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Ancient river systems in the Himalayan foredeep, Chinji Village area, northern Pakistan
Sedimentary Geology, 88 (1993) 1-76
1
Elsevier Science Publishers B.V., Amsterdam
Ancient river systems in the Himalayan foredeep, Chinji Village area, northern Pakistan Brian Willis
*
Department of Geological Sciences, State University of New York at Binghamton, N.Y. 13902-6000, USA Received December 17, 1991; revised version accepted February 3, 1993
ABSTRACT
The Miocene Chinji and Nagri Formations (Siwalik Group) of northern Pakistan record ancient fluvial environments in the Himalayan foredeep basin. Excellent exposures on the Potwar Plateau (Chinji Village area) allowed detailed documentation of the geometry and stacking of sediment bodies that comprise these strata, and of variation of large-scale bedding geometry, grainsize, sedimentary structures and paleocurrent orientations within such bodies. Major sandstone bodies are tens of meters thick and are continuous along strike for many kilometers. They are composed internally of interconnected channel belt deposits, each of which contain several storeys (channel bar and fill deposits). Individual storeys, defined by inclined bedsets dipping down to a major basal erosion surface at up to 11°, are only 5 to 15 m thick within the Chinji Formation but can be up to 30 m thick within the Nagri Formation. Bedsets within storeys reflect sediment accreted during individual flood events. Along-strike variation of bedsets within storeys and the stacking patterns of storeys within channel belt deposits reflect changes in the channel bed through time due to the growth of bars, migration of channels, and channel cutoff. Generally, braided channel patterns are indicated by the relatively large number of storeys within individual channel belt deposits exposed perpendicular to paleoflow, abundant evidence for channel bar superposition due to channel switching, local evidence for mid-channel bars, dominance of coarse-grained channel fills, and low paleocurrent variations. Paleochannel reconstructions from storeys exposed within the Chinji Formation indicate that individual channel segments generally had widths of 80-200 m, maximum depths of 4-13 m, wavelengths of 1.6-2 km and discharges of 400-800 m3/s. Full channel belt widths (1-2 kin) estimated from exposures perpendicular to palcoflow, and evidence for 2-3 coeval channels within channel belts, indicate full channel discharges of 1500-2000 m3/s. Larger channel segments reconstructed from the Nagri Formation had widths of 200-400 m, maximum depths of 15-30 m, wavelengths of 3-5 kin, and discharges of 3000-5000 m3/s. As Nagri Formation channel systems were also clearly braided, full channel dimensions and discharge estimates are probably at least a factor of two greater than for individual channel segments (i.e. order of 10,000 m3/s). Strata between major sandstone bodies are dominated by lobate and wedge-shaped bodies (crevasse splay and levee deposits), minor channel-form bodies (deposits of minor floodplain channels), laminated mudstone bodies (lake deposits) and paleosols. These strata are arranged into meters to 10 m thick stratified sequences that were rapidly deposited, bounded by well developed paleosols recording periods when deposition rates were low. The thickness and grainsize of such paleosol bounded sequences are not directly related to the proximity of a major channel deposit along strike. Instead, such sequences within overbank deposits appear to reflect local progradation of splays and levees into low areas on the floodplain.
* Present address: Bureau of Economic Geology, University of Texas at Austin, University Station, Box X, Austin, TX 78713-7508, USA. 0037-0738/93/$06.00 © 1993 - Elsevier Science Publishers B.V. All rights reserved
2
13.WILLIS
\ Fig. 1. Location map showing area studied.
Introduction
The Miocene Siwalik Group of the IndoPakistan subcontinent provides a long record of fluvial deposition adjacent to the rising Himalayan Mountains (Figs. 1, 2). Understanding the evolution of the fluvial systems that flowed across the Himalayan foredeep basin requires detailed interpretations of these deposits. Spec-
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Fig. 2. Lithostratigraphy of the Siwalik Group exposed on the Potwar Plateau, Punjab, northern Pakistan. Age of formation boundaries based on paleomagnetic data of Johnson et al. (1985) and the paleomagnetic time scale of Berggren et al.
(1985).
tacular exposures in the Chinji Village area on the Potwar Plateau of northern Pakistan extend for tens of kilometers along strike and for kilometers upsection. These exposures allowed detailed documentation of: (1)the dimensions of sediment bodies within these strata; (2) the nature of stacking, lateral truncation and lateral thickness variations of bodies along strike; and (3) spatial variations in bedding geometry, grainsize, sedimentary structures and paleocurrent orientations within bodies. From these data, ancient channels were reconstructed in terms of their dimensions, channel pattern, bar geometry, hydraulics and migration behavior. Ancient floodplains were reconstructed in terms of their physiography, local depositional environments and varying depositional rates. The ultimate objective of this study is to understand how variations in these deposits are related to changing paleoenvironments and the evolution of the foredeep basin. Stratigraphic studies completed on the Potwar Plateau over the past few decades provide a framework for more detailed sedimentological study. Detailed lithostratigraphy (Lewis, 1937; Gill, 1951; Fatmi, 1973; Pilbeam et al., 1979; Johnson et al., 1982), biostratigraphy (Barry et al., 1982, 1985) and chronostratigraphy based on magnetic polarity (Keller et al., 1977; Barndt et al., 1978; Visser and Johnson, 1978; Opdyke et al., 1979, 1982; Tauxe and Opdyke, 1982; Johnson et al., 1982, 1985, 1988; Burbank et al., 1986; and o t h e r s ) have been established in several areas of the Potwar Plateau, allowing the timing of lithologic and faunal changes, and depositional rates,
ANCIENT RIVER SYSTEMS IN THE HIMALAYAN FOREDEEP
to be compared across the basin. Stratigraphic studies demonstrate that deposits filling the Himalayan foredeep basin varied greatly in both space and time. In particular, formations tend to thin and fine to the south and east, their boundaries can be diachronous by millions of years, and deposition rates varied by at least a factor of two between different formations (Gill, 1951; Barry et al., 1985; Johnson et al., 1985; Beck and Burbank, 1987, 1989).
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3
Although stratigraphic studies have allowed fairly detailed interpretations of the relationship between tectonism and basin sedimentation for some stratigraphic intervals within the basin (e.g. Visser and Johnson, 1978; Raynolds and Johnson, 1985; Johnson et al., 1985; Burbank and Raynolds, 1988; Cerveny et al., 1988; Burbank and Beck, 1989; Beck and Burbank, 1989), more detailed interpretation of depositional environments is needed in order to understand relationships between the varying paleoenvironments, evolution of organisms, basin physiography, and extrabasinal controls such as tectonics, eustasy and climate acting on various spatial scales (e.g. Behrensmeyer and Tauxe, 1982; Badgley and Behrensmeyer, 1982; Behrensmeyer, 1987). The ChinjiNagri Formation boundary (Fig. 3) marks a particularly important time of change in the Siwalik Group where deposition rates increase, the succession dramatically coarsens, a major faunal turnover occurs, and there is a dramatic decrease in the abundance of preserved faunal remains (Johnson et al., 1985; Barry et al., 1982, 1985). This paper focuses on documenting variation among deposits along specific stratigraphic intervals within the Chinji and Nagri Formations in order to understand the distribution of depositional environments, and the physical processes controlling deposition within paleochannels and on ancient floodplains. A companion paper (Willis, 1993b) examines larger-scale sediment variations across these formations in order to document the evolution of these fluvial systems during this important change in foredeep basin deposition. The stratigraphic section studied is located along the Gabhir River in the Punjab of northern Pakistan (Fig. 1). This area includes the stratotypes for both the Chinji and Nagri Formations (Fatmi, 1973). Strata here dip to the north at 3° to 18° but are otherwise generally undeformed. The deposits were documented along three different intervals within this 2-km-thick succession (Fig. 3). A map traced from air photos shows positions of measured sections across each stratigraphic interval documented (Fig. 4). Nearly complete east-west oriented exposure for several kilometers along strike at each interval allowed
4
B.WILLIS
construction of bedding diagrams from photo mosaics, vertical logs, and along-strike tracing of beds in the field. Bedding diagrams display the geometry and distribution of sandstone bodies, the geometry of bedsets within sandstone bodies, paleocurrent orientations relative to the outcrop plane, and traces of distinct paleosols. Sedimentary logs provide more detailed information on variations in sedimentary textures, structures, and paleocurrent orientations (relative to north) for specific sections within bedding diagrams. Logs were generally spaced closely enough to document along-strike variations within individual sandstone bodies and mudstone sequences; however, in places vertical cliffs limited access (particularly within the Nagri Formation). A key to symbols on bedding diagrams and sedimentary logs is presented in Fig. 5. No single bedding surface could be assumed perfectly horizontal along strike across an entire stratigraphic interval studied. A datum was established for successive logs measured along each interval by the correlation of several different paleosols, each of which appeared generally horizontal when traced along strike. Successive logs were aligned such that the average vertical dis-
tance of correlated paleosols from the datum was minimized. Thus it is assumed that only local floodplain topography caused paleosol horizons to vary significantly from paleo-horizontal and that down-basin slopes were negligible across distances studied. This method provides a better estimate of how logs align along strike than that based on arbitrarily assigning one traceable surface as horizontal because local relief reflected by any one bedding surface or paleosol is clearly much greater than regional floodplain slopes. Together the three stratigraphic intervals documented provide a sample of upsection sediment variations across the Chinji and Nagri Formations. The lower Chinji Formation has a low proportion of major sandstone bodies relative to mudstone-dominated beds (30%) and relatively low sediment accumulation rates (200 m / M a compacted thickness). The lower Chinji Formation was examined in detail along a 4 km by 60 m thick interval (Fig. 3, Level A, meters 347-407; bedding diagram, Fig. 6; associated sedimentary logs, Fig. 7; selected photos, Fig. 8). Strata near the Chinji-Nagri Formation boundary record a period of dramatically increased sediment accumulation rates (300-500 m / M a compacted thick-
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ANCIENT RIVER SYSTEMS IN THE HIMALAYAN FOREDEEP
ness) that mark the start of a pronounced upsection coarsening within this stratigraphic succession. These strata were also examined over a 4 km by 60 m thick interval (Fig. 3, Level B, meters 865-925; see bedding diagram, Fig. 10; associated sedimentary logs, Fig. 11; selected photos, Fig. 9). Finally the Nagri Formation is characterized by greater proportions of major sandstone bodies (75% of deposits), and continued high overall sediment accumulation rates (300-400 m / M a compacted thickness). Strata within the upper Nagri Formation were documented along a 3 km by 110 m thick interval (Fig. 3, Level C, meters
5
1512-1612; bedding diagram, Fig. 12; associated sedimentary logs, Fig. 13; and selected photos, Fig. 14), and along a second 750 m by 65 m thick interval located immediately below and to the west (left) of the first (Fig. 3, Level C', meters 1462-1517; bedding diagram, Fig. 15; associated sedimentary logs, Fig. 16). The strata can be divided into major gray sandstone bodies, and red to brown mudstonedominated intervals containing thin gray-brown sandstone beds. To facilitate descriptions and interpretations, these two types of strata are addressed separately below.
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6
Major sandstone bodies
Description of major sandstone bodies Most major sandstone bodies continue along strike for the length of exposures (up to tens of kilometers) and vary from 10 to nearly 100 m in thickness. One major sandstone body was documented in the lower Chinji Fm. (Level A; Figs. 6, 7, meters y = 15-35) and at the Chinji-Nagri Fm. boundary (Level B; Figs. 10, 11; generally between meters y = 20-50 but locally thicker), whereas three major sandstone bodies were documented within the upper Nagri Formation (Level C; Figs. 12, 13; roughly between meters y = 0-40, 45-70, and 75-90, respectively). These sandstone bodies are underlain by a major erosion surface and are capped, in most locations, by a paleosol, Sandstone bodies are composed internally of many storeys stacked both along strike and vertically. Pronounced changes in the thickness of major sandstones along strike are generally associated with the termination of storeys aligned across different horizons within the body. Individual storeys vary between 4 and 30 m thick and extend along strike for a few 100 m to over 3 km. They are defined at their base by a major erosion surface and internally are composed of a set of inclined bedsets. Basal erosion surfaces of storeys display meter-scale relief and locally exhibit flute casts, gutter casts and tool marks. Bedset geometries, sedimentary structures, mean grainsize, paleocurrent orientations and the extent of sediment disruption vary in a systematic way both vertically and along strike within storeys (see Fig. 17). Bedding diagrams (Figs. 6, 10, 12) do not show traces of all bedset bounding surfaces observed (e.g. contrast Fig. 17, example 1 with Fig. 12, Storey J; and Fig. 17, example 2 with Fig. 6, Storey O), rather these diagrams present only general trends in bedset dip direction and steepness across each storey. In most storeys, bedsets dip dominantly in a single direction such that they concordantly overly one storey margin and are discordant with the opposite margin (e.g. Fig. 17). Bedsets within such storeys can be divided into .two groups. Those adjacent to the concordant margin are more
B. WILLIS
steeply inclined a n d / o r progressively steepen in dip, whereas those adjacent to the opposite margin progressively decrease in dip and become more concave upwards. Less commonly storeys contain convex upwards bedsets, and in these cases bedsets dip towards both margins of the storey (e.g. Fig. 6, Storey R; Fig. 10, Storey H). Bedsets (sensu Bridge and Diemer, 1983) within storeys are equivalent to lateral accretions of Allen (1983), inclined strata of Thomas et al. (1987), and stratum bounded by third-order surfaces of Miall (1988a). Beds within sets are strata containing similar lithofacies (i.e. similar sedimentary structures and textures) or are associated with individual sedimentary structures (e.g. an individual set of cross-stratification), and are equivalent to strata bounded by first- and second-order surfaces of Miall (1988a). Trends in bedset dip across storeys can be difficult to discern where bounding surface dips are low, and where adjacent bounding surfaces are separated by longer distances along strike (i.e. rarely, within the Nagri Formation, up to several tens of meters). In some locations, dips of bedset bounding surfaces could only be recognized by tracing individual bedding surfaces along strike over hundreds of meters. Furthermore, when outcrops are viewed from a distance, bedset bounding surfaces can occasionally be confused with the bounding surfaces of individual cross-sets, which rarely also extend along strike for several tens of meters (e.g. Fig. 14B). Despite local difficulties, variation of bedset geometry and internal characteristics can be confidentially mapped across most storeys (see Figs. 6, 10, 12). Bedsets are normally composed of coarse to very fine-grained sandstone, although locally decimeter-thick beds of extrabasinal pebbles occur at the base of some bedsets in the Nagri Formation. Bedsets are generally decimeters to several meters thick, and are bounded below by minor erosion surfaces that dip at 1° to 11° relative to the basal erosion surface of the storey. In some cases (particularly within the Nagri Formation) bedsets can be considerably thicker, and rarely they locally comprise the entire thickness of the storey (e.g. Fig. 12, Storey A; Fig. 14B). Bedsets typically fine upwards and show system-
ANCIENT RIVER SYSTEMS IN THE HIMALAYAN FOREDEEP
atic vertical variations in sedimentary structure, Grainsize fining and variation of sedimentary structures vertically across bedsets vary, however, depending on the vertical position within a storey (Figs. 7, 11, 13, 17). For example, bedsets measured lower in a storey generally have intraformational conglomerate at their base, grading upward into large-scale trough cross-stratified (i.e. dm to m thick cross-sets) a n d / o r planar stratified sandstone (Fig. 17, logs a, b and d). Bedsets measured higher in a storey generally decrease in grainsize upward and show an associated change in sedimentary structure from large-scale trough crossstrata (dm thick cross-sets) to cross-lamination (i.e. mm to cm thick cross-sets) or planar strata (Fig. 17, logs c, e and f). The thickness and mean grainsize of individual bedsets, the thickness of individual lithofacies beds within bedsets, and the thickness of individual sets of cross-stratification all tend to decrease upward within storeys. Near the top of storeys, bedsets are commonly dominated by planar stratified, cross-laminated or highly disrupted sandstone, Biogenic structures within major sandstone bodies are generally limited to either the upper parts of storeys, or the fine-grained tops of bedsets lower down within storeys (Figs. 7, 11, 13). Pronounced sediment disruption, obscuring all primary strata, occurs in the upper few meters of all storeys that are not truncated upwards. This represents over a quarter of the thickness of many storeys, but rarely more than half. Sandstone color commonly changes from light gray to pale reddish brown or yellowish gray where bioturbation/sediment disruption becomes pronounced. There are relatively few distinct types of trace fossils within these sandstone bodies, Nearly all are simple, unbranched, tube-shaped burrows of various circular diameters (1 mm to 1 dm, most 1-2 cm). Burrows have unlined walls, and meniscate, massive or mudclast breccia fills (Willis, 1992a). Lateral changes in bedset geometry, mean grainsize, sedimentary structures and paleocurrents across individual storeys vary greatly between different examples, and stacking patterns of storeys within these sandstone bodies is complex (see Figs. 6-14). Therefore, generalized
7
statements summarizing these variations are not useful. Examination of the following attributes in each sandstone body is integral to subsequent interpretations: (1) alignment of storeys across particular horizons within sandstone bodies; (2) occurrence of paleosols within some bodies, or a paleosol capping a body in some locations but truncated by upper storeys of the sandstone body in other locations; (3) truncation relationships between adjacent storeys; (4) along-strike changes in the thickness of individual storeys; (5) alongstrike changes in bedset inclination direction and steepness across each storey; (6) changes in lithofacies comprising bedsets in different parts of each storey; (7) changes in paleocurrent orientations relative to bedset dip trends observed within the plane of the outcrop across individual storeys; and (8) changes in paleocurrent orientations between adjacent storeys (see Figs. 6-14). In the interval documented at the Chinji-Nagri Fm. boundary (Fig. 3, Level B; Figs. 9-11) it is important to note the smaller size of storeys preserved to the west (left) side of the major sandstone body relative to those on the east (right) side. Note also the smaller overall size of storeys documented in the lower Chinji Fm. (Fig. 3, Level A; Figs. 6-8) relative to most examples documented within the upper Nagri (Fig. 3, Level C; Figs. 12-14).
