12. DEEP-MARINE SEDIMENTS AND SEDIMENTARY SYSTEMS

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1 1 12. DEEP-MARINE SEDIMENTS AND SEDIMENTARY SYSTEMS R. William C. Arnott, Department of Earth Sciences, University of Ottawa, Ottawa, ON K1N 6N5 INTRODUCTION The deep marine is a unique sedimentary environment compared to all others because of its inaccessibility and the enormous spatial scale of many of its constituent depositional systems. For, example, the modern Bengal Fan, which has been accumulating sediment for only about 55 m.y., is 3000 km long, over 1400 km wide, >5000 m thick, and contains an estimated sediment volume of 4x10 6 km 3 (Table 1). (For a global inventory of fan systems, see the wall chart of Barnes and Normark, 1985). Accordingly, the principal investigative tool for modern deep-marine environments is seismic, including an array of techniques that range from highfrequency, high-resolution but shallow-penetrating surveys, to lower frequency, more deeply penetrating but lower resolution surveys. The 3-D seismic with its ability to image features in both plan and cross-sectional views has proven to be especially useful. Over the past decade, this work has resulted in the publication of many stunning seismic images that have improved greatly our understanding of the deep-marine sedimentary system (e.g., Weimer et al., 1991; Posamentier and Kolla, 2003; Posamentier and Walker, 2006). Nevertheless, the seismic method generally suffers from limited vertical resolution minimum vertical resolution of industry seismic data is commonly of the order of 10 meters. To fill the gap, the geological community must also utilize outcrop studies, but here too a number of shortcomings are recognized (Fig. 1). Perhaps most importantly, the outcrop record is inherently 2-dimensional with, at best, local 3-D perspectives (i.e., valley cuts). Also, in most cases, the horizontal scale of most outcrops is small compared to the spatial scale of most deep-marine architectural elements, and, more profoundly, their parent sedimentary system. Nevertheless, over the past decade or so, much research, including work on a number of seismic-scale outcrops, is helping to bridge the gap between the ancient outcrop record and modern seismic images, to merge these two independent datasets into a single, coherent package (see, for example, Nilsen et al., 2007). Nevertheless, the emphasis here, as in other chapters in this volume, is on the lithological characteristics of the strata that make up the deep-marine sedimentary record, and also on how these strata are distributed spatially (space) and vertically (time) in the geological record. Excellent reviews of the seismic-scale attributes can be found elsewhere (e.g., Posamentier and Kolla, 2003; Posamentier and Walker, 2006). BRIEF HISTORY OF DEEP-MARINE SEDIMENTOLOGY Turbidites are a ubiquitous feature in the deep-marine sedimentary record, and are deposited from subaqueous turbidity currents. However, because the formative process is largely hidden from direct observation, the connection between process and deposit remained speculative and anecdotal since as early as the late 1800s. Knowledge of deep-marine processes took a major step forward following the laying of the first successful Atlantic seafloor telegraph cable in In 1899, Benest (in Heezen and Ewing, 1952) reported that accidents to cables have already been valuable... in directing attention to hitherto unsuspected forces constantly in action altering the features of the sea bottom. The most famous cable disruption occurred on November 18, 1929, on the southern margin of the Grand Banks off east coast Canada. On that day, an estimated 7.2 magnitude earthquake occurred and disrupted telegraph communication between North America and Europe. Several submarine cables were apparently broken at the time of the earthquake, presumably as a result of seabed movement. More puzzling, however, was the fact that a number of other cables became deactivated successively after the earthquake, some as much as 13 hours later (Heezen and Ewing, 1952; Piper et al., 1999). Furthermore, the cables broke progressively in a single (offshore) direction. It was hypothesized that a turbidity current moving at considerable speed was responsible (Heezen and Ewing, 1952). At about the same time, Kuenen and Migliorini (1950) began experimenting with sediment dispersions released into a basin of still-standing water. Upon release the dispersion formed a bottom-hugging turbidity current that produced a deposit with a characteristic upward decrease in grain size. Based on these results, these authors suggested that the common occurrence of upward-fining beds in the deepmarine geological record might reflect deposition from turbidity currents. Beds showing this upward-fining character were observed in flysch deposits of the European Alps by Bouma (1962). In addition to the upward fining, he also noted a characteristic vertical succession of sedimentary structures, the origin of which was then unknown. Shortly thereafter, experimentalists illustrated the variety of bedforms that formed under unidirectional currents of various speeds (Simons et al. 1965; and now many others). Based on those results the characteristic suite of sedimentary structures observed by Bouma, the so-called Bouma turbidite sequence, reflects deposition from a decelerating turbidity current. At about the same time, Middleton (1966a, b, 1967) published a series of important papers based on flume experiments, describing the various

2 2 ARNOTT Table 1. Dimensional characteristics of five deep-marine turbidite systems. Data from wall chart of Barnes and Normark (1984). Length, Maximum Name Location Age Width Thickness Volume Dominant Range (km) (m) (km 3 ) Grain Size Amazon Brazilian Middle 700, x10 5 mud pebbles margin Miocene 700 (max) to mud Bengal Bay of Eocene 2800 (min), >5000 4x10 6 mud mud to Bengal 1400 (max) medium sand Laurentian coast Quaternary 1500 (max), x10 5 mud to med- mud to eastern 400 fine sand gravel Canada Mississippi Gulf of Pleistocene 540, x10 5 silty mud mud to Mexico gravel Navy coast Late sandy silt mud to southern Pleistocene gravel California numerous overlapping slumps formed by progressive retrogressive failure, moved downslope and in areas of high slope transformed into a debris flow and ultimately into a sustained, fast-moving (~70 km/hr) turbidity current. This current eroded on the continental slope and eventually deposited a turbidite up to 1 m thick on the Laurentian submarine fan. The turbidite contains about 185 km 3 of sediment using railway boxcars to conceptualize volume, 185 km 3 equates to 1.29 billion boxcars, or a train almost 22 million kilometers long, and one that would wrap around the world over 558 times! Figure 1. Vertical and horizontal scales of modern deep-marine systems (Bengal, Monterey and Navy fans), ancient turbidite systems (Butano), some deepmarine sediment features/elements, and outcrop and core (adapted from Barnes and Normark, 1985). Note how the scale of most outcrops is dwarfed by the scale of seismic data sources, but more importantly, the size of deep marine fans and their constituent depositional elements. parts of a turbidity current and the influence of sediment concentration on depositional patterns and characteristics. Today, it has been well established that sediment-gravity flows, principally turbidity currents and debris flows, in addition to mass-movement processes, are the major formative agents of the deep-marine sedimentary record. A superb illustration is the 1929 Grand Banks event, which is now known to have been initiated by widespread failure of a surficial layer about m thick (Piper et al., 1999). The slide, which comprised SEDIMENT TRANSPORT MECHANISMS AND DEPOSITS Unlike continental and shallowmarine sedimentary systems, sand and gravel transport in the deep marine is dominated by sedimentgravity flows and mass-movement processes. Finer grained sediments, principally silt and clay, although present in sediment-gravity flows and their related deposits, are transported mostly in suspension. Mass-Movement Deposits Slides and Slumps The gravity-driven downslope movement of coherent to semi-coherent masses of sediment along discrete

