1 in Fig. 1.21; Mackay and Dalrymple, 2011). These interflood characteristics are a result of the overall lower current speeds and suspendedsediment concentrations during periods when the river discharge was lower. Neither the river-flood nor interflood deposits are bioturbated to any extent (BI values are uniformly 0–1), presumably because depositional conditions were invariably hostile; indeed, evidence of strong tidal action is present in both the river-flood and interflood beds, which suggests that sediment mobility and sedimentation rates were consistently too high to permit colonization of the channel floor where these deposits accumulated. Thus, the depositional conditions were probably similar to those shown

1.5 Discussion

in Fig. 1.9, in which the river flood is subordinate to the tidal currents, which dominate sedimentation at all times, generating a tide-dominated deposit, in which the river-flood sedimentation is cryptic.

1.5 DISCUSSION Observations in modern environments indicate that there should be a temporal change in depositional conditions throughout much of the fluvial–marine transition that is driven by changes in river discharge. The bed-level variations in the relative intensity of tidal and river processes documented above show that this process variation is recorded in a recognizable manner in the deposits of a number of ancient successions (cf. van den Berg et al., 2007). These stratigraphic units were not chosen for the purpose of documenting tidal–fluvial interaction, but were instead chosen because they were either known petroleum reservoirs or considered to be potential analogs for reservoirs. This suggests, therefore, that the types of features documented here are likely to be common, even widespread, in coastal successions. Thus, the scheme outlined above for recognizing deposits of the fluvial–marine transition appears to offer a new tool for recognizing and interpreting coastal-plain deposits. A few of the possible applications and implications are discussed below. The suite of case studies described illustrates that this technique can be used to determine the relative degree of tidal influence within a particular deposit, and perhaps within the larger environment. Of the case studies described above, four of them (case studies 2, 3, 5, and 6) come from mouth-bar deposits. Because the depositional environment is fixed, to at least a first-order approximation, then their relative ranking in terms of tidal influence/dominance can be assessed: case study 1 (lower Lajas Formation) is the least tidally influenced and is clearly river-dominated (Kurcinka, 2014), whereas case study 6 (Tilje Formation) is the most tidally influenced, with case study 6 (lower Neslen Formation) being a close second with regard to the intensity of tidal influence. Whether or not these last two examples are truly tidedominated remains to be seen, given that the river-flood signal is still clearly visible. Case study 7 (Bluesky Formation) shows that the river-flood signal can become even more cryptic; thus, case studies 5 and 6 are presumed to plot somewhere in the middle of the river-tide side of the delta triangle; wave-generated sedimentary structures were not present in case study 5, although they are present but minor in case study 6. The evaluation of how the river-tide structures change with proximal–distal location within tidal–fluvial channels (cf. Figs. 1.5–1.9) cannot be demonstrated with the isolated case studies discussed here, although there is a reason to believe that the proposed gradient shown in Figs. 1.5–1.9 is a reasonable starting point. It should be noted, of course, that systems with overall weak tidal currents will not develop features like those predicted by the conditions shown in Figs. 1.8 and 1.9. Thus, there is no reason to expect the full spectrum to be present in any given deposit. Certainly, of the three tidal–fluvial channel examples discussed here (case studies 2, 4, and 7), the first is the least tidally influenced, whereas case study 7 is the most strongly tidally

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influenced/dominated. However, whether this represents a difference in location relative to the tidal limit (the McMurray example being the closest to the tidal limit in such an analysis), or whether the differences are due to differences in the tidal range or the magnitude of river floods, or both, is impossible to tell from isolated case studies. In order to test the ideas presented here, systematic studies are needed within a series of single, well-exposed ancient systems, or better yet, in modern systems (cf. Johnson and Dashtgard, 2014; La Croix and Dashtgard, 2014), in order to ascertain whether the predicted longitudinal trend (cf. Figs. 1.5–1.9) is reproduced, or not. In any application of these concepts to modern or ancient deposits, care must be taken to evaluate the potential for significant local spatial and temporal variability in the relative importance of river and tidal currents. Local variations are to be expected because of vertical changes in current strength on individual point bars (Gugliotta et al., 2015b), the presence of ebb- and flood-dominated channels (Dalrymple, 2010), and variations in the relative strength of tidal and fluvial currents around individual point bars (Choi, 2010; Dalrymple and Choi, 2007). In addition, there will be stochastic variability in the magnitude of river floods, as well as semiregular variations in the size of the flood peak due to such climatic factors as the El Nin˜o Southern Oscillation and North Atlantic Oscillation ( Jablonski and Dalrymple, 2015), all of which will cause the tidal and fluvial signals to be recorded differently (e.g., the difference between the deposits of larger and smaller floods discussed in case study 2; cf. Fig. 1.16). Such spatial and temporal variations have to be averaged out before the relative importance of river and tidal currents and the proximal–distal location can be determined with confidence. Precisely how the physical evidence of location within the fluvial–marine transition corresponds with the assessment based on the ichnological assemblages that are a function of salinity remains to be determined. It must be noted, however, that the ichnological character of a succession almost certainly reflects primarily the depositional conditions that existed between river floods (i.e., during the lengthy interflood periods when river discharge was low; Gingras et al., 2002). During the relatively short periods when the river is in flood, depositional conditions throughout much of the fluvial–marine transition are likely to be significantly more stressed than during low-flow periods, not only because the salinity would be lower, but also because sedimentation rates, energy levels, and suspended-sediment concentrations would have been higher. River-flood periods might also have been too short for the establishment of an equilibrium community, and any organisms that could have withstood the more stressed conditions might not have had time to bioturbate the sediment to any significant degree. By contrast, interflood periods are generally longer and have less biologic stress, thereby allowing more intense colonization by a greater diversity of organisms. The result is that the deposits of the tidal–fluvial transition are likely to have a binary ichnological character that reflects the two environmental states that characterize this zone. This opens the possibility of being able through the combined use of the physical and biogenic structures to deduce depositional conditions during both river floods and interflood periods, and, by extension, to gain an