Interpretationof major sandstone bodies These major sandstone bodies record deposition within the migrating channel segments of a large river. Each storey within these bodies represents a channel bar deposit (set of inclined bedsets) and normally an adjacent coarse-grained channel fill (concave upwards set of bedsets). This distinction between bar deposits and channel filling deposits is somewhat subjective because successive bedsets vary gradually across storeys. Thus a storey is the fundamental architectural element comprising these sandstone bodies. Storey margins with relatively low dips and generally concordant overlying bedsets record channel incision and subsequent bar formation and migration. Migration of thalweg areas and cutbanks in the channel produced basal erosion
8
surfaces of storeys. Intraformational breccia above basal erosion surfaces represent eroded cut bank material. Steep storey margins with adjacent channel filling (concave-upward to horizontal) bedsets record channel cutbanks, Individual bedsets represent deposits of a single flood event, and bounding surfaces of bedsets record a cross-section through a channel bar or fill. Bedset thicknesses reflect distances the bar and channel migrated, a n d / o r the amount of channel filling that occurred during individual deposition events. Normally only a few meters to a few tens of meters of lateral accretion (i.e. accretion in horizontal directions) parallel to the outcrop plane is indicated. Rarely in the Nagri Fm., however, bedsets record over 100 m of lateral accretion during an individual flood event (e.g. Fig. 14B). Along-strike variation of sedimentary textures and structures within individual bedsets reflect the distribution of grainsize and bedforms across a traverse of a channel bar or fill. Dominance of large-scale trough cross-strata within bedsets lower in storeys indicate most bars were covered by sinuous-crested dunes in the deeper parts of channels. Abundant planar strata and crosslamination in bedsets near the tops of storeys (particularly in the Chinji Fm.) reflect the abundance of upper stage plane beds in shallow bartop areas and ripples within areas of slower flow. Vertical variation of grainsize and sedimentary structures within a bedset reflects varying flow velocities during individual floods. Minor erosion surfaces at the base of bedsets reflect local scouring and modification of bar surface geometry during initial or high flood stages. Upward fining within bedsets and the common transition from large-scale trough cross-strata to planar strata or cross-lamination indicate decreasing flow velocities a n d / o r depths as the flood waned. Bioturbated bedset tops record disruption of the sediment by burrowing organisms and, in the upper parts of storeys, plant roots. Intense disruption and root casts within the upper quarter or so of storeys reflect bioturbation of bar deposits when emergent during low flow periods or following channel abandonment (Stanley and Fagerstrom, 1974; Bridge et al., 1986). If associated with low
B. WILLIS
flow periods, a relatively large decrease in channel flow depth between major floods is suggested. Variations in grainsize and sedimentary structures within bedsets traced upwards across a storey are comparable to those inferred from bar migration within many sandy, low to moderate sinuosity, single-channel and braided rivers (cf. Harms et al., 1963; Sarkar and Basumallick, 1968; Bernard et al., 1970; Bluck, 1971; Singh, 1972; Shelton and Noble, 1974; Jackson, 1976; Schwartz, 1978; Bridge and Jarvis, 1982; Bridge et al., 1986; Singh and Bhardwaj, 1991; also cf. review of Bridge, 1985). Many storeys in the Chinji Fm., however, contain more planar strata than is reported in most of these modern rivers (an exception is Harms et al., 1963). Exceptionally largescale sets of planar cross-strata, associated in modern rivers with chute, tributary or alternate bars, were not observed within the sandstone bodies documented (cf. Smith 1971; Cant and Walker, 1976; Cant, 1978). However, a few major sandstone bodies observed in other locations within the lower Chinji Formation locally contain a thick (3-4 m) cross-set at their base that: constitutes a quarter of the thickness of the storey; contains abundant intraformational breccia; and continues along strike for at least 100 m. More detailed study of such bodies is needed to determine whether such cross-sets simply represent migration of a large dune-scale bedform or whether they record deposition on larger-scale channel bed features.
Lateral variations across storeys and storey stacking patterns Each bedset exposed in an outcrop plane shows the geometry of a channel bar as viewed along a specific two-dimensional cross-section. Alongstrike variations in bedset geometry, internal lithofacies, and paleocurrent orientations across individual storeys reflect the interaction between: (1) temporal and spatial variations of bar geometry and facies; (2) channel bar migration, which controls facies preservation and the three-dimensional geometry of the storey; and (3)the outcrop orientation (see also Bluck, 1971, 1976; Jackson, 1976, 1978; Cant, 1978; Collinson, 1978; Bridge
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21
B
®
Fig. 8. Photos associated with Level A bedding diagram. (A) Minor channel-form sandstone body (base marked by dashed line) overlying the major sandstone body (Fig. 6; right side of Storey A). Scale 10 m. (B) Inclined bedsets (two examples arrowed) within storey decrease in inclination toward a coarse-grained channel-fill (Fig. 6; Storey O). Scale 5 m. (C) Complex coarse-grained channel-fill (base marked by dashed line) consisting of smaller channel-forms with inclined bedding surfaces (small arrow marks base of one channel form within larger scale channel fill). A layer of mudstone locally splits the sandstone body where the thick storey terminates (arrow to the right; dotted line marks start of mudstone beds) (see Fig. 6; left side of Storey P). Scale is 5 m. (D) The right side of a storey with convex upward bedding surfaces. Three small arrows mark one bedset that is inclined to the right relative to the base of the storey (see Fig. 6; right side of Storey R). Paleocurrents are generally out of the outcrop. Scale is 5 m long. (E) Lateral truncation of storeys within the major sandstone body (see Fig. 6; truncation of Storey S by T). Bedsets to the left are generally horizontal and disconformable with the inclined erosion surface (one example marked by dotted line). Bedsets to the right are inclined at a low angle to the right and are generally conformable with the inclined erosion surface (small arrow marks one such inclined bedset).
22
B.WILLIS
O
Fig. 9. Photos associated with Level B bedding diagram. (A) Sandstone body displaying three, vertically stacked storeys (see Fig. 10, Storeys B-D). Storeys are defined by major erosion surfaces (arrows and dashed lines) which bound sets of low angle inclined beds (e.g. beds within the middle storey are inclined to the right). Scale is 5 m. (B) Two-storey sandstone body (dashed lines mark the base of each storey; see Fig. 10, Storeys I and O between meters 750-850). Lower storey displays beds inclined to the left which are truncated upward at the erosional base of the overlying storey (arrow). Only the lower part of the upper storey is exposed within the cliff face. Scale is 5 m. (C) Complex vertical and lateral stacking of Storeys H, J and K within the sandstone body (dashed lines and small arrows mark the base of each storey; see Fig. 10, meters 1300-1700). Lower storey to the left shows inclined bedsets dipping to the left (2 examples marked by the letter b). Scale is 5 m. (D) Lateral truncation of Storey H by the steep margin of Storey O (Fig. 10, meters 1800-2100). Note inclined bedding surfaces dipping to the right within the thinner storey (arrows) and beds disconformable with the storey, margin in the thicker storey. Also note small mudstone-filled channel cutting top of the smaller channel sandstone. Scale is 10 m.
A N C I E N T R I V E R S YS T E MS 1N T H E HIMALAYAN F O R E D E E P
and Jarvis, 1982; Bridge et al., 1986; Willis, 1989, 1992, 1993a). Interpretations of bedset geometry within storeys depend on relating the orientation of these two-dimensional cross-sectional views (i.e. recorded by successive bedsets within a storey) to the changing three-dimensional geometry and migration pattern of an ancient channel bar surface (see discussions in Willis, 1989, 1993a). Such interpretations are analogous to reconstructing the shape of a three-dimensional object from a set of serial cross-sections (Fig. 18). However, here such reconstructions are more complicated because: (1) the bar changes shape over time, and thus along-strike variations of successive bedsets exposed in an outcrop reflect both temporal changes in bar shape and the changing position of the bar relative to the outcrop plane; (2) distances of bar migration vary during different depositional events, and thus successive bedsets do not provide an evenly spaced set of crosssections along the channel bar; (3) deposits from some areas on the bar may be removed by erosion as the bar and channel migrates, and thus a complete record of changing bar shape and position may not be preserved, Although individual outcrop exposures of channel bar deposits do not provide the complete information needed to reconstruct changes in bar geometry and position over time, fairly specific interpretations can be advanced by comparing outcrop variations to modern analogues. Specifically, across storey changes in the geometry, lithofacies, and paleocurrents of successive bedsets need to be directly compared to: (1) variations of channel bed topography, facies, and flow orientations observed in planview across modern river channels; and (2) records of how such channel beds change as bars grow and channels migrate. There are relatively few field studies that report spatial variations of channel bed topography and facies in the detail that is required for such comparisons; however, examples of varying utility include both studies of single-channel rivers (e.g. Fisk, 1944; Smith, 1971; Bluck, 1971, 1976; Jackson, 1975, 1976; Bridge and Jarvis, 1 9 8 2 ; Crowley, 1983; Ferguson and Werritty, 1983; Dietrich and Smith, 1983, 1984) and of multichannel rivers (Cant and Walker, 1976; Cant,
23
1978; Bridge et al., 1986; Bridge and Gabel, 1992). Only Bridge et al. (1986) attempted to combine detailed surveys of the river bed with information about how the channel bed changed through time to predict the characteristics of bedding and facies in vertical cross-sections. Unfortunately there are no studies that document bedding geometry and lithofacies variations in deposits of modern river channels with known bed geometry and migration history. As a starting point, interpretations advanced in this study assume that planviews of ancient channel segments were consistent with sine-generated curves of Langbein and Leopold (1966), that paleocurrent orientations generally record down-stream orientations of the channel planform at the point measured, and that channel segments progressively increased in sinuosity and migrated downstream as deposits formed. The trace of the outcrop plane relative to planform traces, and amounts the channel segment increased in sinuosity and migrated downstream are estimated by considering: (1) length of the storey relative to channel width; (2) vertical trends of lithofacies measured at several locations along each storey; (3) trends in bedset dip direction and steepness within the outcrop plain; and (4) paleocurrent variations across the storey relative to the outcrop plain (cf. Willis, 1989, 1993a). Initial assumptions used here may not be useful for deposits formed in channels with highly asyrnmetrical planforms, or for deposits of bars with steep downstream dipping facies (e.g. "down-stream accreted" deposits of Miall, 1988a). However, the basic methodology of interpreting outcrop exposures in terms of sequential cross-sections recording channel migration and bar growth remains the same regardless of channel planform or bar geometry (e.g. see channel bar facies model of Bridge et al., 1986; paleochannel reconstructions of Todd and Went, 1991, and Willis, 1992, 1993a; and Kocurek et al., 1992 for similar reconstructions of eolian dunes). The basic assertion here is that variations in bedset dip, lithofacies and paleocurrents across a storey provide more information about bar geometry and migration than that provided by any one inclined bedset viewed alone. Vertical and along-strike stacking of storeys
24
within sandstone bodies may reflect: (1)stacking of bars and channels in a single or multichannel river associated with channel migration, switching and cutoff within a channel belt; or (2) superposition of different channel belts (Figs. 19-21). Understanding how bedset geometry and facies vary within storeys and how storeys are stacked is fundamental to interpretations of channel sandstone bodies (e.g. Allen, 1983; Friend, 1983; Ramos and Sopena, 1983; Bridge and Gordon, 1985; Miall 1985, 1988a, b; Ramos et al., 1986; and others; also see many excellent examples in Miall and Tyler, 1992). Several lines of evidence suggest that vertical stacking of storeys within the sandstone bodies documented is due mainly to superposition of different channel belt deposits (cf. Figs. 6, 10, 12-19). Nearly complete preservation of most lower storeys is indicated by similar thickness and facies of upper complete storeys and lower truncated storeys (Figs. 6, 10, 12). Complete preservation of lower storeys is evidence against superposition of deposits formed in channel segments with different erosional depths in a single channel belt (cf. Figs. 19B, 19C). Relatively long periods of time separating vertically stacked storeys is indicated by local preservation of pedogenically modified sandstone a n d / o r overbank mudstones overlying some lower storeys within sandstone bodies (e.g. Fig. 6, Storeys K and L; Fig. 8D), or a paleosol that caps a sandstone body in some locations but is truncated along strike by storeys higher within the body in other locations (e.g. Fig. 10, terminating margins of Storeys B and O). Basal erosion surfaces and top surfaces of storeys do not rise more than a few meters along strike (Figs. 6, 10, 12). This suggests only a few meters of vertical accretion within channel belts and is evidence against the high vertical aggradation rates that would be needed to preserve nearly 10 m thick portions of underlying storeys by lateral bar migration (cf. Fig. 19A). Because most major sandstone bodies within the Chinji and Nagri Formations appear to be composed of multiple channel belt deposits separated from vertically adjacent sandstone bodies by thick intervals dorainated by mudstone, they record the concentra-
B. WILLIS
tion in time of different channel belts on the floodplain (Fig. 3; Willis, 1993b). Bedset geometries, lithofacies and paleocurrent variations within storeys, and lateral stacking patterns of storeys within each sandstone body, vary greatly, thus individual examples are interpreted separately below.
Detailed interpretation of major sandstone body in LevelA Sedimentologic variations observed in this sandstone body of the lower Chinji Fm. (Fig. 3, meters 365-380) are presented in Figs. 6-8. Preservation of pedogenically disrupted mudstone beds within this sandstone body (e.g. Fig. 6, above Storeys K and L) suggests that the six storeys aligned across the top of the sandstone body (Storeys E-J, centered on meters y = 30) were deposited in a different channel belt than those lower within the sandstone body (cf. Fig. 19D). Several other storeys locally cut into the paleosol that caps this sandstone body (i.e. Storeys A, B and D) and these storeys undoubtedly postdate all others within this body. Because these storeys above the capping paleosol are not connected to others along strike, they are discussed along with minor channel sandstone bodies (below). Storey L underlies the paleosol in this sandstone body (Fig. 6, between meters x = 100 to 1200, y = 15 to 20), whereas the other storeys in the lower part of the sandstone body truncate this paleosol. The well developed paleosol capping Storey L reflects a period of minimal deposition a n d / o r erosion on bar top areas for on the order of 103-104 years (Behrensmeyer, 1989; Willis, 1993b). Because such long periods of channel bar stability within channel belts seem unlikely (cf. bars in modern Himalayan foredeep basin rivers described by Coleman, 1969 and Bristow, 1987) and because paleocurrent orientations within Storey L are generally outcrop-parallel in contrast to dominantly outcrop-normal paleocurrent orientations within all other storeys lower within the sandstone body, it is most probable that the storey (L) represents a different channel belt that
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ANCIENT
RIVER SYSTEMS
IN T H E H I M A L A Y A N
FOREDEEP
was deposited well before the other storeys within this sandstone body. The other storeys exposed within the lower part of this major sandstone body can be placed into three groups, each separated along strike by the paleosol, and each group interpreted as a different channel belt (i.e. K, M-O, and P-U). Given the similar stratigraphic position of these three channel belt deposits and their similar mean paleocurrent orientations (generally outcrop-normal), they may record local anastomosing of the river system as is locally observed within modern Himalayan foredeep rivers. However, this hypothesis can not be tested within this single outcrop exposure, The topmost channel belt deposit (Storeys E-J) is exposed generally along its downstream a x i s . The truncation of Storey F by E suggests bar superposition by downstream migration (cf. Fig. 20E). All other storey truncations suggest adjacent bar deposits superimposed by channel switching or cutoff (cf. Figs. 20F, 20G). Storey F was clearly deposited first, followed successively by those to the right. Storey E may also have been deposited after Storey F, or it may have been coeval. The along-strike extent of each storey exposed parallel to paleoflow direction provides a minimum estimate of downstream bar length (0.5-1.5 km; cf. Fig. 20). The dominance of downstream dipping bedsets within all storeys in this upper channel belt records the preservation of only downstream ends of channel bars and the dominance of downstream bar migration (cf. Fig. 20). The fining of successive bedsets upwards within storeys agrees with sediment variations observed on the downstream ends of modern river bars (e.g. Jackson, 1976, 1978; Bridge, 1 9 8 5 ; Bridge et al., 1986). Bedset dips (2-4 °) indicate bar surfaces were inclined at low angles in the downstream direction (evidence for the steepness of bar surface dips in the upstream direction has been removed by erosion). In some storeys, bedsets are initially somewhat steeper (5-7 °) and they gradually decrease in dip along strike toward the channel fill (e.g. Storey J). This suggests downstream dips of bar surfaces decreased slightly as the bar grew and the channel migrated. This may be related to
43
a gradual increase in the channel width-depth ratio and a gradual loss of channel discharge over time. The along-strike extent of individual bedsets (100-300 m) records bar and channel widths measured highly oblique (almost parallel) to mean paleochannel fow. Storey thicknesses (7-13 m) provide an estimate of maximum bankfull channel depth. Channel fills (which contain bedsets that progressively decrease in dip, decrease in mean grainsize, and contain a greater abundance of planar and cross-lamination) record decreased flow depth and velocities as the channel was being abandoned. Since channel filling deposits are generally concordant with and nearly as coarse-grained as bar deposits, and fine-grained (mudstone) channel fills do not occur, ends of abandoned channel segments probably did not become plugged with sediment as channels filled (cf. Bridge et al., 1986). Instead, filling of a channel segment must have been associated with a very gradual diversion of discharge into an adjacent, contemporaneous channel. Paleocurrent variations within storeys (30-90 °) suggest sinuosities of at least 1.05 to 1.3 for individual channel segments. Mean paleocurrent directions of different storeys suggests an average sinuosity for this paleo-river reach between 1.1 and 1.2. Storey L is also exposed generally parallel to paleoflow. The upstream dipping bedsets formed by deposition on the upstream end of a bar, and suggest that channel migration was dominated by increasing channel sinuosity with little downstream bar translation (Figs. 20a, 20d). The downstream length of this bar had to be greater than the 1.1 km length of the storey because much of the downstream accreted part of this storey was removed by erosion associated with deposition of the downstream adjacent storey. Bedset dips suggest that bar surfaces were inclined at low angles in the upstream direction (downstream dipping bar surfaces were not preserved). Thickening of the storey and drop of the basal erosion surface to the left can be related to increasing maximum channel depth (from about 5 to 8 m) as the channel bend increased in sinuosity (see Bridge et al., 1986; Willis, 1989). Similarly, slightly coarser-grained bedsets dominated by large-scale trough cross-strata lower in the storey
44
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METERS Fig. 15. Bedding diagram documenting major sandstone bodies exposed in Level C' (stratigraphically just below Level C) and adjacent mudstone-dominated successions. See Fig. 5 for explanation of symbols. The bedding diagram is presented without vertical exaggeration (upper box), and with 4 × vertical exaggeration (lower box).
to the left are probably related to deeper flow depths. Fining-upward grainsize trends across this storey contrast with the coarsening-upward deposits expected in upstream ends of channel bars (cf. Jackson, 1976; Bridge, 1977, 1978; and others),
However, most of the fine-grained bedsets higher in the storey to the left are clearly channel filling, and thus the vertical log probably does not reflect the normal distribution of grainsizes across the channel bar. Small channel forms capping the
Fig. 14. Photos associated with Level C bedding diagram. (A) Two 20 meter-thick sandstone bodies (left side of photo) become connected to the right to form a single 40 meter-thick sandstone body (see Fig. 12, Storeys A, C and G between meters x = 0 to 50). The sandstone bodies become connected where a third (intervening) storey starts (see arrow; dashed lines mark the base of each storey, dotted lines mark the top of storeys that are not truncated upwards). Scale is 10 m. (B) Sandstone body displaying inclined bedsets dipping to the right (see Fig. 12, Storey A between meters 800 and 900). Bedset bounding surfaces (large arrows) are spaced 80 m apart along strike (apparent shortening of this distance resulted from telephoto lens). Within this thick bedset, cross-set bounding surfaces rise gradually to the right (two examples marked by small arrows). Scale is 10 m. (C) Two-storey sandstone body (bases marked by dashed lines; Fig. 12, Storeys I and K between meters 400 and 500). Basal erosion surface of upper storey (arrow) rises laterally, more fully exposing lower storey. Note single large cross-set at the base of the upper storey (just above arrow). Scale is 10 m. (D) Multi-storey sandstone body (base of each storey marked with dashed lines). Middle storey starts just to the left of this photo and thickens towards the right truncating the storey below (large arrow; see Fig. 12, Storeys K and J between meters 600 and 700). Inclined bedsets with the middle storey (small arrow marks one example) dipping to the right are generally conformable with downcutting basal erosion surface. Scale is 5 m. (E) Two-storey sandstone body (bases of each storey are marked with dashed lines, tops are marked by dotted lines where they are not truncated upwards; Fig. 12, Storeys H and K between meters 1200 and 1400). Inclined bedsets to the right within the u p p e r storey end at the basal erosion surface (arrow marks one example) and successive bedsets decrease in dip into the coarse-grained channel fill (to the right). T h e upper storey ends just to the right of this photograph, but the lower storey continues for about another kilometer. Scale is 10 m. (F) Termination of major sandstone body in a coarse-grained channel fill (major erosion surfaces marked by dashed lines, dotted line marks top of the sandstone body; see Fig. 12, Storey K between meters 2500 and 2700). Note the smaller-scale channel-form (arrow) capping the larger-scale channel-filling sequence. Scale is 5 m.