3 12. DEEP-MARINE SEDIMENTS AND SEDIMENTARY SYSTEMS 3 Figure 2. Mass-movement deposits. A) Slide deposit (strata are vertically dipping due to later (Cordilleran) tectonic deformation and stratigraphic top is toward the left). Base of slide is indicated by the solid black line. Note the extensively deformed and brecciated strata in the lower part of the slide, including a large rotated block (Neoproterozoic Isaac Formation, British Columbia). B) Slump deposit with extensive internal ductile deformation (base of slide indicated by solid white line; Carboniferous Gull Island Formation, Ireland). failure planes is termed mass movement or mass wasting. Such movements, which can range up to hundreds of kilometers in distance, occur when the driving force, gravity, exceeds the tensile strength of the parent sediment pile (which depends upon a wide range of internal conditions, including pore-fluid pressure, sediment composition and consolidation and the occurrence of mechanically weak layers among others). The mechanisms responsible for triggering the initial instability include oversteepening, seismic loading, cyclic storm-wave loading, rapid accumulation and underconsolidation, gas charging, gas-hydrate dissociation, seepage, etc. (e.g., Locat and Lee, 2002). Once initiated, movement continues until the resisting forces, principally friction along the basal failure plane, exceed the gravitational driving force and en masse deposition takes place. Following earlier workers, two end-member kinds of mass-movement deposits are recognized: slides and slumps, where the principal difference between the two is the intensity and nature of internal deformation. In the case of slides, deformation is comparatively minor and is dominated by brittle deformation, mostly in the form of beddingparallel surfaces of detachment, although shear deformation may be intense in a thin zone near the base of the unit (Fig. 2A). In slumps, deformation shows an element of rotation and typically is more intense and ductile in character (Fig. 2B). Softsediment slump folds are abundant, and are commonly tightly folded with the axial plane sub-parallel to planes of internal shear. Sediment-Gravity Flows Gravity currents occur when a moredense fluid moves through and displaces a less-dense fluid. Sedimentgravity currents are a type of gravity current, wherein the density contrast is produced by the presence of suspended sediment. Early classification schemes for sediment-gravity flows recognized four types of endmembers: debris flows, grain flows, fluidized/liquefied flows, and turbidity currents (Middleton and Hampton, 1976), each differentiated by the primary mechanism for suspending sediment and maintaining the density contrast: matrix strength, grain collision, escaping pore fluid and fluid turbulence, respectively. More recently, an alternative scheme by Mulder and Alexander (2001) proposed a classification based on flow properties and sediment-support mechanisms, and recognized two types of end-member flows: cohesive flows and frictional flows. Here, a modified version of this classification will be adopted. Cohesive Flows Debris Flows Cohesive flows, which are more commonly termed debris flows and mud flows, are sediment-gravity flows where the volume concentration of the solid and fluid phases are of the same order of magnitude, and the occurrence of a cohesive matrix imparts a pseudoplastic rheology to the flow (e.g., Mohrig et al., 1998; Mulder and Alexander, 2001). Hereafter the term debris flow, which is entrenched in the geological literature, will be used to refer to all cohesive flows. Particles are principally suspended by cohesive forces provided by a matrix of fluid and finegrained sediment (generally a silt clay mixture. Note that in both, subaerial and subaqueous debris flows (see also Chapter 6), the amount of clay-size particles needed to generate sufficient yield strength to suspend larger particles can be surprising low, possibly of the order of 2 4% by volume. In addition to matrix strength, buoyancy effects, particle-particle interaction (collisions and near misses), hindered settling, elevated pore pressure, and in some cases fluid turbulence may provide additional support for suspended particles. Deposition occurs when one or more of the following exceed the driving gravitational force: intrinsic shear strength of the sedimentwater mixture, grain-contact friction and friction along the flow bound-

4 4 ARNOTT Figure 3. A) Ancient example of a partly debris-flow-filled channel. Base of channel is indicated by dashed black line. After accumulating several meters of sandstone, the channel became partly plugged by a debris-flow deposit (DF; Neoproterozoic Isaac Formation, British Columbia). Thereafter, channel-filling sandstones onlap and then overlap the debris-flow deposit. B) Seismic image of the chaotic reflections of a debris-flow deposit that has exploited a pre-existing seafloor channel (photo courtesy of Henry Posamentier). aries, which then causes the flow to freeze inward, either en masse, or more gradually from areas of lower shear near the flow s surface toward those of higher shear at the base. Debris-flow deposits, or debrites, form sheetlike to lobe-shaped masses that have steep margins as a result of the strength of the moving and deposited sediment mass. Deposits range widely in scale, but can be up to tens of kilometers wide and over 100 m thick, although much thinner beds are more common. At their downflow terminus toe-thrusts are common due to a rapid down-flow reduction in flow speed. Bases of debris-flow deposits are commonly planar and non-erosional, because of the strength of the moving mass and the damping of large-scale fluid turbulence. Nevertheless, deeply scoured bases are observed. In some cases they represent the opportunistic occupation of a pre-existing seafloor channel (Fig. 3), whereas others are thought to be formed by a rigid part of the flow ploughing through the underlying seafloor sediment (see Posamentier and Walker, 2006, their Figures 155, 157) or being dragged along the surface forming linear grooves up to 40 m deep, several hundred meters wide and extending longitudinally for more than 20 km (Posamentier and Kolla, 2003). The upper surface can also be uneven and ranges between flat to highly rugose. Internally, debris-flow deposits range from mud- to sandrich, typically with a disorganized, poorly sorted character (e.g., Nemec and Steel, 1984; Fig. 4A). On seismic images debris-flow deposits, especially those interpreted to be mud rich, exhibit a distinctive chaotic or reflection-free character. Where present, clasts ranging from sand grains to enormous blocks are generally dispersed throughout a fine-grained matrix. In some deposits, clasts of incorporated soft sediment are contorted and commonly show a subtle to well-developed orientation with their longest dimension oriented subparallel to the base of the deposit (Fig. 4B). Preferential particle alignment and particle deformation is likely the result of shearing within the sediment mass during transport, which then may be accentuated by post-depositional compaction and consolidation. Although cohesive, some debris flows are capable of movement over long distances. For example, Gee et al. (1999) reported a modern debris flow off the west coast of Africa with a run-out distance of about 700 km. Such large distances may be attributed to hydroplaning and the attendant reduction of friction between the bed and the overlying flow. Elevated hydrodynamic pressure exerted on the forward part of the flow causes a wedge of ambient fluid (i.e., seawater) to penetrate beneath the flow and separate it from the bed (Mohrig et al., 1998). As a result, frictional resistance at the base of the flow is significantly reduced and the weight of the overlying debris flow becomes borne by the fluid. Importantly, the overriding debris flow must have sufficiently low permeability (i.e., be sufficiently muddy) so that the basal fluid layer does not dissipate. Debris flows are also capable of spawning turbidity currents. Owing to mixing along its leading edge, sediment is eroded from the debris flow and cast into a developing turbulent suspension above the flow. However, owing to the low permeability of most debris flows, water infiltration is minimal and the amount of sediment eroded and transferred into the turbidity current is very small (<1%). Also, so long as the debris flow and overlying turbulent flow are moving at a similar velocity, the amount of erosion along the interface between them is negligible. But, if the debris flow stops, then the turbulent suspension can rework the top of the debris-flow deposit as it becomes detached from the parent debris flow and continues farther downslope (Fig. 4C). Conversely, it has been reported that, at an abrupt reduction in slope, turbidity currents can be partially transformed into a debris flow. As the turbulent flow decelerates rapidly and turbulence is reduced, suspended particles fall rapidly toward the bed where the ubiquitous occurrence of cohesive mud particles eventually increases yield strength to the point where the particles become supported by matrix strength. Mass transport deposits (MTDs)

5 12. DEEP-MARINE SEDIMENTS AND SEDIMENTARY SYSTEMS 5 Figure 4. Debris-flow deposits (debrite). A) Schematic diagram showing typical characteristics of subaqueous debris-flow deposits (from Nemec and Steel, 1984). B) Sharp-based debris-flow deposit overlying a succession of thin- and mediumbedded turbidites (below the dashed line). Note the dispersed quartz pebbles and mudstone clasts, some of which have been folded (arrow) due to shearing within the moving cohesive mass (pencil for scale). C) Pebbly debris-flow deposit overlain by a T bc turbidite (pencil for scale). Black layer is a large mudstone clast. The anomalous concentration of quartz pebbles in the upper part of the debrite is attributed to reworking by a genetically related turbidity current that may have deposited the overlying T bc turbidite. The base of this bed (dashed line) has loaded passively into the underlying debrite, which at the time must have been water saturated and poorly compacted (Neoproterozoic Isaac Formation, British Columbia). are observed often in shallow- and deep-penetrating seismic images. The MTDs are typically erosively based deposits that occur on a range of scales, including enormous features that cover areas up to several 1000 km 2 and are several 100s of meters thick. The dimensions of MTDs, including their length:width ratio, are controlled by the geomorphology and position of the sediment source, wherein large, expansive deposits are related to instability initiated along the upper part of the basin margin, at or near the shelf edge, whereas smaller deposits form in the basin by failure along local slopes (detached sediment sources; Moscardelli and Wood, 2008). Internally MTDs consist predominantly of a complex, typically disorganized assemblage of slump, slide and debris-flow deposits that can locally be interstratified with (organized) channel and overbank strata. The occurrence of large MTDs in the stratigraphic column represents a major change in sedimentation regime within the basin. In many cases these changes have been interpreted to be related to allocyclic processes, and therefore to have sequence-stratigraphic significance (see sequence-stratigraphy section below). Frictional Flows According to Mulder and Alexander (2001), frictional flows form a continuum from mass movements (slumps/ slides see above) to a variety of different kinds of sediment suspensions subdivided on the basis of the dominant mechanisms of sediment support. The most common the friction-