1.6 Conclusions

appreciation for the difference in river discharge between high- and low-flow periods as was done above in a comparison of the observations between case studies 1 and 2. Finally, the tidal deposits within flood–interflood successions hold the potential to act as a clock that might allow duration of river floods and the time interval between them to be estimated. Thus, in case study 4 (Fig. 1.18B) is it possible to calculate that the waning stage of the river flood lasted for 20 tides (i.e., 10 days, assuming a semidiurnal tide, the deposition of only one lamina per tide, and the presence of no missing tides in the neap–tide portion of the succession). By a similar calculation, the river flood shown case study 5 (Fig. 1.19B), which contains layering that is likely of tidal origin, might have had a total duration of approximately 1.5–2 weeks. Determining the interval between floods is perhaps more problematic, given that sedimentation rates are commonly too low to generate recognizable tidal rhythmites and that bioturbation commonly disrupts sedimentation during the interflood period. The nature of the interflood bioturbation (e.g., burrow size, number of spreite, etc.) can in at least some cases constrain the length of the interflood period: Gingras et al. (2002), for instance, have suggested that the bioturbated intervals in the IHS examined by them required many months to form, and a similar temporal estimate has been suggested for the bioturbated horizons in a subsurface example of the McMurray Formation that appears to be similar to those described here in case study 2 (Fig. 1.16; J. MacEachern, 2009, personal communication). Such temporal estimates for the length of the interflood period can be used to imply the existence of only one major river flood each year in these systems. (In the case of the McMurray Formation, this is consistent with the continental scale of the river (Blum and Pecha, 2014), which favors the interpretation that its discharge regime was “climatedominated” as defined above rather than responding predominantly to local weather events.) Admittedly, such calculations of the duration of river floods and interflood periods are subject to considerable uncertainty, but a technique for estimating these quantities has not existed before, so the ultimate utility of this approach to reconstructing fluvial regime and paleoclimate opens new avenues for research.

1.6 CONCLUSIONS In the absence of systematic, detailed studies of modern or ancient coastal successions, this chapter has used general considerations of the temporal variation in the intensity of river and tidal processes in the fluvial–marine transition to develop a process framework for the nature of the deposits within the transition, at the scale of an individual bed. This is then extended to propose a spectrum of deposit types ranging from fluvially dominated to tidally dominated. It is based on the assumption that most rivers experience discharge fluctuations that cause depositional conditions to alternate between river floods (high-flow periods, when tidal influence is less) and interfloods (low-flow periods, when tidal influence is greater). In areas that are overall fluvially dominated (e.g., in the proximal part of the fluvial–marine

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transition or in areas with small tides), local depositional conditions are likely to be fluvially dominated during river floods, and tidally influenced to even dominated during low-flow periods. In areas that are overall tidally dominated (e.g., more distal locations and especially in systems with a large tidal range and strong tidal currents), the river-flood deposits will become progressively overprinted by tidal structures (e.g., tidal rhythmites, bipolar current ripples, etc.), with the tidal signature appearing first in the waning stage of the river flood, and gradually extending throughout the entire river-flood deposit as the system becomes more tide influenced or tide dominated. In the extreme, the river-flood deposit can become cryptic and might be expressed as intervals with subtly coarser sand sizes and a greater abundance of fluid-mud deposits, because of the enhancement and seaward displacement of the turbidity maximum because of the enhanced fluvial discharge. The predictions made in this study are supported by a series of isolated case studies, but much more work is needed to test the predictions, and to ascertain to what extent spatial and temporal variations in the relative intensity of river and tidal currents can overprint the longitudinal trends that are suggested here. The concepts and observations presented suggest that the ichnology of the fluvial–marine transition zone primarily reflects depositional conditions during the low-flow periods when the river system is least active and the salinity is the highest. Thus, the simultaneous use of the physical and biological structures has the potential to reconstruct the nature of the depositional environment during times of both high and low flow, and thus to deduce the difference in discharge between river-flood peaks and base flow.

ACKNOWLEDGMENTS We thank the industrial sponsors of the various projects that have contributed to the analysis presented here: BHP Billiton, Carigali-HESS, Devon, Nexen, Shell, Statoil, Woodside, TOTAL, and VNG Norge. R.W.D. also thanks the Natural Research Council of Canada for continued support of his research on tidal depositional systems. We also thank Marcello Gugliotta for sharing images and insights into the sedimentology of the Lajas Formation and Cornel Olariu for sharing his knowledge of the Neslen Formation. Our thanks are also extended to the reviewers who provided comments that helped to improve the presentation.

REFERENCES Allen, P.A., 1997. Earth Surface Processes. Blackwell Science, Oxford. 404 pp. Allen, G.P., Salomon, J.C., Bassoulet, P., Du Penhoat, Y., De Grandpre´, C., 1980. Effects of tides on mixing and suspended sediment transport in macrotidal estuaries. Sediment. Geol. 26, 69–90. Aschoff, J.L., Steel, R.J., 2011. Anatomy and development of a low-accommodation clastic wedge, Upper Cretaceous, Cordilleran Foreland Basin, USA. Sediment. Geol. 236, 1–24. Ashley, G.M., 1990. Classification of large-scale subaqueous bedforms: a new look at an old problem. J. Sediment. Petrol. 60, 160–172.

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CHAPTER 1 Physical record of fluvial–tidal interaction

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