'
•
46
8. WILLIS
larger channel fill record deposition within smaller channels (1-2 dm deep, meters wide) during the final stages of filling. Along-strike paleocurrent variations (SO”) indicate a channel segment sinuosity of at least 1.2. All of the other storeys lower within the sandstone sheet are exposed generally perpendicular to paleoflow. Storey K (far left) was deposited by a 7 m deep, 120 m wide channel, based on its
I
2
Fig. 16. Sedimentological
3
logs associated
4
thickness and 5” bedset dips. Storeys M and N are dominantly channel filhng, suggesting they were relatively short-lived and did not migrate to any great extent. Storey thicknesses of M and N, and bedset dips suggest that channel segments had maximum depths of at least 3 m and widths of about 60 m. Opposite dip directions of bedsets within these two storeys suggest channel cutoff (with Storey N deposited first) or two coeval
6
5
with Level C’ bedding
diagram.
7
See Fig. 5 for explanation
8
of symbols
ANCIENTRIVERSYSTEMSIN THEHIMALAYANFOREDEEP
47
braid channels (cf. Figs. 21B, 21C). In either case, flow diverged into these two channels and they developed sinuosity in an opposite sense, Storey O contains bedsets that dip in a similar direction to those in Storey N. The truncation of Storey N by O may record deposits from different channels superimposed due to channel switching (Fig. 21D). Alternatively, this bedding truncation may record the migration of a single bar that changed shape between two successive depositional events (related to varying discharge within the adjacent channel). Steeply inclined bedsets (11 °) adjacent to the left margin of Storey O traverse a distance of about 50 m along strike indicating that the channel had a width of about 75 m and a maximum depth of at least 8 m. Decrease of successive bedset dips along strike indicate progressive decrease in channel depth due to gradual channel abandonment as the bar migrated, but also progressively more oblique cross-sections across the channel. Most channel filling bedsets are planar stratified and relatively fine grained; however, a few bedsets are coarser
Variations Example I
0
5'o
grained and large-scale trough cross-stratified (e.g. Fig. 7, log 15, meters 20-23), and probably record erosional scouring of the nearly abandoned channel segment during unusually high discharge floods. Paleocurrent variations within inclined bedsets suggest that the channel segment developed a sinuosity of at least 1.1. Within the channel belt defined by Storeys P - U , Storey Q was deposited first on a bar that migrated to the right within the outcrop plane. Storey R was then deposited. The convex-upward bedding surfaces within Storey R indicate two coeval channels separated by a mid-channel bar (cf. Fig. 21A). Diverging paleocurrent directions and vertical trends in grainsize indicate a crosssection through the upstream end of the bar (cf. Bridge et al., 1986). Right-dipping bedsets and storey thickness on the right suggest that the channel segment became at least 9 m deep and at least 150 m wide. Truncation of Storeys Q and R by Storey P indicates either channel switching (cf. Fig. 21C) or a mid-channel-bar with a somewhat complex history of across-stream accretion (cf.
in bedset geometry
1~)0
150
and lithofacies
2bo
~so
36o
35o
Example 2
0 Scale
25 50 75 in meters (2x v e r t i c a l a
b
c
100 125 exaggeration d
e
150
f
Fig. 17. Bedset geometry and lithofacieswithin two storeys (stippled). Thick lines are major erosion surfaces, thin lines are minor erosion surfaces, dotted lines are gradational contacts with overlyingmudstone beds. Characteristic vertical lithofacies variations across bedsets (each decimeters to meters thick) in different parts of each storey are shown by letters on bedding diagrams and sedimentary logs (a-jr).
48
B. WILLIS
Fig. 21B). The total vertical extent of Storey P and the lateral extent of individual bedsets within this storey suggest a maximum channel depth of about 12 m and a width of about 200 m. T h e r e is little evidence of lateral accretion prior to channel filling. Smaller-scale channel forms within the larger channel fill signal reduced discharge and deposition on smaller bars and channels that were several meters deep and a few tens of meters wide (cf. Bridge et al., 1986; Bristow, 1987). In Storeys S and T, bedset dips (initially 7 °) progressively decrease and bedsets become concave upward adjacent to right hand storey margins, indicating periods of lateral bar migration followed by periods of channel filling. Because left hand storey margins and internal inclined bedsets have similar dip directions, they may be related to changes in bar shape due to discharge variations between depositional events. Thus
Storeys S to T could record deposition on the same migrating bar. Given downstream length estimates of bars for other storeys within this sandstone body (generally 0.5 to 1 km), deposition of Storeys R - T by migration of the same bar would require a single bar deposit to extend in a direction transverse to paleoflow for over a bar wavelength• Such a relationship would indicate highly sinuous channel segments (greater than 1.8), which is unlikely given the restricted paleocurrent variations observed within this sandstone body. Rather, paleocurrent shifts of 30° across the storey margins separating R, S and T, and channel filling bedsets on the right side of truncated storeys, suggest that these storeys formed on different bars and that deposits were superimposed by channel switching (cf. Fig. 21D). Because most of Storey U was removed by erosion, the timing of its deposition relative to others within this sandstone body is ambiguous•
•
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Cross-section
Fig. 18. A channel bend shown in planview is contoured in equal increments of depth (darker stippling reflects deeper areas). Depth contours follow Hooke (1974), but no specific scale is implied here. A set of sequential cross-sections across the channel bend are shown for 11 parallel traverses superimposedon the same cross-sectionalplane. The set of cross-sectionstogether reflect spatial changes in channel bed geometry. Similarly,inclined bedsets within a storey reflect a sequential set of cross-sectionsthrough a migratingchannel bar.
ANCIENT RIVER SYSTEMS IN THE HIMALAYAN FOREDEEP
A
B
49
B and C, suggest channel maximum depths and widths of 10-12 meters and 150-225 m, respectively. Low paleocurrent variations within storeys (less than 30°), but high paleocurrent variation
D
Fig. 19. Schematic examples of storeys vertically stacked within sandstone bodies. Thin lines show bedding surfaces, thick lines show major erosion surfaces, arrows show paleocurrent orientations relative to the outcrop plane (up is out of the outcrop). (A) Vertical superposition of storeys within a channel belt produced by rapid vertical accretion and downstream bar migration. The maximum thickness of the lower truncated storey records the amount of vertical accretion that occurred as the channel migrated by one bar length (see also Bridge, 1975). (B) Vertical superposition of storeys within a channel belt related to channels of varying size and thus erosional depth. (C) Vertical superposition of storeys within a channel belt related to stacking deposits from equal-sized channels where the former channel, that eroded to a greater depth, had a higher sinuosity than the later channel. A similar case would occur for channel bends of equal sinuosity where the outcrop plane exposes the two channel bends across different parts of their planform (see also Willis, 1989). (D) Vertical superposition of storeys from two different channel belt deposits by avulsion and subsequent erosion of the older channel belt by channels in the younger channel belt.
Detailed interpretation of major sandstone body in Level B This sandstone body is located at the ChinjiNagri boundary (Fig. 3, meters 885-908). As in Level A, vertical stacking of storeys is interpreted as largely due to vertical superposition of differe n t channel belt deposits. Storeys A - C each record the edge of a different, vertically stacked, channel belt deposit. Storeys D, E and H define a
fourth channel belt deposit. Storeys I-K, along the base of the sandstone body o v e r m o s t of the left side, constitute a fifth channel belt deposit, The thick storeys on the right side (O-Q) comprise a sixth channel belt deposit. Finally, Storey R constitutes a seventh channel belt deposit, The thicknesses (10-12 m) and the lateral extent of inclined bedsets (100-150 m) in Storeys A,
~ P ' ~ ~ ~ ~ ~" ~ . , d
~ ~ ; ~ " ~ e '
~ - ' ~ ] ~ ~ " 4 [ l l l m l ~
~ i ~ ~ . , ~ ~
" e z - - ~ ~ - - ~ ~ f "
~
-'~Ap/
~ g._~./_.j~_~:.~~ h ~
Fig. 20. Schematic examples of storeys stacked laterally and parallel to mean paleoflow within a channel belt. Each exampie contains a map showing channel planforms preserved through successive migration steps, and a cross-section showing trends in bedding geometry. Paleocurrent directions within planform maps are from left to right (see small arrow). Paleocurrents within cross-sections are given by arrows (up is into the cross-section plane). The larger context of channel segments and bars is not shown in these diagrams. Thus, in each case diagrams might represent a side (point) bar growing into the displayed channel segment, or the half of a mid-stream (braid) bar that is growing into the displayed channel segment. There is assumed to be no net vertical aggradation within the channel belt. (A-C) Along-stream stacking of bars deposited on opposite sides of a channel segment. Bar growth is by sinuosity increase (A), sinuosity increase and downstream bar migration (B), or downstream bar migration alone (C). In each case adjacent bar deposits display similar trends in bedding geometry, but paleocurrents vary in an opposite sense. Note that in (B) and (C) the upstream channel-filling storey margin truncates the downstream adjacent storey. (D and E) Along-stream stacking of bars deposited on the same side of a channel segment. Bar growth is by sinuosity increase (D), or downstream bar migration (E). In each case adjacent deposits display similar trends in bedding geometry and paleocurrents vary in the same sense. Note that in (E) the upstream channel-filling storey margin truncates the down-
stream adjacent storey.(F-H) Along-stream stacking of bars produced by channel cutoff. After cutoff the channel segment can develop sinuosity in an opposite sense (F), or the same sense (G, H) compared with sinuosity development before the cutoff. Note that in both cases the depositional margin of the downstream storey truncates the upstream adjacent storey.
50
B. WILLIS
b e t w e e n d i f f e r e n t c h a n n e l b e l t deposits, suggest
d o w n s t r e a m d i p p i n g b e d s e t s suggest that both
that c h a n n e l s w i t h i n c h a n n e l belts had low sinu-
bars were d e p o s i t e d o n the same side of a chan-
osity, b u t d i f f e r e n t c h a n n e l belts could traverse
nel s e g m e n t (as in the c e n t e r part of Fig. 20D)
this position o n the f l o o d p l a i n f r o m very d i f f e r e n t
a n d that the c h a n n e l s e g m e n t h a d a sinuosity of
directions,
at least 1.17. Storey thickness a n d the along-strike
Storeys D a n d H r e p r e s e n t a c h a n n e l belt
extent of i n c l i n e d b e d s e t s i n d i c a t e c h a n n e l maxi-
deposit exposed generally p a r a l l e l to paleoflow,
m u m d e p t h s of at least 10 m a n d b a r d o w n s t r e a m
T h e p a t t e r n of b e d s e t dips in Storey D record
l e n g t h s of at least 1.5 km. T h e t h i n n e r c h a n n e l -
d o w n s t r e a m a c c r e t i o n o n a b a r toward the west (left side), a n d u p s t r e a m a c c r e t i o n o n a n alongs t r e a m a d j a c e n t b a r to the east (right side). P r e s e r v a t i o n of u p s t r e a m - a c c r e t e d b e d s e t s a n d the conformity of b e d s e t s b e t w e e n the two b a r deposits suggest that c h a n n e l b a r s did n o t migrate d o w n s t r e a m as they grew a n d the c h a n n e l
increased in sinuosity (cf. Fig. 20D). Nearly horizontal b e d s e t s n e a r the c e n t e r of the storey record filling of a wide c h a n n e l cross-over area as the c h a n n e l was gradually a b a n d o n e d . F i n i n g - u p w a r d deposits in d o w n s t r e a m accreting b a r areas b u t little vertical v a r i a t i o n in grainsize in u p s t r e a m accreted areas agrees with t r e n d s o b s e r v e d in m o d e r n rivers (e.g. Bridge et al., 1986). L a t e r a l paleocurrent variations between upstream and
~ al" a~ ~------~:~J~~z:'~'~'~'~'~'~
"
" A~'~a'°"
bl• ~..~.../.~1~.-~:j:~ ,~~~[~.~..~
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yel /~e2 ~ "
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~
Fig. 21. Schematic examples of storeys stacked laterally and perpendicular to mean paleoflow within a channel belt. Each exampledisplays a map showing channel planforms preserved through successive migration steps, and a cross-section showing trends in bedding geometry and paleocurrents. Paleocurrent directions within planform maps are from left to right (small arrow). Paleocurrents within cross-sections are given by arrows (up is into the cross-sectional plane). (A) Growth of a mid-stream (braid) bar dominantly by sinuosity increase. Bedding surfaces dip in both directions away from the center of the bar deposit into the adjacent channel segments. Rightdipping and left-dipping beds define broad convex-upward surfaces.Paleocurrents are generally out of the cross-sectional plane. Paleocurrents are directed toward the center of the bar deposit if the cross-section is through the downstream end of the bar (a2), and away from the center if the cross-section is through the upstream end of the bar (al). (B) Growth of a mid-stream (braid) bar by dominantly down-stream translation. Deposits generally resemble (A); however, some differences occur. Cross-sections through upstream parts of the bar will contain beds that are disconformable with adjacent channel deposits (bl). Right-dipping and left-dipping beds will be disconformable if the bar migrates oblique to paleoflow or discharge is transferred from one channel segment to the other over time (b2). (C) Across-stream stacking of two channel bars due to cutoff. The channel segment developed sinuosity in an opposite sense before and after cutoff. Deposits are quite similar to those formed by migration of a mid-stream bar (compare with b2). (D) Across-stream stacking of two channel bars due to channel switching where the channel segment developed sinuosity in the same sense before and after cutoff. Beds dip in the same direction within adjacent bar deposits; however, a disconformity of bedding and a shift in paleocurrent orientations are expected at the storey margin. Where this paleocurrent variation is slight, the local bedding disconformity could be mistaken as one produced by slight change in channel alignment, possibly initiated by discharge variations between depositional events. (E, F) Acrossstream stacking of bars formed on the opposite sides of a channel segment. In both cases beds within adjacent storeys dip toward each other. Paleocurrent trends within storeys and whether beds within storeys are conformable with their mutual channel fill depends on the orientation of the cross-section and migration of the bars (compare el, e2 and f).
ANCIENT RIVER SYSTEMS IN THE HIMALAYAN FOREDEEP
filling storey (Fig. 10, Storey E) was probably deposited later by a smaller channel that flowed along a depression associated with the larger, underlying, abandoned channel fill. Interpreting the truncation of Storey D by H is somewhat complex because the outcrop plane steps backwards in a direction perpendicular to that shown in the bedding diagram by about 200 m between logs 18 and 19 (Figs. 10, 11, between meters x = 1700 and 1800). Had the outcrop not stepped backwards, the downstream end of the bar just described would probably have been exposed. Convex-upward bedsets within Storey H, with paleocurrents out of the outcrop, indicate two channel segments that were separated by a mid-channel bar. Inclined bedsets and storey thickness suggest the right channel segment was about 8 m deep and 85 m wide. Here, successive bedsets coarsen upward or have constant mean grainsize. Bedsets dipping towards the channel segment on the left are not as well exposed, but channel depths appear to have been only about 4 m, widths are probably also somewhat less than for the channel segment to the right, and bedsets are finer grained and generally channel filling. Storeys I - K along the base of the sandstone body record a channel belt deposit exposed generally perpendicular to paleoflow, and provide a measure of flow transverse channel belt width of about 1.8 kin. Bedsets within Storey I suggest deposition on a mid-stream bar (cf. Figs. 21A, 21B). The thickness of the storey, and the lateral extent of inclined bedsets to the left, indicate a channel segment with a maximum depth of 13 m and a width of about 200 m. Increased bedset clips and storey thickness toward the channel fill to the left record bar growth and development of higher channel sinuosities (cf. Bridge et al., 1986). Local bedding truncations, reflecting slight variations in the dip of adjacent bedsets, record temporally varying discharge. Coarsening-upward trends suggest a cross-section through the upstream end of the bar (cf. Fig. 21A; Bridge et al., 1986). However, a few coarse-grained bedsets capping the channel fill to the left probably reflect short periods of erosion associated with channel segment reactivation following its nearly complete abandonment. Limited variation in pa-
51
leocurrents along strike within left-dipping bedsets suggest, alternatively, a cross-section through the channel bend apex and that the bar did not migrate to any great extent downstream but grew as the adjacent channel segment increased in sinuosity. Storey J truncates Storey I, and must have formed later. Storeys J and K contain bedsets that dip in opposite directions away from a central location, indicating bar(s) that grew into two different channel segments (cf. Figs. 21B, 21C). Both storeys record channel segments about 10 m deep and 100-150 m wide. Local vertical superposition of these two bar deposits dictate that Storey J formed before Storey K; however, a period when both channel segments coexisted can not be ruled out. Determining whether different channel segments were active at the same time or whether they represent deposits separated in time is a persistent problem within these deposits. In this channel belt (Storeys I-K) it appears that two (or briefly three) channel segments could have been active at one time, suggesting a braiding index up to 2 or 3. The channel belt deposit represented by Storeys O - Q is exposed generally perpendicular to paleoflow and contains storeys that are substantially larger than those exposed to the left. Minor along-strike paleocurrent variations within these storeys suggest low channel segment sinuosities, but individual bar deposits exposed in flow transverse cross-sections do not provide a reliable estimate of sinuosity. The thickness of Storey O, and the along-strike extent of its bedsets, indicate channel depths of 23 m and widths of up to 450 m. The base of this storey varies in elevation along strike. The initial drop in the basal erosion surface records channel incision. The gradual rise in the basal erosion surface along strike to the left over the next kilometer is an opposite trend to that predicted by Bridge et al. (1986) and Willis (1989). This may indicate that there was at least 4 m of vertical aggradation within the channel belt as the bar migrated and grew about a kilometer in a direction perpendicular to paleoflow, or that the channel gradually became shallower over time (i.e. reflecting gradually decreasing discharge). Bedsets with lower dips and coarse grainsizes (sandstone) within the
52
channel fill record gradual abandonment of the channel over many flood events, Storey Q is similar in scale to Storey O but contains bedsets that dip in the opposite direction (cf. Figs. 21B, 21C). Storey Q clearly formed after deposition of Storey O and it may reflect deposition within the same channel segment, that switched position and grew in the opposite direction (cf. Fig. 21C). In Storey P, bedset dips of only 1-3 ° suggest a low channel segment sinuosity or a cross-section through the bend cross-over area. Beds with finer grainsize, lower dips, and cross-lamination within the channel fill indicate relatively slow flows as this channel segment was being abandoned. Its smaller size (less than 17 m deep, 200 m wide) and its position on the underlying bar deposits suggest it records a major chute channel cutting across the bar top. The presence of such a chute channel, and the coarse-grained nature of channel fills within these large storeys, suggest that the river had a braided pattern in general. However, the braiding index can not be estimated because the full extent of this channel belt could not be documented, nor could the total number of storeys aligned across stream be assessed, Along-strike variation in thickness, bedding geometry, and paleocurrent directions of Storey R indicate a single bar deposit exposed parallel to the channel belt axis (cf. Fig. 20A). Such along-strike variations suggest a maximum channel depth of 12 m, a width of 200 m, a downstream bar length of a least 1 km, and a sinuosity of at least 1.18. Because this storey is not associated with others along strike, it is interpreted along with minor channel deposits (below).