6 6 ARNOTT Figure 5. A) Well-developed head of an experimental turbidity current. B) Line diagram illustrating the typical shape and velocity profile of a turbidity current (modified after Kneller and Buckee, 2000). Note that, unlike an open-channel flow such as a river, the velocity maximum occurs in the lower part of the flow. Note also the extensive mixing (due to interfacial instability) that occurs along the upper part of the current. al flows, however, are turbidity currents. Turbidity Currents Based on flow and sediment characteristics, turbidity currents consist of three distinct but not sharply bounded parts: head, body and tail (see review by Kneller and Buckee, 2000; Fig. 5). The head of turbidity currents is the sediment-rich part of the flow and the site where most of the mixing with the ambient fluid occurs. It is characterized by a sharp, overhanging nose, above which the head slopes back in the upstream direction due to the resistance of the stationary overlying fluid. This generates a strong shear that sweeps sediment-rich fluid from the head backward toward the body of the current. To sustain the current, the head must be continually provided with new sediment supplied from the body, which moves forward faster than the head. Because of differences in settling velocity, coarser sediment tends to accumulate in the lower part of the head whereas finer sediment is moved upward and backward into the body of the current, the consequence being, that with time, the flow becomes longitudinally differentiated in terms of grain size. At the tail of the flow, sediment concentration is low and, as a consequence, flow speed is slower and eventually decreases to zero. Increasingly it is being recognized that most natural turbidity currents are moderately to highly density stratified. Moreover, most natural turbidity currents are typically of higher sediment concentration and made up of sediment significantly more poorly sorted and coarser grained than that contained in laboratory currents. Unfortunately, the effect of sediment concentration, especially high concentrations, on the nature of the current and how it deposits sediment, remain poorly understood. In the early 1980s, Lowe (1982) published a theoretical classification for turbidity-current deposition, wherein he recognized two kinds of turbidity current and their related deposits: low-density and high-density. Classical turbidites, as described originally by Bouma (1962), are interpreted to be deposited by low-density turbidity currents (see also Mulder and Alexander, 2001). The adjective, low-density, refers to the concentration of suspended sediment in the flow, which, based on the earlier work of Bagnold (1954, in Mulder and Alexander, 2001), is thought to be approximately 9 % sediment volume, or less in lowdensity flows (Fig. 6A). Above that value, the closeness of adjacent grains begins to damp fluid turbulence, and hence turbulence alone is insufficient to suspend sediment fully, especially coarse sediment. Accordingly, additional support mechanisms like dispersive pressure, hindered settling, and buoyancy are needed and become increasingly more effective with high sediment concentration. Turbidites deposited by low-density turbidity currents consist of all, or part, of the idealized succession described by Bouma (1962), and reflects decelerating flow speed (Fig. 7). However, if characteristics of the deposit were formed simply by a decelerating unidirectional shear flow, then for sediment coarser than lower fine sand (>0.15 mm), upper stage plane bed (b-division) should not be succeeded by current ripples (c-division), but instead by dunes (medium-scale cross-stratifica- tion). In addition, in rare instances, dune cross-stratified sandstone occurs where it should sandwiched between upper-stage plane bed and current-ripple cross-stratification. Therefore, what intuitively should be the norm is in fact the exception but why? One explanation that has been advanced is that most of the decelerating turbidity currents passed too quickly through the dune stability field. Although appealing, it has been argued that dunes can be formed from a flat bed in a matter of a few tens of minutes, and that many natural turbidity currents persisted at flow speeds in the dune stability field for much longer periods (Arnott and Hand, 1989). Another suggestion has been that, under high rates of sediment fallout from suspension, upperstage plane bed remains stable because the formation of dunes and in some cases, also ripples, is inhibited (Lowe, 1988). However, in many turbidites the ripples that formed the c-division show a negligible angle of climb, indicating that fallout rates are, in fact, commonly low. An alternative explanation for the absence of dunes might be the effect of high near-bed sediment concentration on the inception of dunes. Under such conditions, bed-surface defects are prevented from being amplified into dunes and, as consequence, plane bed persists, even at flow speeds that in a clearwater flow would form dunes. Finally, when near-bed sediment concentrations have been sufficiently reduced that bed defects can grow into bedforms, the flow is moving too slowly and/or the sediment is too fine to form dunes, and ripples form instead. In the case of a T bd turbidite, where neither dune nor ripple cross-stratifica-

7 12. DEEP-MARINE SEDIMENTS AND SEDIMENTARY SYSTEMS 7 Figure 6. Idealized sediment-concentration profile in a low concentration A) and high concentration B) turbidity current. C) Succession of stacked decimeter- to meter-thick T a /S3 beds. Inset depicts a typical structureless, coarse-tail graded T a /S3 bed. D) Coarse-tail, graded, structureless T a /S3 bed. Arrows in C, D indicate stratigraphic tops. Figure 7. Line diagram illustrating a complete idealized turbidite (after Blatt et al., 1984). The lowermost a-division consists of normally graded sandstone or conglomerate deposited from suspension, which then is overlain by the planar stratified b-division (upper plane bed). This, in turn, is overlain by small-scale cross-stratified sandstone (formed by current ripples) of the c-division, overlain by the subtly to well-interlaminated sandstone/siltstone and mudstone d-division capped finally by mudstone of the e-division. Photo on right is an outcrop example of a T bcde turbidite that ranges from mediumgrained sandstone at its base to silty mudstone at the top. The a-division is missing because deposition started with bedload transport. tion is present, it is argued that sediment concentration remains sufficiently high for long enough that neither dunes or current ripples form and plane bed remained stable until the end of traction transport. A significant part of the sand- and gravel-rich deep-marine sedimentary record consists not of classical Bouma turbidites, but instead of structureless, normally graded and, less commonly, ungraded or inversely graded sandstone and conglomerate, which equate to the T a division of Bouma (1962), or the S3/R3 division of Lowe (1982). In general, such beds are several decimeters to a few meters thick (Fig. 6C) and are either amalgamated or separated by a thin finer grained interval. Also, beds are characteristically poorly sorted and coarse-tail graded, wherein only the coarsest part of the grain-size distribution fines upward, typically with an upward decrease in the abundance of the coarsest grains (e.g., Sylvester and Lowe, 2004; Fig. 6D). Mudstone intraclasts are common and, in many cases, are concentrated near the top of the bed. In addition, water-escape features, including pillar and dish structures, are common in some beds. However, in spite of their abundance, the origin of structureless beds remains a major source of debate in the geological literature (see, e.g.,, Stow and Johansson, 2000), although there is growing consensus that they are deposited rapidly from high-concentration (high-density) suspensions sedimentation being so rapid that any lamination that might normally be produced by bed-load transport is not visible. Under these conditions sediment raining from suspension (i.e., capacity-driven sedimentation; see Chapter 4) entraps fine-grained sediment that otherwise could be maintained in suspension, but nevertheless gets incorporated into the accu-

8 8 ARNOTT mulating bed, producing a relatively poorly sorted deposit (e.g., Sylvester and Lowe, 2004). With time and a reduction in sedimentation rate, beds become better sorted and, in some cases, are overlain by tractional sedimentary structures, especially planarlamination, producing T ab -type successions; but what about the origin of beds that do not grade upward into planar lamination or other tractional structures? The most obvious explanation would be that the finer grained upper part of the bed was eroded by a later transport event. Although appealing, it cannot explain beds overlain by a fine-grained, typically mud-rich, layer. Here, the lack of tractional structures at the top of the bed must be related to highly efficient sediment bypass following the earlier episode of rapid sediment fallout. During bypass, sediment fallout all but ceases and dunes and ripples are prevented from forming by the maintenance of high sediment concentration, which lasts until low-energy conditions and fine-grain sediment fallout takes place from the tail of the flow. DEEP-MARINE ARCHITECTURAL ELEMENTS The origin and characteristics of deep-marine clastic systems depend on a complex assemblage of autogenic and allogenic processes, including: changes of global sea level, tectonics, sediment flux and composition, and the nature of the sediment supply system. For example, sediment supply controls the volume and internal stratigraphy of the system, whereas the number and nature of sediment entry points controls its morphology and sediment distribution. Also, grain size, which is a function of climate and provenance, controls sedimentation patterns. Based on these controls, Reading and Richards (1994) classified turbidite systems based on morphology, recognizing point-source fans, multiple-source ramps, and line-source aprons. Further, based on the dominant grain size, they also distinguished between gravel-rich, sand-rich, mixed sandmud, and mud-rich systems (Fig. 8). Gravel- and sand-rich systems tend to be small (radius of a few to a few tens of kilometers) and grade rapidly to fine-grained basin-floor deposits. Sand-mud systems, however, are much larger (radius up to a few hundreds of kilometers) and exhibit a systematic change in depositional elements and their internal stratigraphy down the transport pathway. It is these systems that make up much of the sandstone-rich part of the ancient deep-marine sedimentary record, and accordingly form much of the subsequent discussion. Mud-rich systems are the largest, and range in radius from several tens to a few thousand kilometers, and today represent the most voluminuous deep-marine sedimentary systems (see Table 1). Collectively, mud-, sand- and gravel-rich deep-marine sedimentary systems have generally been termed deepsea fans because of their common semi-conical shape. However many modern deep-sea fans, and by extension ancient fans, are in fact elongate or irregularly shaped, and therefore the more generic term turbidite system is more appropriate (Bouma et al., 1985). In addition, confusion exists because the various parts of a turbidite system, which in a downflow direction consist of the upper, middle and lower fan, have been variously defined by different authors. The recent classification by Pirmez et al. (2000), which is based on the modern Amazon Channel, subdivides the system based on the spatial patterns of sediment erosion and deposition (Fig. 9). From proximal to distal, the fan subdivisions are: submarine canyon zone of net erosion; upper fan a zone of net sediment bypass, wherein the channel thalweg (the deepest part of the channel), which is bounded on both sides by constructional levees, lies at about the same elevation as the surrounding seafloor (i.e., the area external to the channel); middle fan zone of net sediment deposition caused by flow expansion related to loss of flow confinement, and accordingly, where the thalweg lies generally above the surrounding seafloor; and lower fan the area lying downflow from the middle fan where the rate of net deposition decreases and the thalweg elevation more closely approximates the elevation of the surrounding seafloor. Internally turbidite systems are made up of an assemblage of depositional elements, which according to Mutti and Normark (1991), are the basic mappable components of both modern and ancient turbidite systems and are characterized by a distinctive assemblage of facies and facies associations. In this chapter, four depositional elements are discussed in the context of a point-sourced fan system: channels, levees, overbank/ crevasse splays, and depositional lobes. These elements, and this kind of turbidite system, appear to make up much of the sandstone-dominated part of the deep-marine record (see also Mutti et al., 2003; Wynn et al., 2007). Channels As in many sedimentary systems, channels are a common element in deep-marine settings. A channel is a negative topographical element produced mostly by confined turbidity currents that transport sediment along a major, long-term pathway. However, channels can also be sites of sediment deposition or erosion. Like channels in the continental realm, the condition of erosion, bypass or deposition is controlled by the sediment characteristics and boundary conditions of the system changes in one or both of these parameters will effect a change in the channel system. Deep-marine channels, like fluvial channels, continually seek a longitudinal profile graded to a base level, which in the deep marine is generally taken to be gravity base, but, more practically, is defined as the position where the flow becomes unconfined at the upcurrent end of the terminal lobe (e.g., Pirmez et al., 2000). Where the channel gradient and sediment-transporting flows are in equilibrium, channels migrate laterally along a plane parallel to the equilibrium profile and most of the sediment is bypassed to areas farther downflow. Erosion and deposition, on the other hand, represent conditions where the channel profile lies above and below the graded profile, respectively (see Chapters 2 and 6). Based on observations from modern and ancient systems, three kinds of deep-marine channels are recognized: erosional, depositional and mixed erosional-depositional (Fig. 10). Erosional channels are bounded by a scour surface that clearly trun-

9 12. DEEP-MARINE SEDIMENTS AND SEDIMENTARY SYSTEMS 9 Figure 8. Idealized models for deep-marine sedimentary systems based on volume, caliber and nature of sediment input (from Stow and Mayall, 2000, based on Reading and Richards, 1994). Figure 9. A) Plan view of the Amazon Fan (modified after Flood et al., 1991). Note the basinward change from a singlethread channel to a distributive network, the longest system being the modern Amazon Channel, which extends 900 km beyond the present shelf-slope break). B) Basinward transect along the modern Amazon Channel (from Pirmez et al., 2000). Subdivision of the system is based on the relationship between the elevation of the channel thalweg and levee crest relative to the adjacent seafloor surface. Datum (zero line) is the general level of the seafloor outside of the channel cates older strata. In contrast, depositional channels are bounded by well-developed channel-margin levees that, with time, become progressively elevated above the surrounding seafloor. Mixed channels show a combination of levee deposition and channel-axis erosion, and, in which, the channel floor may lie above, or below, the level of the adjacent seafloor. Channels can also be classified on the degree of confinement of the channel, which typically changes systematically downflow (e.g., Posamentier and Kolla, 2003). Highly confined channels are those in which flow is contained mostly within the

10 10 ARNOTT (see Normark and Carlson, 2003), whereas the examples of ancient canyons described so far are typically much smaller, reaching only slightly more than 1 km deep and 10 km wide. The fill of a submarine canyon is typically stratigraphically complex and lithologically variable. In part, this relates to temporal and spatial differences in sediment source, which varies from local canyon-wall collapse, to an up-dip feeder system sampling a shelf and/or continental sediment source. Fills dominated by local wall collapse are typically finegrained with common mud-rich massmovement (slump and slide) and debris-flow deposits (Fig. 11A, B); an example is provided by the subsurface late Paleocene Yoakum and Lavaca canyon fills (Galloway et al., 1991). Fine-grained sediment deposited from suspension may also occur as a drape infilling part or all of the canyon relief. Coarse sandstone and conglomerate occur as isolated elements in mud-rich fills, but also dominate some fills, especially those formed in tectonically active areas (Fig. 11C). In most cases, coarse sediment occurs as thick-bedded, structureless beds deposited by high-concentration turbidity currents (Fig. 11D). Figure 10. Nomenclature for deep-marine channels (after Pickering et al., 1995). channel. Generally, these channels occur in the upflow parts of a turbidite system and include submarine canyons and erosional channels, the dimensions of which are commonly of the order of kilometers to several kilometers wide and >100 m deep. Farther downflow, these channels change into leveed-channel systems, where channel dimensions range up to several to a few kilometers wide and generally up to about 100 m deep (i.e., the relief between the thalweg and the adjacent levee crests). Here, flows and, in particular, the upper parts of flows, can escape the confines of the channel to build channelbounding levees. These systems, in turn, are succeeded downflow by a complex of poorly confined channels associated with a lobate sedimentary body, commonly termed a depositional lobe or fan, at the downstream end of the channel. Submarine Canyons Submarine canyons are the primary conduits for sediment transport into the deep sea. In the modern oceans, submarine canyons range up to more than 2.5 km deep and 100 km wide Erosional Channels Although the smaller scale erosional channels down-system from submarine canyons also owe their existence to erosion, their fill is different,, and typically consists of a significant proportion of sand- and gravel-rich strata deposited by turbidity currents and other frictional flows. Although the geological literature is replete with erosional channel-fill models (e.g., Beaubouef et al., 1999; Mayall and Stewart, 2000; Samuel et al., 2003), the stages common to most models include: channel inception, sediment bypass, channel filling and channel abandonment. Channel inception is marked by a period of successive flows with high transport efficiency that scour-out a throughgoing topographical feature that serves as the conduit for later flows. Sediment transported during this stage is carried further basinward and deposited in more distal areas. This stage is then succeeded by the channelbypass stage wherein flows range between complete bypass (with no further erosion) to incomplete bypass. Incomplete bypass is commonly indicated by the deposition of laterally discontinuous beds (due to erosion by subsequent flows), intercalation of coarse- and fine-grained deposits, the common patchy occurrence of tractional sedimentary structures, especially dune cross-stratification and coarse-grained lags, and thin drapes of fine-grained sediment, a heterogeneous assemblage of lithofacies herein termed the bypass facies. The next stage, channel fill, is characterized by a change toward flows that have a lower transport efficiency that initiate sediment deposition within the channel, eventually filling part, or all, of the channel. Later, as a result of a diversion of flow at a point upstream (avulsion), or diminution of flow for other reasons (e.g., sea-level rise), the channel system is abandoned and becomes a site of mostly fine-grained deposition that drapes any residual topography. Superimposed on this idealized succession of events are episodes of reactivation, particularly during the channel-filling stage, which