Detailed interpretation of major sandstone bodies in Level C Three major sandstone bodies are documented within this interval. A fourth, its lowermost part documented in the topmost portion of the bedding diagram (Fig. 12, meters y = 100 to 110) will not be discussed. One sandstone terminates along strike (Fig. 12, Storeys H - L between meters y = 50 and 70), whereas the others continue for the length of the exposures (3-4 km). The lowest sandstone body within this interval (Fig. 12,
B. WILLIS
Storeys A - G between meters y = 0 and 40) is clearly composed of two different superimposed channel belt deposits that are connected locally by several stratigraphically intermediate channelfilling storeys. Both channel-belt deposits are exposed generally parallel to paleoflow. The truncation of Storey A by B suggests superimposition of downstream adjacent channel bar deposits by channel switching (cf. Figs. 20F, 20G); however, similar bedset dip directions and paleocurrents in these storeys suggest that truncation may alternatively represent simply a bedding disconformity within a single bar deposit. Minor along-strike paleocurrent variations (less than 45 °) suggest that channel sinuosities were low, but were at least 1.05. Convex upward bedsets, and local upstream dipping bedsets, within Storey A indicate preservation of both upstream and downstream ends of a channel bar. This suggests that the channel bar did not migrate significantly downstream, although it clearly grew preferentially in the downstream direction (cf. Fig. 20). The downstream length of the bar was greater than 3 km. Thickness and lateral extent of bedsets indicate channel segment depths of about 22 m and widths of at least several hundred meters. The alongstrike distance separating adjacent inclined bedset bounding surfaces (many tens of meters) indicates that bars migrated or grew in the downstream direction by tens of meters to 100 m during individual flood events (cf. Coleman, 1969; Bristow, 1987). Storey G indicates a channel less than 15 m deep that is exposed parallel to paleoflow. Very low bedset dips make it difficult to document along-strike bedding trends, and inhibit interpretation of this storey in terms of deposition on channel bar(s). However, low bedset dips and low along-strike variation in paleocurrents suggest deposition within a low sinuosity channel segment. Thin channel forms cutting into the top of Storey G (i.e. Fig. 12, meters x = 200 to 800; meters x = 1200 to 1800) record shallow chute channels, or minor overbank channels flowing along the major channel course after it had been abandoned. Inclined bedsets within such channel forms, and along-strike paleocurrent variations suggest that some of these minor channels had
53
ANCIENT RIVER SYSTEMS IN T H E HIMALAYAN FOREDEEP
high sinuosities (at least 1.5). Channel-form storeys that locally connect the upper and lower channel belt deposits (Fig. 12, Storeys C-E) record channels that were active while the larger channel system had avulsed to another area on the floodplain. Their restricted lateral extent suggests they were relatively short-lived, and did not significantly increase in sinuosity nor migrate laterally. Storey F is thicker and may record the edge of a major channel belt that continues further to the east (left). Storey H records a single-storey channel-belt deposit exposed about 45° oblique to paleoflow, The along-strike extent of this storey suggests that the channel belt was less than 1 km wide. Geometry of bedsets within the storey record a cross-section though a mid-channel bar separating two channel segments, the one to the right about 15 m deep and 150 m wide, and the one to the left about 6 m deep and 100 m wide. Low variability in paleocurrent directions suggests low channel segment sinuosities (less than 1.1). The braiding index of this channel belt was 2. The lower part of this sandstone body (represented by Storeys l - L ) records deposition within a different channel belt. Storeys I - K record lateral stacking of channel bar deposits, whereas the vertical truncation of Storey L records vertical stacking of bar deposits within the channel belt (cf. Figs. 19B, 19C). Paleocurrents between Storeys I and K vary by almost 70° suggesting that channel segment sinuosities were locally as high as 1.3. Downstream-dipping bedsets suggest that channel bars migrated some distance downstream as they increased in sinuosity. Truncation of Storey K by Storey I records superposition of adjacent channel bar deposits due to downstream bar migration (cf. Figs. 20A-20E). Because paleocurrents within these storeys progress in a similar way (i.e. outcrop-parallel to outcrop-normal), they must represent deposition on the same side of a channel segment (cf. Fig. 20E). Storey J truncates Storey K only locally, thus it records channel cutoff across the upstream part of Storey I. Gradual thinning of Storey K along strike to the right, and the associated progression of paleocurrents back to outcrop-parallel orientations, suggest that Storey K is exposed along a cross-
section that passes through a cross-over area before crossing the cutbank margin to the right. Thicknesses of Storeys M and N suggest channel maximum depths of about 15 m, and maximum bedset dips of 4-5 ° suggest that channel widths were less than 250 m. The dip of bedsets in opposite directions indicates a highly oblique cross-section through two channel segments that developed sinuosities in the opposite sense (cf. Figs. 21A-21C). It is not possible to assess whether channel segments were coeval braid channels or represent channel switching of a single channel segment. Preservation of upstreamdipping bedsets within the left-hand storey suggests that bar migration was dominantly by expansion (sinuosity increase). Along-strike variations in paleocurrents suggest channel sinuosities between 1.1 and 1.2.
Paleochannel reconstructions Qualitative reconstructions Generally braided channel patterns within channel belts are suggested by the relatively large number of storeys within individual channel belt deposits, limited along-strike extent of individual storeys exposed perpendicular to paleoflow relative to estimates of bar downstream length, dominance of coarse-grained channel fills (i.e. sandstone- as opposed to mudstone-dominated), low paleocurrent variation within individual storeys, local preservation of convex-upwards bedset surfaces recording a mid-stream bar, and abundance evidence for bar deposit superposition due to channel switching. Channel segments were low sinuosity (1.0-1.3), and channel bars dominantly migrated downstream (but not always). Channel cutoff and switching occurred frequently relative to channel migration rates, but individual channel segments were cutoff gradually and abandoned over many flood events. Bars all had similar geometries, characterized by low-angle channel bed dips (2-4 °) in upstream and downstream directions, and slightly steeper (7-11 °) channel bed dips in cross-stream directions. Lithofacies variations across storeys of the Chinji Formation suggest that during floods much of the channel bed was covered by plane beds
54
B. WILLIS
when channels were straighter (and shallower), but dunes dominated channel bed surfaces in deeper parts of more sinuous channel segments, Lithofacies variations across larger storeys within the Nagri Formation suggest that dunes covered a greater proportion of areas within these chartnels. Exposure of the full width of channel belts displayed perpendicular to paleoflow provide a more complete view of channel bed geometry and lithofacies measured across individual ancient rivers. For example, Storeys P - T (Figs. 6, 7) indicate a river that was about 2 km across and contained two active channel segments (i.e. those that deposited Storeys P and T), thus the river had a braiding index of 2 (as defined by Brice, 1964). Bedsets within these two storeys (P and T) suggest that most of this distance consisted of very shallow bar top areas covered dominantly by plane beds, but by ripples during falling flows. Patterns of sediment disruption suggest that much of this bar top area was exposed during low flows, There are significant differences in the size of storeys compared between different channel belts within some sandstone bodies (e.g. Fig. 10), and between sandstone bodies observed along different stratigraphic intervals (cf. Figs. 6 and 1 2 ) . Two distinct scales are recognized, larger ones (15-30 m thick) are restricted to the Nagri Formation (see east (left) side of Fig. 10, Storeys O-Q; and Fig. 12, Storeys A-B, G, I-N). These two distinct scales of channel deposits are interpreted to reflect deposition within two coeval rivers of different magnitude (Willis, 1993b). More detailed estimates of the dimensions and discharge of these two rivers are presented below,
Quantitative reconstructions Paleochannel geometry and hydraulics of several storeys within these sandstone bodies were reconstructed by comparison with a physical model that predicts the sedimentology of singlechannel bar deposits (Bridge, 1977, 1978, 1 9 8 2 ) . Full descriptions of paleochannel reconstruction methods are presented by Bridge (1978), Bridge and Diemer (1983) and Willis (1992, 1993a); thus only a brief summary will be given here. Given certain initial hydraulic and geometric conditions, the model predicts the local flow pattern, bar
topography, and variation in mean grainsize across the bed at channel forming (bankfull) discharge within channel bends of regular planform. Reconstruction of a paleochannel involves modifying the initial conditions of the model until it simulates the observed bar deposit as closely as possible. Many of these initial conditions can be determined directly from the deposits. For instance, the geometry of successive inclined bedset bounding surfaces, dimensions of channel fills, and lateral variations in storey thickness provide an estimate of the varying paleochannel geometry within a single two-dimensional cross-section. In past reconstructions such geometric parameters were transformed by a linear projection into a plane orthogonal to paleoflow before comparisons were made with simulated bar deposits (e.g. Bridge 1978; Gardner, 1983; Bridge and Diemer, 1983; Bridge and Gordon, 1985; and others). This method has been shown to provide unreliable estimates of channel width and bar shape (see Willis, 1989). Here a new computer program is used that allows comparison of outcrop information with simulated bars in any specified crosssection orientation (see also Willis, 1989, 1992, 1993a).Channel bend sinuosity and wavelength are estimated by considering the along-strike extent of the storey and paleocurrent variations within the storey. These estimates are checked against empirical equations relating channel width and depth to channel planform (as in Bridge, 1978). Hydraulic parameters such as mean water surface slope and friction coefficients are estimated from variations in grainsize, sedimentary structures and channel depth. Once initial estimates are obtained, repeated simulations are performed and estimates of model input parameters are refined to improve the fit of model output to data. Precise error limits associated with this methodology are difficult to estimate. Errors are associated with differences between real and modeled three-dimensional bar geometries, and inherent uncertainties associated with using depth/grainsize/sedimentary-structure relationships to estimate flow velocities. Reconstructions of the same channel deposit using maximum and
ANCIENTRIVERSYSTEMSIN THE HIMALAYANFOREDEEP
55
minimum reasonable input estimates generally produced discharge estimates that varied by less than 25% of the total value. This bar model was developed for single-channel meandering rivers (i.e. point bars) and it has been tested successfully against both modern rivers (Bridge and Jarvis, 1982; Bridge, 1984a) and ancient deposits (Bridge and Diemer, 1983; Bridge and Gordon, 1985; Willis, 1992, 1993a). The model also appears to provide adequate agreement with single curved channel segments in braided rivers where bars do not have steep down-stream dipping faces (Bridge and Gabel, 1992). Significantly, bedding geometries predicted
by this model are also broadly similar to those observed within the deposits documented here (cf. Figs. 6, 10 and 12 with Willis, 1989). It must be emphasized, however, that reconstructions of multi-channel rivers will provide only an estimate of conditions within individual channel segments. Therefore, dimensions of the entire river channel and details of its hydraulics can only be obtained if the channel braiding index can be determined. Tables 1-3 summarize estimates of paleochannel geometry and hydraulics reconstructed from storeys in each of the three stratigraphic intervals documented. Two distinct scales of channel deposits are recognized within these deposits (as
TABLE 1 Paleochannel reconstructions, Level A (obstacle sandstone) Storey number
1
2
3
Apparent bar width (m)
110
100
130
Estimated channel bankfull width (m)
160
140
Mean depth (m)
4.5
Maximum depth (m)
4
5
6
7
8
45
135
120
95
105
200
70
200
180
140
160
5.0
4.5
4.0
6.5
5.5
5.0
7.0
7.5
8.5
8.5
8.0
11
10.5
10
13
Centerline channel length in one bend (m)
1110
1010
1470
490
1300
1360
1160
1190
Centerline sinuosity
1.1
1.1
1.2
1.15
1.1
1.2
1.3
1.2
Mean Darcy Weisbach type friction coefficient
.04
.06
.06
.06
.06
.06
.06
.06
Dynamic friction coefficient
0.6
0.6
0.6
0.6
0.6
0.6
0.6
0.6
Mean friction slope (ram/m)
.059
.063
.055
.064
.051
.050
.052
.049
Mean velocity (m/s)
0.59
0.63
0.57
0.59
0.62
0.61
0.62
0.64
Bankfull discharge (m3/s)
430
440
510
170
810
610
430
710
Position of storeys within Fig. 6. 1: Storey E, meters x = 0-600, y = 20-30.2: Storey K, meters x = 0-250, y = 15-22.3: Storey L, meters x = 110-1200, y = 12-21.4: Storey O, meters x = 860-1140, y = 17-25.5: Storey P, meters x = 1170-1520, y = 17-29.6: Storey R, meters x = 1800-2140, y = 18.5-30. 7: Storey T, meters x = 2020-2550, y = 18-27.5.8: Storey J, meters x = 2550-4000, y = 18-31.
56
B. WILLIS
discussed above, and in Willis, 1993b), channel dimensions and flow reconstructed from these two different scale deposits are discussed separately below (reconstructions of larger channels are marked by an asterisk in Tables 2 and 3). Smaller channels were reconstructed from the lower Chinji Fm. (Fig. 3, Level A; Table 1), the west (left) side of the interval documented at the Chinji-Nagri boundary (Fig. 3, Level B; Table 2) and Storey H documented within the upper Nagri Formation (Fig. 3, Level C; Table 3). Estimated channel dimensions and flow parameters reconstructed from these smaller-scale deposits are: (1) depths 7.5-13 m; (2) widths 70 to 200 m; (3) downstream lengths of individual curved channel
segments 0.5 to 1.5 kin; (4) sinuosities of individual curved channel segments 1.0-1.3; (5) downstream flow velocities 0.60-0.80 m / s ; (6) downstream slopes 0.05 and 0.08 m m / m ; and (7) discharge of individual channel segments 150 and 1050 m 3 / s (although most channel segments had discharges between 500 and 800 m3/s). Full river discharges associated with these smaller-scale channel deposits can be estimated by adding discharges of all potentially coeval channels within an individual channel belt (this assumes only one channel belt was active, i.e. the system was not anastomosing). Three channel belts provide such estimates: (1) Storeys P - T in the Lower Chinji Formation (Fig. 6) indicate a flow-transverse
TABLE 2 Paleochannel reconstructions, Level B (GB 1 sandstone) Storey number
1
2
3
4
Apparent bar width (m)
140
130
90
100
Estimated channel bankfull width (m)
205
195
140
Mean depth (m)
6.0
7.0
M a x i m u m depth (m)
12
Centerline channel length in one bend (m)
5
6
7"
8
55
105
280
140
150
80
160
400
205
7.0
4.5
4.0
5.0
12
6.5
13
13
9
7.5
10
23
12
1430
1590
1360
1110
550
1220
2900
1480
Centerline sinuosity
1.1
1.3
1.25
1.15
1.1
1.25
1,1
1.18
Mean Darcy Weisbach type friction coefficient
.04
.06
.04
,04
.06
.06
.06
.04
Dynamic friction coefficient
0.6
0.6
0.6
0.6
0.6
0.6
0.6
0.6
Mean friction slope (ram/m)
.05
.05
.05
.06
.08
.06
.03
.04
Mean velocity (m/s)
0.59
0.76
0.70
0.65
0.60
0.62
0.72
0.71
Bankfull discharge (m3/s)
730
1040
680
440
200
480
3560
950
Position of storeys within Fig. 10. 1: Storey C, meters x = 0-500, y = 31-45.2: Storey I, meters x = 400-600, y = 15-27. 3: Storey I, meters x = 700-900, y = 16-25. 4: Storey J, meters x = 1300-1700, y = 26-35.5: Storey H, meters x = 1600-1950, y = 23-32. 6: Storey K, meters x = 1550-1960, y = 11-21.7: Storey O, meters x = 2000-3000, y = 20-43.8: Storey R, meters x = 2800-3900, y = 40-53.
57
A N C I E N T R I V E R SYSTEMS IN T H E H I M A L A Y A N F O R E D E E P
channel belt width of about 2 km, and two coeval channels with a combined discharge just under 1500 m 3 / s (reconstructions 5 and 7 in Table 1); (2) Storeys I - J at the Chinji-Nagri boundary (Fig. 10) indicate a flow-transverse channel belt width of about 1.8 km, channel braiding index of about 2 (but perhaps temporarily 3), and full channel belt discharge of about 1200 m 3 / s (reconstructions 3 and 4 in Table 2); (3) finally, Storey H in the u p p e r Nagri (Fig. 12) records a braid bar separating two channel segments with a combined width of just over a kilometer (measured 40 ° oblique to paleo-flow) with a combined discharge of just over 1350 m 3 / s (reconstructions 4 and 5 in Table 3).
Larger channel deposits were observed in the right side of the sandstone body documented at the Chinji-Nagri Formation boundary (Fig. 10, Storeys O - Q ) , and within the sandstone bodies documented in the u p p e r Nagri Formation (Fig. 12, Storeys A - B , G, H - N ) . Estimated channel dimensions and flow p a r a m e t e r s reconstructed from these larger-scale deposits are: (1) depths 8-23 m; (2) widths 180-400 m; (3) downstream lengths of individual curved channel segments 1.4 to 2.9 km; (4) sinuosities of individual curved channel segments 1.1-1.2; (5) downstream flow velocities 0.7-0.8 m / s ; (6) downstream slopes 0.03 and 0.06 r a m / m ; and (7) discharge of individual channel segments 1390 and 3560 m 3 / s
TABLE 3 Paleochannel reconstructions, Level C (GB 112 and 13 sandstones) Storey number
1"
2"
3"
4
5
6"
App~u'ent bar width (m)
180
120
230
110
150
230
Estimated channel bankfull width (m)
260
180
340
170
230
340
Mean depth (m)
9.5
10
8.0
3.5
7.5
11
Maximum channel depth
17
19
16
6
15
20
Centerline channel length in one bend (m)
1660
1390
2150
2800
2800
2150
Centerline sinuosity
1.15
1.2
1.15
1.15
1.2
1.1
Mean Darcy Weisbach type friction coefficient
.06
.06
.06
.06
.06
.06
Dynamic friction coefficient
0.6
0.6
0.6
0.6
0.6
0.6
Mean friction slope (mm/m)
.05
.047
.053
.06
.072
.045
Mean velocity (m/s)
0.76
0.76
0.74
0.52
0.62
0.79
Bankfull discharge (m3/s)
1920
1390
2180
310
1050
3390
Position of storeys within Fig. 12. 1: Storey I, meters x = 0-300, y = 54-70. 2: Storey J, meters x = 540-920, y = 50-69. 3: Storey K, meters x = 1300-1700, y = 52-65. 4: Storey H, meters x = 1000-1200, y = 60-75. 5: Storey H, meters x = 800-1400, y = 19-39. 6: Storey A, meters x = 500-700, y = 69-75.
58
B. WILLIS
O
ANCIENT RIVER SYSTEMSIN THE HIMALAYANFOREDEEP
(discharges between 2000 and 3000 m3/s are most common). Exposures are not extensive enough to estimate the full channel-belt widths nor braiding indexes for these large channel deposits, and thus full river discharges can not be estimated in detail. However, sedimentological evidence suggests that these larger channel belts had braided channel patterns (see above) and comparisons to modern large rivers of the Himalayan foredeep basin (Coleman, 1969; Bristow, 1987; Wells and Dorr, 1987; Singh and Bhardwaj, 1991) suggest braiding indexes less than 4 or 5. Thus full channel belt discharges recorded by these thicker channel belt deposits were probably on the order of 10,000 m3/s (i.e. at least a factor of four larger than reconstructed for thinner channel belts described above), Minor sandstone- and mudstone-dominated bodies
Strata separating major sandstone bodies are mudstone-dominated overall, but contain small laterally discontinuous sandstone bodies (Figs. 6-16). These strata can be divided into: (1) minor channel-form bodies, (2) lobate and wedge-shaped bodies, and (3)laminated mudstone bodies. These sediment bodies can be grouped into meters to ten meter thick sequences delineated by alternation of sediments containing primary stratifica-
59
tion with highly disrupted sediments defining paleosols. Sediment bodies and paleosol bounded sequences are described and interpreted separately below.
Minor channel-form bodies Description Most minor channel-form bodies are dominantly sandstone, are generally one to a few meters thick, and are continuous along strike for several tens of meters to less than a kilometer (e.g. Fig. 6, Storeys B, V and others; Fig. 10, Storeys M, N and others; Fig. 15 between meters 10 and 15). A few are thicker, however, and have dimensions similar to individual storeys within major sandstone bodies (e.g. Fig. 6, Storeys C and D; Fig. 10, Storey R). Most bodies are composed of a single storey in the plane of the outcrop, are flat topped with a concave-upward basal erosion surface, and have apparent width to thickness ratios between 10 and 50 (cf. ribbon bodies of Friend, 1983). However, some channel bodies are more extensive along strike and rarely contain two adjacent storeys (e.g. Fig. 6, Storey D). These have greater apparent width to thickness ratios (over 100) and are bounded below by an erosion surface that is generally horizontal for some distance between more steeply dipping margins (cf.