11 12. DEEP-MARINE SEDIMENTS AND SEDIMENTARY SYSTEMS 11 Figure 11. A, B) Submarine-canyon strata consisting of contorted mudstone-rich slump deposits most probably sourced from local canyon-wall collapse (Pennsylvanian Jackfork Formation, Arkansas). C) 135-m-deep canyon incised into thinbedded and very thin-bedded continental-slope turbidites. Canyon fill consists mostly of graded, structureless coarsegrained sandstone and conglomerate. D) (Hector Formation, Lake Louise, Alberta). serve to temporarily rejuvenate the system but not reverse its long-term filling trend (e.g., Samuel et al., 2003). Erosional channels are commonly reported from seismic images, and much less commonly from the ancient outcrop record. This disparity may be the result of two factors: 1. many seismically resolved examples of erosional channels are very large, commonly measuring more than 100 m deep (thick) and a kilometer to many kilometers wide (e.g., Mayall and Stewart, 2000; Deptuck et al., 2003; Abreu et al., 2003) and therefore are on a scale significantly larger than most outcrops; and 2. if strata inside and outside the channel margin are of similar lithology, or are poorly exposed, recognizing the channel-bounding surface in outcrop may be difficult, even though the channel fill and the surrounding deposits may have distinctly different acoustical properties and therefore are easily differentiated on seismic. Seismic images also show that erosional channels are commonly bordered by well developed levees. Deptuck et al. (2003) recognized two end-member kinds of erosional channels and their related levee deposits. Large-scale channels and their related outer levees represent the master erosional channel that is typically several kilometers wide and >100 m deep (Fig. 12). These features were infilled, partly or completely, by smaller scale channels with their associated inner levees, and are described in the next section. Heterogeneous strata deposited by incompletely bypassed flows form the basal unit of the channel fill (Fig. 13). This is commonly succeeded by contorted seismic reflectors interpreted to represent slump/slide (massmovement) and debris-flow deposits produced by collapse of the channel margins. These strata are overlain by a thick succession of sandstone or, less commonly, conglomerate that forms a tabular or sheetlike unit comprising amalgamated small-scale channel fills. During this early stage of channel fill, these smaller scale channels tend to be poorly confined with low to moderate sinuosity and a high width-to-depth ratio (see leveed channels below). Upward, these channels are succeeded by channels with higher sinuosity and a lower

12 12 ARNOTT Figure 12. Uninterpreted and interpreted seismic profile of a channel-levee system in the Indus Fan, Bay of Bengal (after Deptuck et al., 2003). Outer levees bound master channels that are several kilometers wide, which are filled with the deposits of smaller channels (HARs) and their associated inner levees. width-to-depth ratio (see later: confined sinuous channels). Furthermore, these channels commonly show an upward increase in angle of climb, reflecting an increased rate of channel aggradation relative to the rate of lateral migration (Peakall et al., 2000). Eventually, the entire channel system, which at this point has been either partly or completely filled, is abandoned and blanketed by a layer of thin-bedded turbidites and hemipelagic suspension deposits! Leveed Channels Leveed channels are more commonly reported from outcrops because of their smaller size compared to large erosional channels,. Although smaller, leveed-channel fills still range up to a few kilometers wide and 100 m thick (Fig. 14). In addition to occurring as an independent channel element, leveed channels are also important stratal components in larger scale erosional channel fills as described above. Leveed channels typically have a sinuous planform (Fig. 14A, B) with levees of varying development along their margins. Based on the degree of confinement, two end-member types of leveed channels are recognized: poorly confined and highly confined. In both types, channel-fill strata terminate abruptly along an erosion surface defining the outer-bend margin of the channel. The strata, cut by the erosion surface, are either genetically related levee deposits or strata related to an older channel. Along the inner-bend side, however, channel-fill strata either grade continuously into levee deposits in poorly confined channels, or onlap them in highly confined channels. Also, owing to the effects of (fluid) inertia and the resulting tendency of a current to continue along a straight line while the channel floor bends beneath it, levees and levee deposits are always best developed along the outer bend of all channels. Poorly Confined Leveed Channels Channel deposits The base of poorly confined leveed channels is commonly asymmetric, with a steeper margin along one side (analogous to the cut bank of a sinu- ous fluvial channel; Fig. 15). In addition, the channel base is often characterized by a step-flat morphology, indicating the episodic but systematic step-like lateral migration of the entire channel system (e.g., Eschard et al., 2003; Navarro et al., 2007). This asymmetry is present also in the nature of the relationship between the channel-fill and adjacent levee strata. Along the steep margin, levee strata are either in erosional contact with channel strata (see outer-bend levee deposits below) or are separated by a thin, fine-grained bypass unit (Beaubouef et al., 1999). In contrast, on the opposite side of the channel (analogous to the point bar of a sinuous fluvial channel), channel-fill strata either onlap or grade laterally into levee deposits (see inner-bend levee deposits below). The fill of leveed channels is, in fact, composed of the fill of myriad smaller channels, which, because of extensive amalgamation, are difficult to trace individually in outcrop. Nevertheless, the fill of an individual channel is of the order of a few to several meters thick and tens to, at most, a few hundred meters wide. The 3-D amalgamation of these channels forms a channel unit, which is probably the most readily identified channel succession in outcrop. Channel units range from several meters to a few tens of meters thick and commonly show a well-developed finingand thinning-upward trend. As noted above, their base is typically marked by the presence of coarse sediment and abundant mudstone intraclasts. Also, in many cases, coarse, amalgamated strata grade abruptly laterally into finer, more stratified deposits. Strata in the axial part are generally thick- to very thick-bedded, coarse-tail normally graded or structureless conglomerate or sandstone. Beds are typically amalgamated with variable lateral continuity (ranging from tens to hundreds of meters laterally). These coarse-grained deposits are T a turbidites deposited by gravel- and sand-rich, high-concentration turbulent flows. Toward the margin of channel units, especially those higher in the leveed-channel fill, coarsegrained strata tend to thin rapidly (generally over <100 m), and grade from amalgamated sandstone to less amalgamated, more thin- to thick-

13 12. DEEP-MARINE SEDIMENTS AND SEDIMENTARY SYSTEMS 13 bedded sandstone with intervening mudstone. Sandstone beds consist of T a and T ab turbidites and are intercalated with thin-bedded T cd, T cde Bouma turbidites. The thinning and fining of strata toward the margin of channel units suggest that the higher-energy, axial part of the channel is flanked by lower-energy conditions. Levee deposits Subaqueous levees build upward by the addition of sediment when flows overtop the margins of the channel. Flow overspill and related levee aggradation occurs in three ways: flow stripping, inertial overspill, and continuous overspill. Flow stripping and inertial overspill occur at channel bends, and preferentially along the outer bank. In the case of flow stripping, the upper fine-grained part of the flow becomes separated from the lower, coarse-grained part that remains confined to the channel. Inertial overspill occurs when an energetic flow is unable to follow the sinuous thalweg and runs-up the channel margin, allowing even the Figure 13. Idealized model for the fill of an erosional channel (after Mayall and Stewart, 2000). lower parts of the flow to escape the channel. Continuous overspill takes place where the thickness of the flow exceeds the depth of the channel, leading to loss of the flow above the height of the levee along both sides of the channel. Levee growth along the inner-bend side of the channel and the straight segments between channel bends comes about by continuous overspill. Upon escaping the channel, the flow expands rapidly and collapses, resulting in elevated rates of sedimentation immediately adjacent to the channel, with rapidly decreasing rates of deposition farther from the channel. This lateral variation in average sedimentation rate gives levee deposits their distinctive gull wing geometry on seis- Figure 14. Sinuous leveed channels. A) Low sinuosity, large-scale (width >1 km) (Pleistocene) Einstein Channel in the Gulf of Mexico (modified after Posamentier and Walker, 2006). B) High sinuosity, small-scale (100s m wide) modern Amazon channel-levee system in about 3500 m water depth (from Pirmez et al., 2000). Seismic profiles across multiple (C) and single (D) channel-levee complexes in the Amazon Fan (from Hiscott et al. (1998) and Piper and Normark (2001) respectively). In C, note the organized offset stacking, termed compensational stacking, of successive complexes (i.e., younger channels are located preferentially in the topographic lows between two older systems). Note also the extensive blanket of mass-transport deposits that underlies the uppermost stack of channel-levee complexes.