Fig. 22. Minor sandstone bodies. (A) Minor channel-form sandstone body (base marked by dashed line, top by dotted line) displaying inclined bedsets that grade upward to mudstone (arrows mark two examples). Scale is 2 m. Outcrop is located within the lower Chinji Formation (Fig. 3, stratigraphically adjacent to meters 498-502). (B) Minor channel-form sandstone body (base marked by dashed line) with inclined beds which pass into a mudstone-dominated channel fill (boundaries of channel fill marked by dotted lines). Scale is 10 m. Outcrop is located within the lower Nagri Formation (Fig. 3, stratigraphically adjacent to meters 1231-1235). (C) Minor sandstone body (base marked by dashed line) with inclined beds (e.g. arrow) that fill a channel-form preferentially from the right. Scale is 5 m. Outcrop is located within the upper most Chinji Formation (Fig. 3, stratigraphically adjacent to meters 863-866). (D) Minor channel-form body (base marked by dashed line) that contains left dipping bedsets and fines from dominantly sandstone to the right to dominantly mudstone to the left. Scale is 1 m. Outcrop is located within the lower Chinji Formation (Fig. 3, stratigraphically adjacent to meters 340-345). (E) Minor lobate sandstone body (base marked by dashed line, top marked by dotted line) with low angle inclined beds dipping in the direction the wedge thins (to the right). Small arrows mark one bedset that drops on elevation relative to the base of the wedge-shaped body about 2 m over 50 m. Scale is 5 m. Outcrop is located within the lower Chinji Formation (Fig. 3, stratigraphically adjacent to about meters 1497-1501). Figure 24 documents sediment variations along this sandstone body in more detail. (F) Wedge-shaped sandstone body that passes laterally into a major sandstone body (dashed lines mark the base of the major sandstone body and the base of the wedge-shaped body). Note inclined bedsets in the wedge-shaped body, dipping in the direction the wedge thins (to the right). Scale is 5 m. Outcrop is located within the upper Nagri Formation (Fig. 3, stratigraphically adjacent to meters 1497-1502). Sediment variations along this sandstone body are documented in Fig. 16, logs 2-6, meters 35-45.
60
B. WILLIS
0
30 _
.
k ; , ~ ' ~ ~
/
2
-~'- ~"-'~'-i. ~ ~:~__,..
c
OMOSISs
i
X"v"
'~-"
_.
,~'~', 1 ,~.,,,,X'v-
I . . . . . . . . . .
e
4
I"
c''''~'r
I ....
5
C
I ....
3
E5
I0
00 METER5
60
120
180
ANCIENT RIVER SYSTEMSIN THE HIMALAYANFOREDEEP
sheet bodies of Friend, 1983). Bodies with greater extent along strike relative to their thickness (i.e. sheets) tend to have higher proportions of inclined bedsets relative to channel-filling (concaveupward to horizontal) bedsets (e.g. Fig. 22A) and more commonly end in mudstone-dominated channel-fills (e.g. Fig. 22B). Minor channel sandstone bodies comprise less than 10% of deposits that separate major sandstone bodies, and across most stratigraphic horizons they are spaced many kilometers apart, Erosion surfaces at the base of these bodies have decimeter-scale relief, and are overlain by intraformational conglomerates of mud clasts, pedogenic nodules and (in places)vertebrate fossils, Oncolites (up to centimeters in diameter and concentrically laminated about a central clast) occur within such basal conglomerates locally in adjacent areas of the Chinji Formation. Cylindrical stromatolites (decimeters to a meter in diameter) were observed along the base of one minor channel body in the Nagri Formation (Behrensmeyer, 1987; Willis, 1992). The base of many channel-form bodies correspond to upper parts of paleosols or to nodular calcite layers within paleosols, Bedsets within minor channel sandstone bodies are fine- to very fine-grained, erosionally based, centimeters to decimeters in thickness, and fine upward (Figs. 7, 11, 13, 16). Lower in a body, bedsets grade upward from intraformational conglomerates into cross-laminated or less commonly large-scale trough cross-stratified
61
sandstone. Higher in a body, bedsets are mainly cross-laminated or intensely disrupted sandstone. Bedsets fine upward to mudstone in the upper parts of some bodies (e.g. Fig. 22A). Rarely, in the Chinji Formation, symmetrical ripple marks cap bedsets. In general, successive vertically superjacent bedsets fine and thin toward the top of the sandstone. Sediment disruption is extensive within most minor channel bodies, and in some cases stratification within the body is completely disrupted (particularly in the Nagri Formation). The disruption of sediment is complex and can not always be related to distinct structures; however, less than 10 mm diameter, tube-shaped burrows are common and are generally similar to those within upper parts of major sandstone bodies. Vertical cracks (up to 1 m in depth) and root casts are also common near the top of these minor bodies and throughout channel filling parts. In bodies with mainly channel filling bedsets (i.e. ribbon-shaped), a few bedsets are inclined concordant with one channel margin but most become concave upward, or horizontal and discordant with the other margin (e.g. Fig. 6, Storeys B and V; Fig. 22C). Paleocurrents in such bodies are generally perpendicular to the outcrop and vary little between bedsets (less than 30°). Sandstone bodies that extend for greater distances along strike, and are similarly exposed generally perpendicular to paleo-flow, display comparable lateral changes in bedset geometry to those just described but are dominated by inclined bedsets adjacent to the bedset-concordant margin that
Fig. 23. Details of a minor channel sandstone body exposed generally parallel to paleoflow. The sandstone body is exposed on a series of extensive bedding surfaces and within adjacent inclined rock faces. The outcrop is located within the upper Chinji Formation (Fig. 3, stratigraphically adjacent to meters 785-800). This sandstone body constitutes paleontology site Y76, which is one of the most prolific vertebrate fossil localities ever discovered in the Siwalik Group (for discussion of taphonomy see Behrensmeyer, 1987). (A) Map of paleocurrent orientations derived from large-scale trough cross-strata and parting lineations on planar strata measured along bedding surfaces near the base of sandstone body. Paleocurrents progress from southerly orientations on the east end of the exposure to westerly orientations on the west end of the exposure (arrows, south to top of page). Circled numbers indicate positions of measured sections. (B) Extensive bedding surfaces exposing edges of individual beds which curve around in a similar way as paleocurrent orientations (A). (C) Sedimentological logs measured at selected locations along exposure (see A). See Fig. 5 for explanation of symbols. Arrows represent paleocurrent orientations (up to the north). (D) Photo to the north of steep slope exposure between sections 4 and 5. Inclined bedding surfaces dip from top of the sandstone to its base. Beds grade to mudstone near the top of the body. Paleocurrents are toward the observer. (E) Schematic bedding diagram of sandstone body with two times vertical exaggeration. Beds dip in both directions away from the center of the sandstone toward laterally adjacent, fine-grained channel fills. Numbers mark positions of measured sections. Arrows display paleocurrent orientations relative to the orientation of the vertical exposure and a line connecting measured sections.
62
13.WILLIS
dip toward the channel filling margin. Individual inclined bedsets normally extend along strike for only a few tens of meters. These more extensive bodies also commonly display higher along-strike paleocurrent variations (up to 70°). Where channel sandstone bodies are exposed generally parallel to paleoflow, bedsets either (1) dip in both directions away from the center of the sandstone body toward marginal channel fills (e.g. Fig. 6, Storey C; Fig. 10, Storey R; see also unusually well exposed example documented in Fig. 23), or (2) are dominated by bedsets that dip toward a marginal channel fill at the down-stream end of the body and are truncated by a marginal channel fill at the upstream end of the body (e.g. Fig. 6, Storey D). Such bodies generally show large along-strike paleocurrent variations (commonly 100_150o). The character of sediments within minor channel bodies varies somewhat compared between examples observed within the Chinji and the Nagri Formations. In the Nagri Formation, most examples contain bedsets entirely of sandstone that are completely disrupted throughout the body. In the Chinji Formation, most contain some bedsets that fine upwards to mudstone; disruption of stratification becomes pervasive only in the upper half or so of the body; bodies showing extensive variation of paleocurrents (i.e. greater than 90 °) along strike are more common; some examples contain oncolites, and vertebrate fossil remains are more common,
Interpretation Most minor channel-form sandstone bodies are interpreted as deposits of crevasse, minor tributary, or floodbasin drainage channels (cf. Fisk, 1944, 1947; Smith et al., 1989; Tye and Coleman, 1989). However, where the thicknesses of these bodies are similar to those of individual storeys within major sandstone bodies (e.g. Fig. 6, Storeys B, C, D, V and others; Fig. 10, Storey R), interpretations are more problematical. Such exampies exposed parallel to paleoflow (e.g. Fig. 6, Storeys C and D; Fig. 10, Storey R) may record single-channel segments within major channel belts cut tangentially by the outcrop plane; however, this interpretation can not be tested within
single outcrop exposures. Thick channel bodies exposed perpendicular to paleoflow (e.g. Fig. 6, Storeys B and V; Fig. 10, Storeys M and N) that are not connected to a major channel belt deposit along strike, must be related to either large crevasse channels or to major tributary channels. The low proportion of minor channel sandstones within these deposits (less than 10%) suggests a paucity of overbank channels, and that these minor channels were commonly spaced at least kilometers apart on the floodplain. Nearly all channel bodies are composed of a single storey in the outcrop plane suggesting that they were deposited by single-channel river systerns. These channels were generally a few meters to 10 m deep, and their apparent channel widths varied from 10 m to less than 100 m. Most of these channels have generally SSE paleocurrent directions that are somewhat oblique to those within subjacent major sandstone bodies, suggesting that they may have been flowing gradually toward or away from a major channel belt. Many channel deposits are directly underlain by carbonate-rich paleosol horizons or laminated claystones suggesting that these sediments were more resistant to erosion (cf. Gibling and Rust, 1990). Bedsets (decimeters to centimeters thick) within bodies clearly indicate that deposition on bars and in adjacent fills occurred gradually over many flood events. Thin bedsets, dominance of cross-lamination, and fine to very fine grainsizes within most bodies suggest generally slower flows and lower deposition rates than within major channel systems. Mudstones capping some bedsets suggest very slow channel flows between flood events. Rare bedsets capped by symmetrical rippie marks in the Chinji Fm. indicate that sediments were locally reworked by waves between depositional events. Oncolites observed within the Chinji Fro. must have formed along basal erosion surfaces of channels (as they are never observed within adjacent deposits), and record the ponding of water and low deposition rates within channel fills between flood events (Schafer and Stapf, 1978; Nickel, 1983). The singular occurrence of cylindrical stromatolites at the base of a minor channel sandstone within the Nagri Formation similarly reflects ponding of water and the accre-
63
A N C I E N T R IVE R SYSTEMS IN T H E H I M A L A Y A N F O R E D E E P
tion of carbonate around submerged logs or tree trunks (cf. Kraus, 1987b). In general, sandstonedominated channel fills suggest that overbank channels continued to be conduits for flood waters until they where completely filled, whereas bodies with mudstone fills must have been abandoned completely and then gradually filled with suspended sediment deposited from slow moving flood waters. Pervasive sediment disruption by mudcracks and root casts within many bodies suggest that much of the channel area was emergent between flood events. However, restriction of completely disrupted sediments to upper parts of most bodies within the Chinji Fm., suggest that most of these channels contained some flow throughout the year until they began to be abandoned. Given the abundance of vertebrate fossils, much of the intense disruption within these channel deposits may have been due to trampling by large mammals (cf. Wells and Dorr, 1987). However, bioturbation by roots and burrowing organisms, and desiccation of the sediment clearly played an important role in sediment disruption, Trends of bedset inclination, lithofacies, and paleocurrents across bodies can be interpreted
and fill deposits exposed generally parallel to mean downstream direction. The proportion of upstream to downstream dipping bedsets within storeys exposed parallel to paleoflow reflects the distance the bar migrated downstream as it grew (cf. Fig. 6, Storeys C and D, and Fig. 20) and along-strike variations in paleocurrent directions reflect the channel sinuosity (e.g. about 1.3 for Storeys C and D). In many cases differences in minor channel body geometry (i.e. ribbon versus sheets) probably reflect only the orientation of the outcrop plane across the channel deposit. However, sheet-like bodies occur more commonly directly overlying major sandstone bodies, suggesting that moderately higher sinuosity overbank channels may occur along the abandoned courses of major rivers. Perhaps topography along abandoned major channel belts is associated with local areas of very low relief. Because many of the sheet-like minor channel bodies have dimensions similar to individual storeys within major channel belt deposits, some examples may reflect deposits associated with the avulsion of the major river channel (i.e. avulsion belt deposits of Smith et al., 1989).
employing the same methods used to interpret individual storeys in major sandstone bodies (see above; Fig. 18; Willis, 1989). Minor channel sandstone bodies with paleocurrent orientations generally perpendicular to the outcrop, and inclined bedsets filling to one side, record minor channel deposits exposed perpendicular to mean downstream direction. Apparent width to thickness ratios of such bodies (less than 50, i.e. ribbons), low proportion of lateral-accretion bedsets relative to channel-filling bedsets and limited variation in paleocurrent directions, suggest that these channels did not migrate laterally nor increase in sinuosity to any great extent before they became abandoned and filled. This suggests that most channels were relatively short-lived, or that channel banks were stabilized by vegetation or erosion-resistant overbank sediment. Bodies with inclined bedsets dipping in both directions away from the center of the body, have paleocurrent orientations generally parallel to the outcrop (cf. Fig. 20), and are interpreted as single-channel bar
Lobate and wedge-shaped bodies
Description Most deposits between major sandstone bodies occur in lobate or wedge-shaped bodies, several decimeters to several meters in thickness, defined by vertical variations in bedset thickness, mean grainsize and extent of sediment disruption (Figs. 7, 11, 13, 16). These bodies commonly appear to be horizontal beds in outcrop, but they gradually vary in thickness and internal lithofacies when traced along strike. These bodies continue along strike for only a few hundred meters to rarely a few kilometers. Some are dominated by sandstone where they are thickest, whereas many are dominated by mudstone throughout. In general, those dominated by sandstone are thicker and extend for greater distances along strike. Bodies contain decimeter to centimeter thick bedsets internally. Bedset inclinations and along-strike changes in bedset geometry could be recognized
64
B. WILLIS
only in the thickest and coarsest-grained examples (see below), Although correlation of individual lobate/ wedge-shaped bodies between adjacent measured logs is possible in many cases (Figs. 7, 11, 13, 16), individual bodies are not included on bedding diagrams (Figs. 6, 10, 12, 15). This is because the limits of bodies can become difficult to define where they thin and become highly disrupted, and correlations are inaccurate unless individual bodies were traced across the numerous gullies cutting exposures. Time limitations precluded such detailed tracing. Scale limitations allowed only generalized trends in bedset thickness and grainsize to be presented on sedimentary logs (Figs. 7, 11, 13, 16). A well exposed and unusually thick example of a lobate body was documented several kilometers to the west of the deposits presented in Fig. 6, at a similar stratigraphic level within the Chinji Formation (Figs. 22E, 24). This body is dominated by fine- to very fine-grained sandstone. It has a horizontal basal erosion surface and a convex-upward top that grades upward into overlying mudstone beds. The body is nearly 10 m thick (where thickest) and gradually thins to only a few meters thick in both directions over distances of about a kilometer. Bedsets are underlain by minor erosion surfaces, and fine upward. Nearly all bedsets are dominated by cross-laminated sandstone, although planar stratified sandstone or disrupted mudstone is common in their fining tops. Bur-
rows, root casts, and mudcracks disrupt all bedsets to some extent throughout the body. Individual bedsets thin, fine, contain less planar strata and become more disrupted in the direction the body thins. Where the body is thickest, bedsets dip at 1-2 ° in the direction the body thins (e.g. Fig. 22E). In places, bedsets vary slightly in dip, producing very low-angle truncations of adjacent bedsets. Bedset dips gradually decrease to horizontal toward the thinned edges of the body. Successive bedsets also tend to thin, fine, become dominated by cross-lamination, and become more disrupted vertically within the body. However, in locations where the body is thinner, successive bedsets coarsen and thicken upward in the lowest few decimeters (Fig. 24, log 2). Further along strike in the direction it thins, the body divides into multiple, thinner, fining-upward sets of bedsets that feather into adjacent mudstonedominated deposits (Fig. 24, log 1). Paleocurrents progressively vary by at least 30° along strike (Fig. 24). Wedge-shaped bodies display similar alongstrike and vertical variations in bedset geometry, grainsize, sedimentary structures, and disruption as lobate examples. However, wedge-shaped sandstone bodies thin away from an adjacent channel sandstone. Two types of wedge-shaped bodies are recognized: those with bedsets that pass conformably into the adjacent channel sandstone, and those that are truncated by the adjacent channel sandstone. Relatively thick and
-~70 meters~-~50 m e t e r s ~ 3 5 meters-4-~35 meters-~-
1
2
a 0
3
al_ .....
ss
0
,
4
5
a _, '
sl's's
0"~
.....
,as
ss
0
Fig. 24, Sedimentological logs measured at several locations along a thick lobate sandstone body. Figure 22E presents a photo of this same sandstone body taken where the body begins to thin in the opposite direction (about 100 m to the right of the area covered by these logs).
ANCIENT RIVER SYSTEMS IN THE HIMALAYAN FOREDEEP
coarse-grained examples of both types are documented in Figs. 15 and 16. In the upper example (Fig. 15, adjacent to Storey A), the wedge-shaped body thins and fines to dominantly mudstone within several hundred meters along strike (Fig. 16, logs 1-6, meters 37-42). Bedsets within this wedge-shaped body are horizontal and appear conformable with those in the top of the channel sandstone body to the left, inclined at about 3° just to the right of the major sandstone body margin, and gradually decrease in dip to the right as the wedge-shaped body thins (Fig. 22F). Where the wedge-shaped body is thickest, bedsets are dominated by planar strata, but contain a few decimeter-thick sets of large-scale cross-strata lower and cross-lamination higher. Bedsets gradually fine, become more disrupted and are increasing dominated by crosslamination to the east (right) as the wedge-shaped body thins. Paleocurrent orientations within the wedge-shaped body are generally parallel to the outcrop and nearly perpendicular to those within the adjacent major sandstone body to the left, but they gradually shift to become more outcrop-perpendicular along strike in the direction the body thins. Small 1-3 m thick and several meters wide channel-forms cut into the top of this wedgeshaped body where it is thickest to the left (e.g. Fig. 16, log 3, meters 40-42). Bedsets within these small channel forms are slightly coarser than those within the underlying wedge-shaped body, and contain large-scale cross-strata lower down but dominantly planar strata and crosslamination higher up. Paleocurrent orientations within these small channel forms are generally parallel to those within the adjacent channel sandstone body and thus are at a high angle to those within the underlying wedge-shaped body. In the second example (adjacent to Storey C in Fig. 15; Fig. 16, log 8, meters 21-28) bedsets within the wedge-shaped body are truncated by the upper part of the channel-filling margin of the adjacent major sandstone body. In this second example bedsets also dip to the right at a very low angle away from the channel margin, and paleocurrents are oriented generally perpendicular to those of the adjacent major sandstone body. Bedsets are finer grained and thinner (1 to
65
a few cm thick) than those adjacent to the channel in the first example, and are dominated by cross-lamination throughout. Most lobate and wedge-shaped bodies are thinner (1 to a few meters thick) and finer grained than the examples just described, and thinning along strike can only be noted when tracing the body over several hundred meters (Figs. 7, 11, 13, 16). Some examples do not appear to have erosional bases, and the base of these bodies are defined only by an abrupt vertical change in bedset thickness, grainsize and the level of disruption. Many bodies fine upward, but some coarsen upward or initially coarsen and then fine upward (e.g. Fig. 5, logs 21-23, meters 39-45, and many others). Bedsets within these thinner bodies appear horizontal in the field and sandstone-dominated bedsets occur only where the body is thickest. Coarser bedsets can have erosional bases with locally developed mud-clast breccia, but finer bedsets are not clearly erosional. Internally, coarser bedsets are similar to those described above, whereas finer-grained bedsets commonly contain horizontally laminated mudstones. Some finer-grained bedsets also contain symmetrical cross-lamination. Individual bedsets commonly can be traced along strike for many tens of meters, but in places they also thin laterally. Shallow channel forms, decimeters deep and meters wide, occur rarely within these bodies. These shallow channel forms are bounded below by a minor erosion surface, overlain by a thin layer of intraformational breccia, that truncates several underlying beds. Such channel forms are normally filled by horizontal mudstone bedsets. The thickness of individual bedsets within lobate and wedge-shaped bodies normally decreases vertically and in the direction the body fines, but the degree of disruption always increases upward. Therefore, in fining-upward bodies, thicker and coarser bedsets lower down contain relatively undisrupted cross-laminated siltstone to very fine sandstone, but thinner and finer bedsets higher up have highly disrupted mudstone internally and are commonly cut by abundant mudcracks (e.g. Fig. 7, log 5, meters 41-47). Conversely, in coarsening-upward bodies,
66
B. WILLIS
thin and fine bedsets lower down display well preserved horizontal lamination and only moderate disruption, but thicker and coarser bedsets higher up are completely disrupted internally (e.g. Fig. 7, log 5, meters 4-9).
A
~
~
~ " - . ~
Interpretation Meters-thick, lobate and wedge-shaped bodies are similar to modern levee or crevasse-splay deposits (Fisk, 1947; Kruit, 1955; Coleman, 1969; Singh, 1972; Ray, 1976; Farrell, 1987; Tye and Coleman, 1989; Smith et al., 1989). Lobate bodies record a cross-section through crevasse splays that were hundreds of meters to kilometers across, Wedge-shaped bodies, thinning away from an adjacent channel deposit, may record cross-sections through splay lobes or through levees extending along the adjacent channel. Where bedsets within the wedge appear conformable with those in the top of the adjacent channel body (e.g. Fig. 15, adjacent to Storey A) the s p l a y / l e v e e must have developed adjacent to the depositional channel mar'-in on the inside of a channel bend (Fig. 25a). Where the wedge body is truncated by the channel filling margin of the adjacent channel body (e.g. Fig. 15, adjacent to Storey C), the river channel must have continued to migrate laterally in the direction the splay/levee was prograding, eroding earlier formed channel proximal splay/ levee deposits (Fig. 25b). Along-strike thinning of bedsets, decrease of mean grainsize, increase in sediment disruption and change in sedimentary structures record a decrease in flow strength and deposition rate laterally in the direction the body thins. Crevasse channels were relatively rare as demonstrated by the scarcity of small channel-fills cutting into the top of lobate and wedge bodies. Thus much of the s p l a y / l e v e e deposition must have been associated with shallow sheet flows. Thicker, coarsergrained bodies record deposition by more rapid flows, presumably in crevasse channel-proximal locations. Horizontal basal erosion surfaces with scour structures record local erosion prior to deposition of the splay/levee. The high proportion of overbank deposits composed of these metersthick bodies (perhaps 70%) suggest that most of the sediments on the floodplain were deposited
B a ~ J b ~
a' '
"
~
"
--~ " " ""
.... .--"-- " - ' : - - ~
-13' - ••
Fig. 25. Bedding relationships between channel bodies and adjacent wedge shaped bodies. (a) A crevasse splay deposit formed on the inside of the channel bend will overly (and
perhapstruncate) the bedset concordant storey margin of the underlying channel sandstone body. (b) A crevasse splay/levee deposit formed on the outside of the channel bend will be truncated by the channel filling margin of the adjacent chan-
nel sandstone body, and channel proximal deposits may be removed as the channel continues to migrate.
by progradation of crevasse splays and levees, and not by gradual vertical accretion in floodbasins. Bedsets within these bodies record individual flood events. Minor erosion surfaces underlying bedsets record slight erosion of the splay/levee surface during initial or high flood stages. Inclined bedsets, which dip down to meet the basal erosion surface of the body at up to 3 °, indicate gradual progradation of the s p l a y / l e v e e onto the floodplain over many flood events. Bedset slopes suggest that prograding splays/levees could produce a local floodplain topography of nearly 10 m and local relief of nearly 1 : 20; however, such slopes were quite unusual. More usual examples (see Figs. 7, 11, 13, 16) suggest a floodplain topography of only a few meters and local relief less than 1: 100. Decreased bedset dips from channel-proximal to distal locations reflect decreased slopes away from the channel. The dominance of cross-lamination within most bedsets suggests most of these splays/levees were cov-
67
ANCIENT RIVER SYSTEMS IN T H E HIMALAYAN FOREDEEP
ered by ripples during episodes of deposition, Planar strata within fine-grained bedset tops may record rapid shallow flows over upper stage plane beds or ripple migration under conditions of a very low deposition rate. Upward decrease in grainsize within bedsets records waning overbank flows and deposition of mud from suspended load. Bioturbation, mudcracks and root casts record post-flood disruption of sediment, Paleocurrents within wedge-shaped bodies suggest flows parallel to directions of bedset dip (contrast with bedsets in channel deposits). Paleocurrents within wedge-shaped bodies, which are initially perpendicular to those within the adjacent channel body but become more parallel to those within the channel where the wedge thins along strike, record overbank flow away from the channel where splay/levee slopes are steep, but gradual diversion of flow to down-basin directions further onto the floodplain where splay/ levee slopes decrease (e.g. Fig. 16, adjacent to Storey A; cf. Smith et al., 1989). Paleocurrents within channel-belt-proximal crevasse channels suggest that they tend to flow obliquely away from the major channel belt, preferentially feeding new areas in the down-basin direction (e.g. Fig. 15, wedge deposit adjacent to Storey A). Thinner, mudstone-dominated bodies record deposition by slower flows, presumably in crevasse channel distal locations. Nearly horizontal bedsets suggest low floodplain relief in such locations. Rare, shallow channel-forms and local mud-clast breccia at the base of bedsets record local scouring of the floodplain surface during flood events. However, the absence of erosion surfaces underlying some bodies, and underlying many bedsets within bodies, indicate that scouring by flood waters was less extensive than within channel proximal locations. Rare symmetrical cross-lamination within bedsets records locally ponded water and the formation of wave ripples, Fining upward bodies record a progressive decrease in deposition rate and flow velocity over successive flood events. Coarsening-upward bodies probably record gradual progradation of crevasse splays or levees into crevasse channel distal areas over many flood events. Bedsets lower in bodies generally have well preserved internal
stratification, suggesting relatively rapid burial or deposition in areas of ponded water. The progressive upwards increase in bioturbation/disruption observed in most bodies may record gradual decrease in deposition rate with time. However, because bedsets become more disrupted upward even in coarsening-upward bodies, sediment lobes were probably deposited rapidly and then were followed by a period of non-deposition during which more deeply buried bedsets were protected from disruption.
Laminated mudstone bodies Description Laminated mudstone bodies (composed dominantly of claystone) are several decimeters to meters thick, have non-erosional bases, and have horizontal tops (e.g. see relatively thick examples in Fig. 7, logs 5-27, meters 30-40; Fig. 16, logs 2-8, meters 0-10). The base of these bodies conform to the topography of underlying deposits (most commonly the top surface of a paleosol). Some bodies grade upwards into an overlying horizontal paleosol, but more commonly these bodies are truncated upwards by overlying channel deposits or lobate bodies. Laminated mudstone bodies can continue along strike for kilometers. A few examples thin along strike and gradually become more disrupted until they are indistinguishable from the surrounding deposits, However, most examples continue until they are truncated along strike by a channel body. Laminated mudstone bodies can not be divided into bedsets, but show changes upwards. Lower down in bodies, sediments are relatively undisrupted (but contain sparse vertical burrows and root casts), most sediments are finely laminated (but centimeter-thick beds can be massive and calcified), and isolated lenses of symmetrical cross-laminated siltstone can occur. Root casts and mudcracks become abundant upwards in the body. Initially thin centimeter-deep mudcracks are restricted to thin horizons separated by beds of laminated mudstone, but near the top of the body, wider decimeters-deep cracks filled with clasts of laminated claystone completely disrupt the sediments.
68
B. WILLIS
Rarely, light gray calcareous laminated mudstones also occur, most commonly above and adjacent to channel filling deposits (e.g. Fig. 5, logs 3-7, meters 39-41). These mudstones occur in decimeters-thick bodies that can be traced along strike for only a few hundreds of meters before they become completely disrupted. Within such bodies there is a vertical alternation between centimeter-thick layers containing fine laminae and layers that are highly disrupted by burrows, roots, and mudcracks. These layers can be horizontal but also can be contorted and recumbently folded along irregular underlying disrupted surfaces, or where cut by large mudcracks. Fine, equant sparry to micritic carbonate concentrations occur along the walls of some mudcracks, burrows and root casts,
Interpretation Bodies of horizontally laminated claystone record deposition within shallow ponds on the floodplain that were at least kilometers across (e.g. Tye and Coleman, 1989). The irregular nonerosional bases and horizontal tops of bodies, and the horizontal laminae within, suggest lake sediments gradually filled topographic low areas. Root casts and lenses of symmetrical cross-laminated siltstone lower down within bodies suggest that ponds were shallow enough for bottom sediments to support plants and become modified by water surface waves, respectively. Sparse vertical burrows imply that ponds filled rapidly or were biologically restricted. Thin beds that are massive and highly calcified resemble hardgrounds, and suggest periods of low suspended sediment input and low wave reworking. Upward increase in disruption by bioturbation and desiccation cracks records shallowing of the ponds and periods of subaerial exposure of sediments as ponds filled, Calcareous laminated mudstone beds record ponded areas where deposition rates where low. High carbonate content, fine laminae conformably covering irregular surfaces and recumbent folding of some beds suggest they may have been covered by cyano-bacterial mats. Local concentration of calcite along mudcracks, burrows and root casts record periods of subaerial exposure and sediment desiccation,
Paleosols Nearly all mudstone beds show evidence of disruption by roots, and thus must have been disrupted to some degree by pedogenesis. However, many meters-thick layers in these strata are distinct from superjacent deposits due to their heightened red hues, intense disruption, and complex vertical variations of structure and fabric (Figs. 7, 11, 13, 16). Such disrupted layers cornprise nearly a third of mudstone-dominated deposits within this stratigraphic succession and are widely recognized to be distinct paleosols (Johnson et al., 1982, 1985; Retallack 1985, 1986, 1992; Behrensmeyer 1987, 1989; and others). Paleosols are not sediment bodies in the same sense as others described here because they are not defined by variations in primary sedimentary structures and textures (cf. sediment bodies described above). Rather, they are mappable layers of strata defined by distinct patterns of sediment disruption superimposed on underlying sediment bodies. Paleosols continue across entire documented intervals, except where they are truncated by overlying sediment bodies. Many paleosols stay nearly parallel along strike for kilometers (Figs. 6, 10, 12, 15). However, in some cases, paleosols gradually converge along strike, and some join laterally to form composite paleosols. In most cases, paleosols can be divided into two distinct horizons: an upper horizon that lacks disseminated groundmass carbonate, and a lower horizon that contains abundant carbonate concentrations. The thickness of these horizons and the abundance of various pedogenic structures within paleosols can vary greatly, both between different paleosols and along strike within individual paleosols (Figs. 7, 11, 13, 16). Detailed studies of along-strike variations in paleosol profiles and micromorphology in relation to variations in lithofacies and floodplain physiography are in progress (Behrensmeyer et al., 1992; Behrensmeyer and Willis, 1992; Willis and Behrensmeyer, 1992a, b; Behrensmeyer and Willis, unpubl, data), thus only a very generalized description of paleosols will be presented here. Higher within a paleosol, deposits are generally characterized by: (1) bright red-brown to
69
A N C I E N T R IVE R S YS T E MS IN T H E H I M A L A Y A N F O R E D E E P
strong-brown hues; (2) groundmass that lacks elfervescence in dilute hydrochloric acid (interpreted to reflect the leaching of carbonate from the upper part of the soil); (3) gray mottles (cornmonly reflecting local reduction along root casts and burrows); (3) coarse-grained concentrations (interpreted to reflect eluviation of fines from the upper part of the soil); (4) slickensided clay films (interpreted to record clay particles aligned along surfaces faulted due to clay swelling); (5) small carbonate or sesquioxide concentrations. Carbonate concentrations can occur along cracks but more commonly occur as micritic or fine crystalline calcite nodules, normally clustered along root casts, burrows and mudcracks. Sesquioxide concentrations occur as millimeter-diameter concretions and coatings on carbonate nodules. Ped structure is observed only rarely within these paleosols, and the structure of most paleosols is dominated by pervasive burrows, Lower in paleosols, abundant carbonate concentrations typically occur. Carbonate concentrations normally occur as large, composite nodules (several centimeters across), but also occur as smaller isolated nodules aligned along mudcracks, burrows and root casts. In some layers, composite nodules are elongate vertically along large mudcracks (reflecting preferential carbonate precipitation along preexisting mudcracks; cf. Blodgett, 1988; Joeckel, 1991). Sesquioxide nodules can be broadly disseminated within the paleosol, but also can occur densely packed directly above intervals dominated by carbonate nodules, Rarely argillic concentrations are observed along burrows and root casts, particularly in the top parts of layers containing abundant carbonate concentrations. Complex textural relationships within carbonate concentrations, and superposition of carbonate and sesquioxidic concentrations, suggest complicated histories of accretion and changing pedogenic conditions over time. Siwalik paleosols display vertical variations in structure and fabric similar to that observed within modern alluvial soils of the Indo-Gangetic basin (Sehgal et al., 1968; Sehgal and Stoops, 1972; Ahmad et al., 1977; Brammer and Brinkman, 1977; Brinkman, 1977; Sidhu and Gilkes, 1977; Sidhu et al., 1977). Estimates of time sepa-
rating paleosols based on paleomagnetic dating (order of 104 years; see Behrensmeyer, 1987; Willis, 1993b) are comparable to estimates for development of well differentiated soils formed within the modern Indus Basin (Ahmad et al., 1977). Sesquioxide concentrations along carbonate nodules and within small concretions appear most commonly in modern soils where the soil profile becomes seasonally waterlogged then desiccated causing repetitive cycles of reduction and reoxidation (Sehgal and Stoops, 1972; Brinkman, 1977; Sobecki and Wilding, 1983; Wright et al., 1991). Paleosols of the Chinji Formation contain sesquioxidic concentrations more commonly then paleosols of the Nagri Formation (80% versus 35% of paleosols, respectively), suggesting that seasonal waterlogging of floodplain surfaces was more common during deposition of the Chinji Formation. Siwalik Group paleosols have been interpreted as oxisols to alfisols underlying tropical, thickly wooded to savanna vegetation because of their red coloration and the presence of sesquioxide nodules (Johnson et al., 1981; Retallack, 1985, 1986, 1992). These interpretations also reflect fossil flora evidence extrapolated from the Siwaliks of India, and fossil faunal evidence from the Siwaliks of Pakistan. Paleosol-bounded sequences
Strata between well defined paleosols are composed of multiple lobate/wedge-shaped bodies, minor channel-form bodies, and bodies of |aminated claystone (Figs. 7, 11, 13, 16). Whereas lobate/wedge-shaped bodies and minor channelform bodies extend along strike for only a few tens of meters to a few kilometers (as described above), most paleosols extend for at least tens of kilometers along strike. Although some strata between well developed paleosols are intensively disrupted by roots, burrows and mudcracks, most contain primary stratification and are comparatively undisrupted relative to paleosols. This alternation between completely disrupted paleosols separating strata with preserved primary bedding defines larger-scale (1-10 m thick) sequences within these mudstone-dominated deposits. Understanding the geometry of paleosol-bounded
70 sequences relative to lithologic variations within, and relative to terminal margins of major channel belt deposits, is fundamental to interpretations of overbank deposition, Paleosols record ancient floodplain surfaces that received minimal deposition for prolonged periods of time (see above). Conversely, stratified sediments between paleosols record periods when overbank deposition rates were generally high and long hiatuses in deposition did not occur, Paleosol-bounded sequences thus reflect an alternation between periods when overbank deposition rate was generally high and continuous, and periods when overbank deposition rates were minimal. Because such sequences contain many bedsets, they can not be related to individual deposition events. Further, because sequences contain multiple minor channel-form and lobate/ wedge-shaped bodies, they can not be related to progradation of individual splays or levees onto the floodplain. Individual crevasse splays can be deposited within just several tens of flood events, and complex overbank sequences including overlapping crevasse lobes and channel deposits can form in on the order of a hundred years (e.g. Smith et al., 1989; Tye and Coleman, 1989). Similarly high deposition rates within these deposits are supported by relatively undisrupted sediments lower within paleosol-bounded sequences, Paleosol-bounded sequences within the Chinji and Nagri Formations represent on the order of 104 years on average (Willis, 1993b). Evidence for rapid deposition rates of sediments between paleosols suggests that most of the time recorded within these strata is associated with development of paleosols. Rapid deposition of overbank sequences have been reported in modern river systems associated with: (1) build-up of multiple, overlapping crevasse splays directly adjacent to the major channel belt (e.g. Fisk, 1947; Coleman, 1969; Farrell, 1987; see also Bown and Kraus, 1987; Smith, 1990); (2) progradation of sediment into low areas on the floodplain further away from the major channel belt associated with tributary, distributary or crevasse channels (e.g. Parkash et al., 1983; Wells and Dorr, 1987; Tye and Coleman, 1989); (3) rapid deposition of overbank deposits
B.WILLIS following river avulsion but directly preceding development of a new channel belt (e.g. Smith et al., 1989). Each case involves the deposition of lobate overbank deposits as flows pass to areas of lower gradient and decelerate. In the first case, overbank deposition would largely be associated with the build up of an alluvial ridge and deposition of coarse-grained sediments close to the major channel belt. In the latter two cases, overbank deposition would be associated with the filling of low areas on the floodplain, and thus the reduction of floodplain relief. In the latter cases, grainsize and thickness variations within sequences would not necessarily reflect the proximity of the major channel belt. In these deposits, most paleosol-bounded sequences vary in thickness in a complex way along strike (Figs. 6, 10, 12, 16). Some contain minor channel and coarser-grained crevasse-splay deposits where they are thicker, and they thin and fine along strike (e.g. Fig. 7, logs 6-23, meters y = 45 to 50; Fig. 11, logs 11-19, meters 3-9; amongst others). In most cases, however, the thickness of paleosol-bounded sequences is not related to internal lithologic variations. For exampie, the paleosol-bounded sequence to the left of Storey V (Fig. 6; Fig. 7, logs 1-12, meters 5-10) does not vary in thickness in a systematic way with distance from the minor channel body, and the sequence adjacent to Storey M (Fig. 10; Fig. 11, logs 13-24, meters 10-15) dramatically fines to the right but changes little in thickness. Laminated mudstone beds occur most commonly lower within finer-grained parts of sequences (e.g. Fig. 7, meters 30-40; Fig. 11, logs 18-24, meters 1015). In places, paleosols converge defining the margins of paleosol-bounded sequences (e.g. Fig. 6, between meters 0-10; Fig. 11, logs 15-19, meters 5-10). In most cases, however, paleosolbounded sequences continue across entire documented intervals despite local variations in thickness and internal lithology. Strata directly overlying the major sandstone body documented in Figs. 6 and 7 show variations along a paleosol-bounded sequence. This sequence is discussed in more detail below to emphasize relationships between sediment bodies within paleosol-bounded sequences. This se-