14 14 ARNOTT Figure 15. Poorly confined leveed-channel deposits. A) Channel fill 1 is a sharp-based, up to about 80 m- thick sandstone/conglomerate unit sharply overlain by a second channel system (Channel fill 2), the base of which is indicated by the dashed orange line. Note the sharp, terraced channel base (solid orange line) along the base of Channel fill 1 that ascends obliquely upward toward the right, and is the result of combined vertical and lateral channel migration. Lateral channel migration and aggradation also causes deposits of Channel 1 to terminate abruptly as they overstep strata of their genetically related outer-bend levee B). Along the opposite (left) side, channel strata fine and thin continuously into strata of the inner-bend levee (Neoproterozoic Isaac Formation, western Canada). (C, D) Sharp, terraced margin along the outer bend of laterally migrating, vertically aggrading channels (Upper Cretaceous Tres Pasos, southern Chile). mic images (Fig. 14C). As a levee aggrades, the relief between the channel floor and levee crest increases, eventually allowing only the upper, more dilute portion of the throughgoing flows to overspill. This process is responsible for the upward thinning and fining trends associated with many levee deposits. Owing to their fine-grained nature, levee deposits are typically not well exposed, although notable exceptions exist. Along the outer-bend side of the channel, levee deposits tend to be thick and sand-rich, whereas those on the inner-bend side are significantly finer and thinner. Paleocurrents, typically measured from the c-division of turbidites, are generally oriented very oblique to paleoflow in the main channel. Closest to the channel and along the outer bend, (the proximal levee facies) strata consist mostly of thin- to medium-bedded, fine to mediumgrained sandstone T bc turbidites interstratified with thin-bedded, very fineto fine-grained sandstone/siltstone T cde turbidites and common thick T a sandstones (Fig. 16A). In many modern and interpreted ancient proximallevee deposits, small-scale (ripple) cross-lamination commonly shows evidence of vertical aggradation (i.e., climbing) because of sediment fallout from suspension as the flow expands, with the angle of climb generally decreasing away from the channel margin. However, the presence of climbing ripples is not a universal feature of levee deposits as shown by the Windermere Supergroup (Navarro et al., 2007). This may be due to a somewhat coarser grain size and/or poorer sorting, but the cause remains uncertain. With increasing distance from the channel, the strata become thinner. Typically, thicker beds thin rapidly over 100s of meters whereas thinner beds show little lateral change. As a consequence, strata occurring a few to several 100s of meters from the channel in the distal levee are composed predominantly of thin-bedded, very fine to fine-grained sandstone/ siltstone T cde turbidites (Fig. 16B). The c-division consists of one to at most a few sets of non-climbing ripple crossstratification. Locally, distal levee strata are intercalated with overbank or crevasse-splay deposits (see below). Levee strata on the inner side of channel bends are distinctly thinner and generally finer grained than those on the outer side of the bend. Deposits consist predominantly of fine sandstone and siltstone T cde turbidites, which, locally, are interbedded with thicker sandstone beds. These latter beds consist typically of medium- to thick-bedded, fine to medium sandstone T c turbidites composed of multiple (3 or 4) ripple cross-stratified sets. Like on the opposite side of the channel, strata thin and fine away from the channel, but the rate of change is much greater and reflects the lower-energy nature of overspilling flows on the inner-bend side of the channel. An important difference between inner- and outer-bend levee strata is that, in many places, inner-bend strata are continuous with and grade laterally into channel-fill strata, indicating a continuum between channel and levee deposition. Highly Confined Leveed Channels Channels of this variety are significantly smaller than those described above channel width and depth are of the order of tens to a few hundreds of meters and a few tens of meters, respectively; sinuosity is generally higher too (Fig. 17). In outcrop, highly

15 12. DEEP-MARINE SEDIMENTS AND SEDIMENTARY SYSTEMS 15 confined channels occur mostly as disconnected channels fills that locally are clustered laterally and/or vertically. Clustering is attributed to younger channels exploiting remnant seafloor topography formed by incompletely filled older channels (Fig. 18). In addition, lateral-accretion deposits are well developed, indicating systematic deposition on the inner bend of laterally migrating sinuous channels (Arnott, 2007a; Fig. 19A, B). Channel bases are generally planar and horizontal with only local small-scale (few cm) scours. In places, however, the surface shows a step-like geometry, rising abruptly upward by a few to several meters. The top of the sandy channel fill interfingers with thinly bedded turbidites (Fig. 19A, B). The strata within the channel fill show a distinctive and consistent dip of up to about 7 to 12º, and, like similar features observed in meandering fluvial systems, are interpreted to be lateralaccretion deposits (LADs) formed on the inner bend of a laterally migrating sinuous channel. The fill of the channel consists of lower and upper parts. Beds in the lower part are typically thick- and very thick-bedded, and generally amalgamated (Fig. 20). They consist mostly of sharp-based, graded sandstone and less common granule or fine pebble conglomerate (Fig. 19C). Bases of beds are commonly scoured and, in places, completely erode underlying beds. Mudstone generally occurs as localized patches of intraclasts. Stratigraphically upward, sand-/gravel-rich strata change little in grain size but thin, typically becoming medium bedded; mudstone intraclasts are absent. In the upper part of the channel fill, finegrained strata (mudstone and thinly bedded turbidites) become interstratified with the coarse-grained deposits (Figs. 19A, B). It is noteworthy that coarse-grained beds terminate abruptly up-slope (Figs. 19D, E, 20), which contrasts markedly with the gradual lateral trend observed in poorly confined leveed channels. Also, very near the terminus of each bed, coarse strata consist of a small number of graded, poorly sorted beds capped by planar laminated or dune cross-stratified sandstone. Figure 16. A) Proximal outer-bend levee deposits composed of medium- and thin-bedded turbidites. Thicker beds typically consist of T bcde turbidites that thin rapidly away from the channel margin, whereas thinner beds consist of T cde turbidites that change little in thickness laterally. B) Thin to very thin-bedded T cde and T de turbidites of the distal levee (field notebook for scale). As in the proximal levee, thin beds in the distal levee show little thickness change laterally. Photos A) and B) are from Neoproterozoic Isaac Formation, western Canada. Figure 17. Highly sinuous small-scale (<1 km wide) leveed channel system, the Joshua Channel (Pleistocene), Gulf of Mexico (Posamentier and Kolla, 2003; Posamentier and Walker, 2006). A) Seismic horizon slice illustrating the highly sinuous nature of the channel. Note also the meander cut-offs (oxbows) indicated by arrows. B) Seismic curvature map showing the well-developed levees that bound the channel. Note the common slump scars along the levees on both sides of the channel. Figure 18. Organized channel pattern created by the lateral offset pattern of successive channel-fill elements (images courtesy of Henry Posamentier). Inset sketch from Posamentier and Kolla, (2003). Down-dip from the terminus of each coarse bed, fine-grained strata are truncated as the coarse beds amalgamate (Figs. 19E, 20). Near their termination, fine-grained strata consist of almost complete Bouma-division turbidites (Fig. 19E) that up dip thin, fine, and become dominated by upper-division turbidites (T cde ). The coarse-grained beds in the LADs represent coarse sediment deposition on the lower part of the channel margin. In the LADs finegrained beds fine and thin abruptly