A N C I E N T R I V E R SYSTEMS IN T H E HIMALAYAN F O R E D E E P
quence is defined by two paleosol horizons that extend across the entire area studied: a lower one capping the major sandstone body (Fig. 6, about meters y = 30 to 35), and an upper one about 10 m higher (Fig. 6, meters y = 40 to 45). The lower paleosol records soil formation on an elevated channel belt following its abandonment by avulsion. Laminated mudstones across the base of this sequence suggest that this area later became a floodplain low that was initially ponded. Thickening of the sequence to the left can be related to filling of the depositional topography on the submerged channel belt deposit. This low ponded area was then locally invaded and filled by multipie levees or crevasse splays associated with minor channels. Multiple splay/levee deposits adjacent to Storey C (Fig. 6), and thinning of these deposits to the left, record decreased deposition rates away from the minor channel deposit (Fig. 7, logs 25-32, meters 35-45). However, because these splay/levee deposits are truncated to the right by Storey C, it is not possible to distinguish whether the channel eroded into a previously deposited splay lobe, or whether the splays/ levees formed away from the channel and lower beds were truncated as the channel enlarged or migrated. Storey B (to the left) truncates mainly laminated mudstones, thus clearly not all minor channels are associated with splay deposits (see also numerous other examples in Figs. 7, 11, 13, 16). Calcareous mudstones capping this channel sandstone and adjacent laminated mudstones to the right (Fig. 5, logs 3-8, about meter 40), suggest lower deposition rates occurred after this channel (B) was filled. Storey A also truncates mainly laminated mudstones; however, a metersthick levee deposit is associated with its top. Following deposition of Storey A, deposition rates of this sequence slowed and the deposits became capped by the upper paleosol. The deposits within this sequence suggest a rather complex set of depositional environments: including areas of ponded water that received only suspended muds, the development of overlapping crevasse splay lobes formed where flows decreased in velocity into ponded areas, and minor channels that locally had adjacent levees. Similarly complex histories of deposition are suggested by sediment vari-
71
ations across other paleosol-bounded sequences within the Chinji and Nagri Formations. Termination of storeys in major sandstone bodies against overbank deposits show how overbank deposits vary away from the edges of major channel belts. Most major channel margins truncate several paleosol-bounded sequences; however, only the topmost truncated sequence can be related to overbank deposition coeval with the existence of, and active deposition within, the adjacent channel belt. In some cases sequences adjacent to the top of a major channel margin are coarser-grained levee/splay deposits (e.g. Fig. 16, meters 40-45); however, such cases are rare and such splay/levee deposits normally thin and fine to mudstone along strike over less than a kilometer. More commonly, topmost sequences truncated by a major channel margin are n o t unusually thick nor coarse-grained, and they do not vary systematically away from the channel deposit (e.g. Fig. 11, logs 5-25, meters 40-55; logs 21-25, meters 15-20; and many others). In fact, sequences truncated by major channel margins can be quite fine grained (e.g. sequence dominated by claystone directly adjacent to the terminating margins of Storeys O and P in Fig. 11). The thickness of paleosol-bounded sequences adjacent to channel belt margins suggests that only a few meters of sediment was deposited on the floodplain adjacent to channel belts before they became abandoned, and thus channel belts were associated with relatively subdued alluvial ridges (less than a few meters in elevation). The absence of thick, coarse-grained overbank sequences adjacent to most major channel belt margins, and pronounced variations in the thickness and grainsize of both splay/levee deposits and paleosol-bounded sequences over distances of only a few hundred meters, do not support overbank deposition models that relate the thickness and grainsize of meter-scale overbank sequences to the proximity of coeval major river channels (cf. Johnson et al., 1981; Bridge, 1984b; Kraus, 1 9 8 7 a ;Smith, 1990). There is also little evidence that paleosols vary systematically away from channel-belt margins (cf. Bown and Kraus, 1987; Smith, 1990). Paleosol-bounded sequences clearly record a
72
complex distribution of environments including local levees, crevasse splays and ponded areas that do not vary in a simple way across the floodplain. Finely laminated mudstones lower within many paleosol-bounded sequences suggest that high overbank deposition rates were associated with the filling of low ponded areas on the floodplain. Abundance of paleosols with iron nodules (particularly in the Chinji Formation) suggest that local ponding of water in low areas on the floodplain was common for at least part of the year. It is not possible to tell whether there was continuous filling of different localized low areas on the floodplain (e.g. Tye and Coleman, 1989), versus wide spread (but short-lived) periods of rapid deposition in low areas associated with avulsion of major channel belts (e.g. Smith et al., 1989). Relatively thick sequences, that contain minor channel-form bodies of similar thickhess to individual storeys in major sandstone bodies, are the best candidates for avulsion-related deposits (e.g. Fig. 7, meters 30-40). Conclusions (1) The Miocene Chinji and Nagri Formations of northern Pakistan record fluvial deposition adjacent to the rising Himalayan Mountains. Rivers were predominantly braided with low channel sinuosities (1-1.3). Individual channel segments of the Chinji Formation were generally 100-200 m wide, 5-10 m deep and carried discharges of about 400-800 m3/s. Full channel-belt widths were less than 2 km, braiding index was about 2, and full river discharges were on the order of 103 m3/s. Deposits of a larger river are recorded within the Nagri Formation, where individual channel segments were 200-400 m wide, 15-30 m deep and carried discharges of about 3000-5000 m3/s. Full channel widths and braiding index of these larger rivers are difficult to estimate; however, there was clearly more than one coeval Channel and thus full channel discharges appear to have been on the order of 10,000 m3/s. (2) Channel bars were characterized by low angle (2-5 °) along-stream dipping surfaces and slightly steeper (5-11 °) cross-stream dipping sur-
B. WILLIS
faces. Bars had along-stream lengths of 0.5 to 1.5 km within Chinji Formation rivers, but 1.5-3 km within the larger Nagri Formation rivers. Most bars migrated downstream as they grew and increased in sinuosity, but not all. Channel switching was clearly an important process within channel belts as indicated by the downstream stacking of storeys and the dominance of coarse-grained channel fills. In a few locations, unambiguous evidence for braid bars is recorded by diverging paleocurrents around concave-upward bedding surfaces. (3) Floodplains had low relief, that was generally less than meters per kilometer, but higher local relief was associated with abandoned channel belts, small overbank channels, crevasse splays and levees. Most floodplain surfaces were covered by well developed soils. Overbank deposits are characterized by meters-thick stratified sequences separated by paleosol horizons. These sequences are dominated by overlapping crevasse splay and levee deposits. Finely laminated mudstones were deposited in ponded areas away from active splays/levees. The thicknesses and grainsize of these overbank sequences are not typically related to the proximity of coeval major river channels, in contrast to recently published overbank deposition models proposed for other basins (e.g. Bridge, 1984b; Kraus, 1987a; Smith, 1990). Instead most overbank deposition appears to be related to localized crevasse channels that carried sediment to low areas on the floodplain, or to rapid periods of floodplain deposition associated with channel belt avulsion. Acknowledgements This paper is part of the authors Ph.D. dissertation supervised by John S. Bridge at the State University of New York at Binghamton; it was greatly improved by his instructive comments throughout the research. I am grateful to A.K. Behrensmeyer of the Smithsonian Institution for introducing me to the Siwalik exposures. This research was supported by grants to J.S. Bridge and A.K. Behrensmeyer from N.S.F. and the Smithsonian Institution, and a grant to J.S. Bridge from the Donors of the Petroleum Research Fund
73
ANCIENT RIVER SYSTEMS IN THE HIMALAYAN FOREDEEP
(administered by the American Chemical Society).
J.S. Bridge, A.K. Behrensmeyer, S. Gabel, B. Demicco, T. Lowenstein, and J. Kradyna are thanked for their comments on the initial manuscript. A. Miall, P. Friend, D. Burbank, and an
anonymous
reviewer provided helpful
corn-
ments via reviews for Sedimentary Geology. This research benefitted from discussions in the field with J. Barry, A.K. Behrensmeyer, P. Friend, M. Raza, and M. Zaleha (amongst others). This research could not have been completed without the logistical support of the Geological Survey of Pakistan (Stratigraphy and Paleontology Branch, Islamabad), with specific thanks to officers M. Raza, K. Sheikh, I. Khan, and B. Jimal for their generous efforts that allowed me to stay in the field for extended periods. References Abroad, M., Ryan, J. and Paeth, R.C., 1977. Soil Development as a function of time in the Punjab River plains of Pakistan. Soil Sci. Soc. Am. J., 41: 1162-1165. Allen, J.R.L., 1983. Sedimentary Structures: their Character and Physical Basis. Elsevier, New York, N.Y., 1160 pp. Badgley, C. and Behrensmeyer, A.K., 1982. Paleoecology of Middle Siwalik sediments and faunas, northern Pakistan. Palaeogeogr., Palaeoclimatol., Palaeoecol., 30: 133-155. Barndt, J., Johnson, N.M., Johnson, G.D., Opdyke, N.D., Lindsay, E.H., Pilbeam, D.R. and Tahirkheli, R.A.K., 1978. The magnetic polarity stratigraphy and age of the Siwalik Group near Dhok Pathan village, Potwar Plateau, Pakistan. Earth Planet. Sci. Lett., 44: 355-364. Barry, J.C., Lindsay, E.H. and Jacobs, L.L., 1982. A biostratigraphic zonation of the Middle and Upper Siwaliks of the Potwar Plateau of northern Pakistan. Palaeogeogr., Palaeoclimatol., Palaeoecol., 37: 95-130. Barry, J.C., Johnson, N.M., Raza, S.M. and Jacobs, L.L., 1985. Neogene Mammalian faunal change in southern Asia: correlations with climatic, tectonic and eustatic events, Geology, 13: 637-640. Beaumont, C., 1981. Foreland basins. Geophys. J.R. Astron. Soc., 65: 291-329. Beck, R.A. and Burbank, D.W., 1987. Basin-scale isochronous paleocurrent study of Chron 9 interval, Middle Siwaliks, Potwar Plateau, northern Pakistan. Soc. Econ. Paleontol. Mineral., Annu. Midyear Mtg., Abstr., 4: 6. Beck, R.A. and Burbank, D.W., 1989. Regional fluvial facies variability within the C5N magnetic polarity interval, Siwalik Group, northwest Himalayan foredeep, Pakistan. 4th Int. Conf. Fluvial Sedimentology. International Association of Sedimentologists, Barcelona-Sitges, Progr. Abstr., p. 70.
Behrensmeyer, A.K., 1987. Miocene fluvial facies and vertebrate taphonomy in northern Pakistan. In: F.G. Ethridge, R.M. Flores and M.D. Harvey (Editors), Recent Developments in Fluvial Sedimentology. Soc. Econ. Paleontol. Mineral., Spec. Publ., 39: 169-176.
Behrensmeyer,A.K., 1989. Overbank facies and paleosols of the Chinji Formation, Northern Pakistan. 4th Int. Conf. Fluvial Sedimentology. International Association of Sedimentolgists, Barcelona-Sitges, Progr. Abstr., p. 77. Behrensmeyer, A.K. and Tauxe, L., 1982. Isochronous fluvial systems in Miocene deposits of Northern Pakistan. Sedi-
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Am., Abstr. Progr., 24: 229. Behrensmeyer, A.K., Quade, J. and Willis, B.J., 1992. Floodplains and paleosols in the siwalik Miocene sequence:
implicationsfor tophonomy and paleoecology of faunas. 5th North American Paleontological Convention, Abstr. Progr., p. 22. Berggren, W.A., Kent, D.V., Flynn, J J . and Van Couvering, J.A., 1985. Cenozoic geochronology. Geol. Soc. Am. Bull., 96: 1407-1418. Bernard, H.A., Major, C.F., Parrott, B.S. and Le Blunc, R.J., 1970. Recent sediments of southeast Texas: a field guide to the Brazos alluvial and deltaic plains and the Galveston Barrier Island complex. Bureau of Economic Geology, University of Texas, Austin, Guidebook, 11. Blodgett, R.H., 1988. Calcareous paleosols in the Triassic Dolores Formation, southwestern Colorado. In: J. Reinhardt and W.R. Sisleo (Editors), Paleosols and Weathering through Geologic Time: Principles and Applications. Geol. Soc. Am., Spec. Pap., 216: 103-121. Bluck, B.J., 1976. Sedimentation in some Scottish rivers of low sinuosity. R. Soc. Edinburgh Trans., 69: 425-456. Bluck, B.J., 1971. Sedimentation in the meandering River Endrick. Scott. J. Geol., 7: 93-138. Bown, T.M. and Kraus, M.J., 1987. Integration of channel and floodplain suites, I. Development sequence and lateral relations of alluvial paleosols. J. Sediment. Petrol., 57: 587-601. Brammer, H. and Brinkman, R., 1977. Surface-water gley soils in Bangladesh: environment, landforms and soil morphology. Geoderma, 17: 91-109. Brice, J.C., 1964. Channel patterns and terraces of the Loup Rivers in Nebraska. U.S., Geol. Surv., Prof. Pap., 422-D. Bridge, J.S., 1975. Computer simulation of sedimentation in meandering streams. Sedimentology, 22: 3-43. Bridge, J.S., 1977. Flow, bed topography, grain size and sedimentary structures in open channel bends: a three-dimensional model. Earth Surficial Process., 2: 401-416. Bridge, J.S., 1978. Paleohydraulic interpretation using mathematical models of contemporary flow and sedimentation in meandering channels. In: A.D. Miall (Editor), Fluvial Sedimentology. Can. Soc. Pet. Geol., Mem., 5: 723-742. Bridge, J.S., 1982. A revised mathematical model and Fortran
74 IV program to predict flow, bed topography and grainsize in open-channel bends. Comput. Geosci., 8: 91-95, Bridge, J.S., 1984a. Flow and sedimentary processes in river bends: comparison of field observations and theory. In: M. Elliott (Editor), River Meandering, Proceedings of Rivers '83, American Society Civil Engineers, pp. 857-872. Bridge, J.S., 1984b. Large-scale facies sequences in alluvial overbank environments. J. Sediment. Petrol., 54: 583-588. Bridge, J.S., 1985. Paleochannel patterns inferred from alluvial deposits: a critical evaluation. J. Sediment. Petrol., 55: 579-589. Bridge, J.S. and Diemer, J.A., 1983. Quantitative interpretation of an evolving ancient river system. Sedimentology, 30: 599-623. Bridge, J.S. and Gabel, S.L., 1992. Flow and sediment dynamics in a low-sinuosity, braided river: Calamus River, Nebraska Sandhills. Sedimentology, 39: 125-142. Bridge, J.S. and Gordon, E.A., 1985. Quantitative reconstructions of ancient river systems in the Oneonta Formation, Catskill Magnafacies. In: D. Woodrow and W. Sevon (Editors), The Catskill Delta. Geol. Soc. Am., Spec. Pap., 201: 163-181. Bridge, J.S. and Jarvis, J., 1982. Dynamics of a river bend: a study of flow and sedimentary processes. Sedimentology, 29: 499-542. Bridge, J.S., Smith, N.D., Trent, F., Gabel, S.L. and Bernstein, P., 1986. Sedimentology and morphology of a lowsinuosity river: Calamus River, Nebraska Sand Hills. Sedimentology, 33: 851-870. Brinkman, R., 1977. Surface-water gtey soils in Bangladesh: genesis. Geoderma, 17: 111-144. Bristow, C.S., 1987. Brahmaputra River: channel migration and deposition. In: F.G. Ethridge, R.M. Flores and M.D. Harvey (Editors), Recent Developments in Fluvial Sedimentology. Soc. Econ. Paleontol. Mineral., Spec. Publ., 39: 63-74. Burbank, D.W. and Beck, R.A., 1989. Early Pliocene uplift of the Salt Range; temporal constraints on thrust wedge development, northwest Himalaya, Pakistan. In: L.L. Malinconico and R.J. Lillie (Editors), Tectonics of the Western Himalayas. Geol. Soc. Am., Spec. Pap., 232: 113-128. Burbank, D.W. and Raynolds, R.G.H., 1988. Stratigraphic keys to the timing of deformation; an example from the northwestern Himalayan foredeep. In: C. Paola and K. Kleinspehn (Editors), New Perspectives in Basin Analysis. Springer-Verlag, New York, N.Y., pp. 331-351. Burbank, D.W., Raynolds, R.G.H. and Johnson, G.D., 1986. Late Cenozoic tectonics and sedimentation in the northwestern Himalayan foredeep, If. Eastern limb of the northwest syntaxis and regional synthesis. In: P.A. Allen and P. Homewood (Editors), Foreland Basins. Int. Assoc. Sedimentol., Spec. Publ., 8: 293-306. Cant, D.J., 1978. Development of a facies model for sandy braided river sedimentation: comparison of the South Saskatchewan River and the Battery Point Formation. In: A.D. Miall (Editor), Fluvial Sedimentology. Can. Soc. Pet. Geol., Mem., 5: 627-639. Cant, D.J. and Walker, R.J., 1976. Development of a braided
B. WILLIS fluvial facies model for the Devonian Battery Point Sandstone, Quebec. Can. J. Earth Sci., 13: 102-119. Cerveny, P.F., Naeser, N.D., Zeitler, P.K., Naeser, C.W. and Johnson, N.M., 1988. History of uplift, relief and provenance of the Himalaya during the past 18 million years: evidence from fission track ages of detrital zircons from Siwalik Group sediments. In: C. Paola and K. Kleinspehn (Editors), New Perspectives in Basin Analysis. SpringerVerlag, New York, N.Y., pp. 43-61. Coleman, J.M., 1969. Brahmaputra River: channel process and sedimentation. Sediment. Geol., 3: 129-239. Collinson, J.D., 1978. Vertical sequence and sand body shape in alluvial sequences. In: A.D. Miall (Editor), Fluvial Sedimentology. Can. Soc. Pet. Geol., Mem., 5: 577-586. Crowley, K.D., 1983. Large-scale configurations (macroforms), Platte River Basin, Colorado and Nebraska: primary structures and formative process. Geol. Soc. Am. Bull., 94: 117-133. Dietrich, W.E. and Smith, J.D., 1983. Influence of the point bar on flow through curved channels. Water Resour. Res., 19: 1173-1192. Dietrich, W.E. and Smith, J.D., 1984. Bed load transport in a river meander. Water Resour. Res., 20: 1355-1380. Farrell, K.M., 1987. Sedimentology and facies architecture of overbank deposits of the Mississippi River, False River Region, Louisiana. In: F.G. Ethridge, R.M. Flores, and M.D. Harvey (Editors), Recent Developments in Fluvial Sedimentology. Soc. Econ. Paleontol. Mineral., Spec. Publ., 39:111-120. Fatmi, A.N., 1973. Lithostratigraphic units of the KohartPotwar Province, Indus Basin, Pakistan. Mem. Geol. Surv. Pak., 10: 1-80. Ferguson, R.I. and Werritty, A., 1983. Bar development and channel changes in the gravelly River Feshie, Scotland. In: J.D. Collinson and J. Lewin (Editors), Modern and Ancient Fluvial Systems. Int. Assoc. Sedimentol., Spec. Publ., 6: 181-194. Fisk, H.N., 1944. Geological investigation of the alluvial valley of the lower Mississippi River. Mississippi River Commission, Vicksburg, Miss., 78 pp. Fisk, H.N., 1947. Fine grained alluvial deposits and their effects on Mississippi River activity. Mississippi River Commission, Vicksburg, Miss., 82 pp. Friend, P.F., 1983. Towards the field classification of alluvial architecture or sequence. In: J.D. Collinson and J. Lewin (Editors), Modern and Ancient Fluvial Systems. Int. Assoc. Sedimentol., Spec. Publ., 6: 345-354. Gardner, T.W., 1983. Paleohydrology and paleomorphology of a Carboniferous meandering fluvial sandstone. J. Sediment. Petrol., 53: 991-1005. Gibling, M.R. and Rust, B.R., 1990. Ribbon sandstones in the Pennsylvanian Waddens Cove Formation, Sydney Basin, Atlantic Canada: the influence of siliceous duricrusts on channel-body geometry. Sedimentology, 37: 45-66. Gill, W.D., 1951. The stratigraphy of the Siwalik Series in the northern Potwar, Punjab, Pakistan. Q.J. Geol. Soc. London, 107: 375-394. Harms, J.C., McKenzie, D.B. and McCubbin, D.G., 1963.