16 16 ARNOTT Figure 19. A, B) Lateral-accretion deposits (LADs) in a highly confined leveed channel from the Neoproterozoic Isaac Formation, western Canada. Note the dipping LADs, which are inclined at an angle of 7-12º, and are interpreted to have accumulated on the inner-bend margin (point bar) of a laterally accreting sinuous channel (from Arnott, 2007a). Also, note the sharp, planar basal contact, but the interfingering of the sandy beds with thin, mud-rich deposits at the top. C) Structureless, graded sandstone near the base of LADs. At their upper end, such coarse beds pinch-out abruptly upward (open arrow in E) into thin-bedded T cde and T de turbidites D), which in turn become truncated downward by the coarse beds (solid arrows in E). Figure 20. Idealized model of lateral-accretion deposits (LADs) formed by a laterally migrating deep-marine sinuous channel. Each coarse and fine LAD consists of several beds and indicates that there were longer term repetitive alternations in the nature of the flows. fills, reflects recurring changes in sediment transport through the channel system. Deposition of the fine beds most likely represents periods of highly efficient turbidity currents that bypass the channel bend and transport much of their coarse sediment load farther downdip. Periodically, these conditions are interrupted by episodes of less efficient turbidity currents that result in deposition of a small number of beds that make up each coarse interval. Currently, the cause for the rhythmic alternation of fine and coarse beds in the LADs is poorly understood, but most likely relates to repetitive changes in local or regional flow and/or channel conditions (Arnott, 2007a). beyond the termini of the coarse LADs, and represent the fine-grained part of the inner-bend levee onto which the coarse-grained LADs onlap. On the opposite, or outer-bend margin of the channel, coarsegrained LADs typically terminate abruptly against fine, thin-bedded turbidites of a slightly older levee system. The rhythmic interstratification of coarse- and fine-grained beds, which is especially well developed and preserved in the upper part of channel Overbank and Crevasse Splays Overbank splays are lobate and sheetlike features formed by energetic flows that overtopped and escaped the channel in an unconfined manner. Crevasse splays, on the other hand, are larger scale lobe-

17 12. DEEP-MARINE SEDIMENTS AND SEDIMENTARY SYSTEMS 17 Figure 21. A) Seismic profile and interpretive sketch of channel-levee complexes in the Amazon Fan (Flood et al., 1991). High amplitude, sheet-like packages, termed HARPs, commonly underlie the high-amplitude reflectors (HARs) that are interpreted to be channel deposits. HARPs are interpreted to be crevassesplay deposits (B) formed during the initial stages of an avulsion, which in many cases are overlain by their genetically related channel (C). shaped features formed immediately downflow of a crevasse channel that is incised into the channel margin and proximal levee. To date, few unequivocal examples of crevasseor overbank-splay deposits have been reported from the geological record. Work on the Amazon Fan (e.g., Flood et al., 1991) identified a distinctive seismic facies that commonly underlies channels (Fig. 21A). Termed a HARP (high-amplitude reflection package) because of its high acoustic impedance, these strata have a sheet-like geometry and are assumed to be sand-rich. These packages are interpreted to represent crevasse splays formed downflow of a breach in the levee of an adjacent active channel (Fig. 21B, C). In the Windermere Supergroup, sand-rich strata interpreted to be crevasse-splay deposits are interstratified with fine-grained, thin-bedded basin-floor and distal-levee deposits (Fig. 22; Arnott 2007b). These crevasse-splay successions, which are up to several meters to a few tens of meters thick, consist of decimeter to several meter-thick units of mediumto thick-bedded structureless sandstone containing common mudstone intraclasts. In places, these strata are interbedded with units composed of upper division (T cde ) sandstone turbidites. Structureless sandstones are poorly sorted and coarse-tail graded with a matrix of fine sandstone, siltstone and mudstone, and are interpreted to have been deposited rapidly by capacity-driven deposition immediately downflow of an area of rapid flow expansion. The thinner graded beds are deposited on the periphery of the collapsing sediment cloud. The intercalation of the two types of beds is related to the lateral wandering of the zone of flow expansion. At its headward end, a crevasse splay is joined to the breach in the adjacent parent channel by a crevasse channel, which forms a sharpbased, distinctly coarser grained unit within a background of fine-grained levee deposits. Crevasse-channel deposits tend to be thin (up to a few meters thick) and consist of amalgamated thick-bedded, massive or normally graded sandstone or (less commonly) conglomerate. Discontinuous beds of single-set-thick dune cross-stratified sandstone are common also. Overbank splays, on the other hand, form from large magnitude, high concentration, coarse-grained turbidity currents that overspill the adjacent channel without confinement (Fig. 22). Their deposits consist of single to multiple beds forming units up to 2 5 m thick. Internally, units consist of amalgamated thickbedded, medium-grained sandstone turbidites that commonly comprise complete Bouma turbidites. Basin-floor deposits In the proximal part of the basin floor, leveed channels terminate downflow in a thick, laterally extensive sediment body variously termed a depositional lobe, distributary-channel complex, sand-sheet deposits, and frontal-splay complex, for it is here that highly confined flows emanating from the leveed channels become unconfined and depositional (Fig. 23). The depositional lobes range up to several tens of meters thick and 100 km wide. Farther basinward, they become finer and thinner, and eventually are replaced by hemipelagic and pelagic deposits. Loss of confinement and the lateral spreading of the flow can be the result of a reduction in slope and/or loss of sufficient fine-grained sediment to build channel-margin levees (Posamentier and Kolla, 2003). Characteristics of the transition from channels to depositional lobes are principally controlled by grain size. In coarse-grained systems, the depositional lobe connects directly with the leveed channel, but in mud-rich systems, the lobe is separated by a transitional zone marked by large scours, sediment waves and sediment mounds. Interpreted transition-zone deposits in the ancient stratigraphic record consist also of an array of lithofacies suggesting both sediment bypass and sediment deposition. Bypass features include shallow scour surfaces draped by fineand/or coarse-grained (lag) sediment that commonly is planar laminated or dune cross-stratified, and units up to about 5 m thick composed of compensationally stacked scour-based lenticular sandstone. Depositional features include sheetlike sandstones that are similar to sheetlike splay deposits described below. Depositional Lobes The planform geometry and size of a depositional lobe depends principally on the sand-mud ratio of the sediment supply. In sand- and gravelrich systems, these features tend to be areally restricted because of rapid deposition, and form discrete elements on the order of a few kilome-

18 18 ARNOTT Figure 22. Aerial photo of strata from the Neoproterozoic Isaac Formation, western Canada showing overbank-splay deposits (OB1.1-OB1.4), and crevasse-splay deposits (CS2.1-CS2-3) and their genetically related crevassechannel fills (C1, C2) note, strata are vertically dipping (see Arnott, 2007b). Overbank-splay deposits occur as units, one to a few beds thick, comprising medium- to thick-bedded, coarse-grained, more complete turbidites (T bcde ) interbedded with thin-bedded, upper-division turbidites (T cde, T de ). This succession is then overlain abruptly by a thick crevasse-splay deposit consisting of three several-meter-thick packages of matrix-rich structureless sandstone (CS2.1 CS2-3) intercalated with few-meter-thick packages composed of classical turbidites. Crevasse-splay deposits are then sharply overlain by a lateralaccreting channel deposit that comprises two separate channel fills (C1, C2). ters wide and a few meters to several tens of meters thick that decrease in thickness and grain size rapidly downflow. Recent work by Deptuck et al. (2007) showed that such late Pleistocene features off the coast of Corsica are of the order of 2 19 km 2 in area, 9 20 m thick, have a length-towidth ratio of <1 2, and show no evidence of progradation but instead backstep upslope. (A similar backstepping pattern is noted in proximal mouth-bar deposits; see Chapter 10). Based on seismic expression and shallow piston cores, strata are dominated by amalgamated sand that represents 50% of the planform area and 75% of the total volume of the depositional lobe. In contrast, more mudrich systems form broader, more sheetlike features that are up to about 10 km wide and become finer and thinner more gradually downflow. Also, their stratigraphic composition is significantly more complex than those in sand-rich systems. Based on seismic interpretation and observations in the rock record, such lobe systems consist commonly of three recurring architectural elements: deep channels, shallow channels, and sheetlike splay deposits (Figs. 23 B, C, D; 24). Deep channels show up to a few tens of meters of incision, with channel margins that are locally steep (few to several meters of relief over a lateral distance of only a few tens of meters; e.g., Meyer and Ross, 2007). These channels are filled with a variety of lithofacies, including heterogeneous assem- blages of coarse- and fine-grained strata related to incomplete bypass, and also thick-bedded, amalgamated sandstone (e.g., Johnson et al., 2001; Fig. 24C). Deep channels are interpreted to represent the principal conduit that supplies sediment to the depositional-lobe complex. Shallow channels occur downflow of the deep channels; these channels show minor relief along their base (scour only a few meters deep over a few hundred meters laterally), and exhibit a consistent internal stratigraphy, which mostly is aggradational. In the channel axis, strata range up to about 5 10 m thick and consist of thick- to very thick-bedded amalgamated sandstone (T a beds), which, when traced laterally over a few hundred meters, show a gradual but systematic change to thick-/medium-bedded complete turbidites that become progressively thinner, finer and less complete turbidites, and eventually pass into single-set thick-, thin-bedded, fine-sandstone T cde turbidites (Fig. 24B). Infilling of a shallow channel is probably because of reduced efficiency caused by deposition farther downflow in the sheetlike splay element, which then is followed by, or is coeval with, the initiation of a new channel and lobe element elsewhere. Outboard of the terminus of the shallow channels, flows become unconfined and highly depositional, forming the sand-rich sheetlike splay element that farther downflow forms a distal apron of thin- and very thin-bedded finegrained turbidites intercalated with hemipelagic and pelagic mudstone. Currently, details of this transition are poorly understood. Nevertheless, in the more proximal sand-rich part of the depositional lobe, strata consist of laterally extensive, tabular units that range from a few meters to a few tens of meters thick, but typically are of the order of 15 m thick (e.g., Meyer and Ross, 2007; Fig. 24D). Typically, the base of each unit is marked by a sharp increase in grain size compared to the underlying strata, which, in many cases, consists of a few meter-thick succession of structureless, mudstone-intraclast-rich, coarse-tail-graded sandstone beds that resemble overbank and crevasse-splay deposits. The sheetlike splay element consists almost entire-

19 12. DEEP-MARINE SEDIMENTS AND SEDIMENTARY SYSTEMS 19 Figure 23. A) Seismic image of a Pleistocene depositional-lobe complex developed at the downflow terminus of a leveed channel, Gulf of Mexico (Posamentier and Kolla, 2003). B) Isopach map of a Pleistocene depositional-lobe complex, Indonesia (Saller et al., 2008). C) Map illustrating the internal stratigraphic complexity of the lobe complex in (B), which consists of 18 (A-R) discrete lobe deposits, which formed during the lowstand systems tract. During the late lowstand to early transgressive systems tract, however, a change to more mud-rich flows caused the leveed channel at the headward end of the lobe complex to extend basinward ( final upper fan channel ), which caused lobe deposition to shift basinward too. D) Detail of a single depositional lobe (Lobe D outlined in red in C) showing that it too is composed of several even smaller, discrete splay elements (labeled 1-6). It is these smaller splay elements that are generally observed in the ancient record and here are termed sheet-like splay elements. ly of amalgamated sandstone that can be mapped over several kilometers. Although difficult to discern because of amalgamation, beds are generally thick to very thick bedded and seemingly form a random pattern of bed-scale cut-and-fill with no organized internal architecture such as compensational stacking of beds. CONTOURITES Although first theorized in the mid- 1930s to exist, the occurrence of contour-following deep-sea currents, or contour currents, was first demonstrated some 30 years later on the continental rise off eastern North America. Deposits of these currents, which are known as contourites, were shown to be characterized by features that distinguish them from better known turbidites, and were later discovered to be the principal constituent in large (tens to hundreds of km long, few tens of kilometers wide, and up to over 1 km high), elongate, slope-parallel sediment bodies termed sediment drifts. These deposits owe their existence to deep-water bottom currents that form part of the global thermohaline or wind-driven circulation system. These semi-permanent currents generally flow parallel to the slope, but locally, especially because of topographical effects, can be diverted obliquely up or down the slope. In polar regions, cold surface water and also more saline water formed by surface-water freezing descends to the basin floor initiating a large-scale (global) flow system. As

20 20 ARNOTT Figure 24. A) Ancient basin-floor deposits (Neoproterozic upper Kaza Group, western Canada) comprising 3 principal depositional elements: shallow channels, deep channels, and sheetlike depositional lobes note strata are vertically dipping. Yellow arrows indicate location of figures B, C, D. In all photos stratigraphic top is to the left. B) Shallow, erosionally based channel with only subtle relief along its base (black dashed line). From the margin toward the axis (direction indicated by red arrow), strata show a rapid increase in sandstone:mudstone ratio as thin-bedded turbidites pass laterally into amalgamated sandstone in the channel axis. C) Deep channel with prominently scoured base (solid black line), overlain by a heterogeneous bypass facies. D) Sheetlike depositional lobes of the order of 15 m thick that form laterally extensive bodies composed of amalgamated sandstone (hammer (circled) for scale). Base and top of lobe indicated by double-headed arrow. of muddy to silty to sandy contourites overlain by a unit that fines upward to muddy contourites (Gao et al., 1998; Stow et al., 2002). This upward change suggests a systematic temporal change in flow speed and/or sediment supply, and which recent evidence suggests occurs on time scales that closely parallel Milankovitch periodicities, suggesting a relationship between orbital forcing of climate and changes in bottom-current velocity (Stow et al., 2002). Where turbidites and contourites coexist, they may be difficult to differentiate. However in an interpreted Plio- Pleistocene turbidite contourite succession, Ito (1996) suggested that contourites can be identified based on minor inverse grading within the ripple cross-stratified unit, in addition to intercalated layers or drapes of mud within ripple cross-stratified sandstone/siltstone units (Fig. 25B). Moreover, contourites commonly contain internal erosion surfaces, typically lack an ordered vertical succession of features like a classical turbidite, and, where interstratified with turbidites, are bounded sharply on their base and top. Collectively these features indicate fluctuating bottom-current speed, and, in many cases, the oscillation between traction and suspension sedimentation. a consequence of the rotation of the Earth and the Coriolis effect, these currents, which have speeds of about 1 2 cm/s, become deflected toward the western side of ocean basins. There, they are constrained by the continental slope and their speed increases to about cm/s and, where locally constricted, can reach speeds of over 2 m/s. However, even at their generally lower speeds, contour currents are approximately at the threshold velocity for very fine and fine-grained sand, and hence are an important sediment-transporting agent in the deep sea. The characteristics of contourites are controlled largely by sediment supply (Stow et al., 2002). They can be composed of terrigeneous clastic, volcaniclastic, or carbonate sediment, and grain size can range from mud to sand and admixtures of both, although mud and silt are most common. In addition, gravel particles occur, but are restricted to local areas of high energy and attendant seafloor reworking and winnowing. (Note that in most cases gravel is brought to the area by glacial ice rafting). Sorting is generally moderate to good, except in areas with low current speeds and slow sedimentation where bioturbation has mixed the deposit. Moreover, bioturbation, which generally is dominated by forms of the Nereites inchnofacies (see Chapter 3), can be moderate to intense, and can destroy primary sediment layering. Traction structures are common and, depending on flow speed and grain size, include small-scale (current ripple) and large-scale (dune) cross-stratification, and also scour marks. Paleocurrents are commonly, but not strictly, parallel to the slope. Contourites often occur as composite units cm thick (Fig. 25A). Most successions consist of a basal upward-coarsening interval consisting SEQUENCE STRATIGRAPHY OF DEEP-WATER DEPOSITS Early sequence-stratigraphic models for the deep-marine siliciclastic deposits were based on wide passive continental margins with a well-developed shelf slope break. Along such margins, the supply of continentderived sediment into deep water depends on the state of the continental shelf, which, in large part, is controlled by the position of relative sea level. During highstand when the shelf is wide, clastic sediment is sequestered in marginal-marine and continental settings and sediment flux into the deep sea is much reduced. Lowstand, on the other hand, and especially when rivers reach the shelf slope break, is a time of voluminous sediment supply and active deposition in deep water. It has been pointed out, however, that in situations where the continental shelf is narrow, for example along the coast