ANCIENT RIVER SYSTEMS IN THE HIMALAYAN FOREDEEP Stratification in modern sands of the Red River, Louisiana. J. Geol., 71: 566-580. Hooke, R.L., 1974. Shear stress and sediment distribution in a meander bend. UNGI Rap. 30, Univ. Uppsala, 58 pp. Jackson, R.G., 1975. Velocity-bedform-texture patterns of meander bends in the lower Wabash River of Illinois and Indiana. Geol. Soc. Am. Bull., 86: 1511-1522. Jackson, R.G., 1976. Depositional model of point bars in the Lower Wabash RNer. J. Sediment. Petrol., 46: 579-594. Jackson, R.G., 1978. Preliminary evaluation of lithofacies models for meandering alluvial streams. In: A.D. Miall (Editor), Fluvial Sedimentology. Can. Soc. Pet. Geol., Mem., 5: 543-576. Joeckel, R.M., 1991. Paleosol stratigraphy of the Eskridge Formation; Early Permian pedogenesis and climate in southeastern Nebraska. J. Sediment. Petrol., 61: 234-255. Johnson, G.D., Rey, P.H., Ardrey, R.H., Visser, C.F., Opdyke, N.D. and Tahirkheli, R.A.K., 1981. Paleoenvironments of the Siwalik Group, Pakistan and India. In: G. Rapp and C.F. Vondra (Editors), Hominid Sites: their Geologic Settings. Westview Press, Boulder, Colo., pp. 197-254. Johnson, N.M., Opdyke, N.D., Johnson, G.D., Lindsay, E.H. and Tahirkheli, R.A.K., 1982. Magnetic polarity stratigraphy and ages of Siwalik group rocks of the Potwar Plateau, Pakistan. Palaeogeogr., Palaeoclimatol., Palaeoecol., 37: 17-42. Johnson, N.M., Stix, J., Tauxe, L., Cerveny, P.F. and Tahirkheli, R.A.K., 1985. Palaeomagnetic chronology, fluvial processes and tectonic implications of the Siwalik deposits near Chinji Village, Pakistan. J. Geol., 93: 27-40. Johnson, N.M., Sheikh, K.A., Saunders, E.D. and McRae, L.E., 1988. Use of magnetic reversal time lines in stratigraphic analysis: a case study in measuring unsteady and nonuniform sedimentation. In: C. Paola and K. Kleinsphen (Editors), New Perspectives in Basin Analysis. Springer-Verlag, New York, N.Y., pp. 191-200. Keller, H.M., Tahirkheli, R.A.K., Mirza, M.A., Johnson, G.D., Johnson, N.D. and Opdyke, N.D., 1977. Magnetic polarity stratigraphy of the upper Siwaliks deposits, Pabbi Hills, Pakistan. Earth Planet. Sci. Lett., 36: 187-201. Kocurek, G., Knight, J. and Havholm, K., 1992. Outcrop and semi-regional three-dimensional architecture and reconstruction of a portion of the Eolian Page Sandstone (Jurassic). In: A. Miall and N. Tyler (Editors), The ThreeDimensional Facies Architecture of Terrigenous Clastic Sediments and its Implication for Hydrocarbon Discovery and Recovery. Soc. Econ. Paleontol. Mineral. Concepts in Sedimentology and Paleontology, 3: 25-43. Kraus, M.J., 1987a. Integration of channel and floodplain suites, II. Vertical relations of alluvial paleosols. J. Sediment. Petrol., 57: 602-612. Kraus, M.J., 1987b. Nodular evidence of ancient forests. Geol. Soc. Am., Abstr. Progr., 19: 733. Kruit, C., 1955. Sediments of Rhone delta: grainsize and microfauna. Verh. K. Ned. Geol. Mijnbouwk. Genoot., Geol. Ser., 15: 357.
75 Langbein, W.B. and Leopold, L.B., 1966. River meanders-theory of minimum variance. U.S. Geol. Surv., Prof. Pap., 422H: 1-15. Lewis, G.E., 1937. A new Siwatik correlation. Am. J. Sci., 33: 191-204. Miall, A.D., 1985. Architectural-element analysis: a new method of facies analysis applied to fluvial deposits. Earth Sci. Rev., 22: 261-308. Miall, A.D., 1988a. Architectural elements and bounding surfaces in fluvial deposits: Anatomy of the Kayenta Formation (lower Jurassic), southwest Colorado. Sediment. Geol., 55: 233-262. Miall, A.D., 1988b. Reservoir heterogeneities in fluvial sandstones: Lessons from outcrop studies. Am. Assoc. Pet. Geol. Bull., 72: 682-697. Miall, A. and Tyler, N. (Editors), 1992. The three-dimensional facies architecture of terrigenous clastic sediments and its implication for hydrocarbon discovery and recovery. Soc. Econ. Paleontol. Mineral. Concepts in Sedimentology and Paleontology, 3, 309 pp. Nickel, E., 1983. Environmental significance of freshwater oncoids, Eocene Guarga Formation, southern Pyrenees, Spain. In: P.M. Peryt (Editor), Coated Grains. SpringerVerlag, Berlin, pp. 308-329. Opdyke, N.D., Lindsay, E.H., Johnson. G.D., Johnson, N.M., Tahirkheli, R.A.K. and Mizra, M.A., 1979. Magnetic polarity stratigraphy and vertebrate paleontology of the Upper Siwalik Subgroup of northern Pakistan. Palaeogeogr., Palaeoclimatol., Palaeoecol., 27: 1-34. Opdyke, N.D., Johnson, N.M., Johnson, G.D., Lindsay, E.H. and Tahirkheli, R.A.K., 1982. Paleomagnetism of the Middie Siwalik Formations of Northern Pakistan and rotation of the Salt Range Decollement. Palaeogeogr., Palaeoclimatol., Palaeoecol., 37: 1-15. Parkash, B., Awasthi, A.K. and Gohain, K., 1983. Lithofacies of the Markanda terminal fan, Kurukshetra district, Haryana, India. In: J.D. Collinson and J. Lewin (Editors), Modern and Ancient Fluvial Systems. Int. Assoc. Sedimentol., Spec. Publ., 6: 405-420. Pilbeam, D.R., Behrensmeyer, A.K., Barry, J.C. and Shah, S.M.I., 1979. Miocene sediments and faunas of Pakistau. Postilla, 179: 1-45. Ramps, A. and Sopena, A., 1983. Gravel bars in low-sinuosity streams (Permian and Triassic, central Spain). In: J.D. Collinson and J. Lewin (Editors), Modern and Ancient Fluvial Systems. Int. Assoc. Sedimentol., Spec. Publ., 6: 301-312. Ramps, A., Sopena, A. and Perez-Arlucea, M., 1986. Evolution of Buntsandstein fluvial sedimentation in the northwestern Iberian Ranges (central Spain). J. Sediment. Petrol., 56: 862-875. Ray, P.K., 1976. Structure and sedimentological history of the overbank deposits of a Mississippi River point bar. J. Sediment. Petrol., 46: 788-801. Raynolds, R.G.H. and Johnson, G.D., 1985. Rates of Neogene depositional processes, northwestern Himalayan
76 foredeep margin, Pakistan. In: N.J. Snelling (Editor), The Chronology of the Geologic Record. Geol. Soc. London, Mem., 10: 297-311. Retallack, G.J., 1985. Fossil soils as grounds for interpreting Ihe advent of large plants and animals on land. Philos. Trans. R. Soc. London, B.309: 105-142. Retallack, G.J., 1986. The fossil record of soils. In: V.P. Wright (Editor), Paleosols: their Recognition and Inter• pfetation. Princeton University Press, Princeton, N.J., pp. 1-57. Retallack, G.J., 1992. Miocene Paleosols and Ape Habitats of Pakistan and Kenya. Oxford Monographs on Geology and Geophysics, 19, Oxford Press, 346 pp. Sarkar, S.K. and Basumallick, S., 1986. Morphology, structure and evolution of a channel island in the Barakar River, Barakar, West Bengal. J. Sediment. Petrol., 38: 747-754. Schafer, A. and Staff, K.R.G., 1978. Permian Saar-Nahe basin and recent Lake Constance (Germany): two environments of lacustrine algal carbonates. In: A. Matter and M.E. Tucker (Editors), Modern and Ancient Lake Sediments. Int. Assoc. Sedimentol., Spec. Publ., 2: 83-107. Schwartz, D.E., 1978. Sedimentary facies, structures, and grain size distribution: the Red River in Oklahoma and Texas. Trans. Gulf Coast Assoc. Geol. Soc., 28: 473-492. Sehgal, J.L. and Stoops, G., 1972. Pedogenic calcite accumulation in arid and semi-arid regions of the Indo-Gangetic Alluvial Plain of Erstwhile Punjab (India)--their morphology and origin. Geoderma, 8: 59-72. Sehgal, J.L., Sys, C. and Bhunbla, D.R., 1968. A climatic soil sequence from the thar desert to the Himalayan Mountains in Punjab (India). Pedologie, 18: 351-371. Shelton, J. and Noble, R., 1974. Depositional features of braided-meandering streams, Cimarron River, northcentral Oklahoma. Am. Assoc. Pet. Geol. Bull., 58: 742749. Sidhu, P.S. and Gilkes, R.J., 1977. Mineralogy of soils developed on alluvium in the Indo-Gangetic Plain (India). Soil Sci. Soc. Am. J., 41: 1194-1201. Sidhu, P.S., Sehgal, J.L., Sinha, M.K. and Randhawa, N.S., 1977. Composition and mineralogy of iron-manganese concretions from some soils of the Indo-Gangetic Plain in northwest India. Geoderma, 18: 241-249. Singh, A. and Bhardwaj, B.D., 1991. Fluvial facies model of the Ganga River sediments, India. Sediment. Geol., 72: 135-146. Singh, I.B., 1972. On the bedding in the natural-levee and the point-bar deposits of the Gomti River, Uttar, India. Sediment. Geol., 7: 309-317. Smith, N.D., 1971. Transverse bars and braiding in the lower Platte Ri.er, Nebraska. Geol. Soc. Am. Bull., 82: 34073420. Smith, N.D., Cross, T.A., Dufficy, J.P. and Clough, S.R., 1989. Anatomy of an avulsion. Sedimentology, 36: 1-23. Smith, R.M.H., 1990. Alluvial paleosols and pedofacies sequences in the Permian lower Beaufort of the southwestern Karoo Basin, South Africa. J. Sediment. Petrol., 60: 258-276.
B. WILLIS Sobecki, T.M. and Wilding, L.P., 1983. Formation of calcic and argillic horizons in selected soils of the Texas Coast Prairie. Soil Sci. Soc. Am. J., 47: 707-715. Stanley, K.O. and Fagerstrom, J.A., 1974. Miocene invertebrate trace fossils from a breaded river environment, western Nebraska, U.S.A. Palaeogeogr., Palaeoclimatol., Palaeoecol., 15: 63-82. Tauxe, L. and Opdyke, N.D., 1982. A time framework based on magnetostratigraphy for the Siwalik sediments of the Khaur area, northern Pakistan. Palaeogeogr., Palaeoclimatol., Palaeoecol., 37: 43-61. Thomas, R.G., Smith, D.G., Wood, J.M., Visser, J., Calverley-Range, E.A. and Koster, E., 1987. Inclined heterolithic stratification--terminology, description and significance. Sediment. Geol., 53: 123-179. Todd, S.P. and Went, D.J., 1991. Lateral migration of sand-bed rivers: examples from the Devonian Glashabeg Formation, SW Ireland and the Cambrian Alderney Sandstone Formation, Channel Islands. Sedimentology, 38: 997-1011. Tye, R.S. and Coleman, J.M., 1989. Depositional processes and stratigraphy of fluvially dominated lacustrine deltas: Mississippi Delta Plain. J. Sediment. Petrol., 59: 973-996. Visser, C.F. and Johnson, G.D., 1978. Tectonic control of Late Pliocene Molasse sedimentation •in a portion of the Jhelum Re-entrant. Geol. Rundsch., 67: 15-37. Wells, N.A. and Dorr, J.A., 1987. A reconnaissance of sedimentation on the Kosi alluvial fan of India. In: F.G. Ethridge, R.M. Flores and M.D. Harvey (Editors), Recent Developments in Fluvial Sedimentology. Soc. Econ. Paleontol. Mineral., Spec. Publ., 39: 51-61. Willis, B.J., 1989. Palaeochannel reconstructions from point bar deposits: a three-dimensional perspective. Sedimentology, 36: 757-766. Willis, B.J., 1992. Evolution of Miocene Fluvial Environments in Chinji Area, Potwar Plateau, Northern Pakistan. Ph.D. diss., State University of New York at Binghamton, 297 pp. (unpublished). Willis, B.J., 1993a. Interpretation of bedding geometry within ancient point-bar deposits. In: M. Marzo and C. Puigdefabregas (Editors), Alluvial Sedimentation. Int. Assoc. Sedimentol. Spec. Pap., 17: 101-114. Willis, B.J., 1993b. Evolution of Miocene fluvial systems in the Himalayan foredeep through a two kilometer-thick succession in northern Pakistan. Sediment. Geol., 88: 77-121. Willis, B.J. and Behrensmeyer, A.K., 1992a. Morphology of Siwalik paleosols exposed in northern Pakistan. Geol. Soc. Am., Abstracts with program, 24:349 Willis, B.J. and Behrensmeyer, A.K., 1992b. Relationship of overbank deposition and pedogenesis within the Chinji Formation. Geol. Soc. Am., Abstr. Progr., 24: 229. Wright, V.P., Vanstone, S.D. and Robinson, D., 1991. Ferrolysis in Arundian alluvial paleosols: evidence of a shift in the early Carboniferous monsoonal system. J. Geol. Soc., London, 148: 9-12,
1
Elsevier Science Publishers B.V., Amsterdam
Ancient river systems in the Himalayan foredeep, Chinji Village area, northern Pakistan Brian Willis
*
Department of Geological Sciences, State University of New York at Binghamton, N.Y. 13902-6000, USA Received December 17, 1991; revised version accepted February 3, 1993
ABSTRACT
The Miocene Chinji and Nagri Formations (Siwalik Group) of northern Pakistan record ancient fluvial environments in the Himalayan foredeep basin. Excellent exposures on the Potwar Plateau (Chinji Village area) allowed detailed documentation of the geometry and stacking of sediment bodies that comprise these strata, and of variation of large-scale bedding geometry, grainsize, sedimentary structures and paleocurrent orientations within such bodies. Major sandstone bodies are tens of meters thick and are continuous along strike for many kilometers. They are composed internally of interconnected channel belt deposits, each of which contain several storeys (channel bar and fill deposits). Individual storeys, defined by inclined bedsets dipping down to a major basal erosion surface at up to 11°, are only 5 to 15 m thick within the Chinji Formation but can be up to 30 m thick within the Nagri Formation. Bedsets within storeys reflect sediment accreted during individual flood events. Along-strike variation of bedsets within storeys and the stacking patterns of storeys within channel belt deposits reflect changes in the channel bed through time due to the growth of bars, migration of channels, and channel cutoff. Generally, braided channel patterns are indicated by the relatively large number of storeys within individual channel belt deposits exposed perpendicular to paleoflow, abundant evidence for channel bar superposition due to channel switching, local evidence for mid-channel bars, dominance of coarse-grained channel fills, and low paleocurrent variations. Paleochannel reconstructions from storeys exposed within the Chinji Formation indicate that individual channel segments generally had widths of 80-200 m, maximum depths of 4-13 m, wavelengths of 1.6-2 km and discharges of 400-800 m3/s. Full channel belt widths (1-2 kin) estimated from exposures perpendicular to palcoflow, and evidence for 2-3 coeval channels within channel belts, indicate full channel discharges of 1500-2000 m3/s. Larger channel segments reconstructed from the Nagri Formation had widths of 200-400 m, maximum depths of 15-30 m, wavelengths of 3-5 kin, and discharges of 3000-5000 m3/s. As Nagri Formation channel systems were also clearly braided, full channel dimensions and discharge estimates are probably at least a factor of two greater than for individual channel segments (i.e. order of 10,000 m3/s). Strata between major sandstone bodies are dominated by lobate and wedge-shaped bodies (crevasse splay and levee deposits), minor channel-form bodies (deposits of minor floodplain channels), laminated mudstone bodies (lake deposits) and paleosols. These strata are arranged into meters to 10 m thick stratified sequences that were rapidly deposited, bounded by well developed paleosols recording periods when deposition rates were low. The thickness and grainsize of such paleosol bounded sequences are not directly related to the proximity of a major channel deposit along strike. Instead, such sequences within overbank deposits appear to reflect local progradation of splays and levees into low areas on the floodplain.
* Present address: Bureau of Economic Geology, University of Texas at Austin, University Station, Box X, Austin, TX 78713-7508, USA. 0037-0738/93/$06.00 © 1993 - Elsevier Science Publishers B.V. All rights reserved
2
13.WILLIS
\ Fig. 1. Location map showing area studied.
Introduction
The Miocene Siwalik Group of the IndoPakistan subcontinent provides a long record of fluvial deposition adjacent to the rising Himalayan Mountains (Figs. 1, 2). Understanding the evolution of the fluvial systems that flowed across the Himalayan foredeep basin requires detailed interpretations of these deposits. Spec-
~ ~
Z
SIWALIKUPPER ~RotrP
~DHOK ~ PATI~AN ,,, F O ~ T I O N z~ m
~ ~
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FORMATION ~ ~
z 17
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Fig. 2. Lithostratigraphy of the Siwalik Group exposed on the Potwar Plateau, Punjab, northern Pakistan. Age of formation boundaries based on paleomagnetic data of Johnson et al. (1985) and the paleomagnetic time scale of Berggren et al.
(1985).
tacular exposures in the Chinji Village area on the Potwar Plateau of northern Pakistan extend for tens of kilometers along strike and for kilometers upsection. These exposures allowed detailed documentation of: (1)the dimensions of sediment bodies within these strata; (2) the nature of stacking, lateral truncation and lateral thickness variations of bodies along strike; and (3) spatial variations in bedding geometry, grainsize, sedimentary structures and paleocurrent orientations within bodies. From these data, ancient channels were reconstructed in terms of their dimensions, channel pattern, bar geometry, hydraulics and migration behavior. Ancient floodplains were reconstructed in terms of their physiography, local depositional environments and varying depositional rates. The ultimate objective of this study is to understand how variations in these deposits are related to changing paleoenvironments and the evolution of the foredeep basin. Stratigraphic studies completed on the Potwar Plateau over the past few decades provide a framework for more detailed sedimentological study. Detailed lithostratigraphy (Lewis, 1937; Gill, 1951; Fatmi, 1973; Pilbeam et al., 1979; Johnson et al., 1982), biostratigraphy (Barry et al., 1982, 1985) and chronostratigraphy based on magnetic polarity (Keller et al., 1977; Barndt et al., 1978; Visser and Johnson, 1978; Opdyke et al., 1979, 1982; Tauxe and Opdyke, 1982; Johnson et al., 1982, 1985, 1988; Burbank et al., 1986; and o t h e r s ) have been established in several areas of the Potwar Plateau, allowing the timing of lithologic and faunal changes, and depositional rates,
ANCIENT RIVER SYSTEMS IN THE HIMALAYAN FOREDEEP
to be compared across the basin. Stratigraphic studies demonstrate that deposits filling the Himalayan foredeep basin varied greatly in both space and time. In particular, formations tend to thin and fine to the south and east, their boundaries can be diachronous by millions of years, and deposition rates varied by at least a factor of two between different formations (Gill, 1951; Barry et al., 1985; Johnson et al., 1985; Beck and Burbank, 1987, 1989).
METERS 21002000180~~L'----0 I~-"-"
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Fig. 3. M e a n grainsize of section measured along the Gabhir Kas, Chinji Village area, northern Pakistan (grainsize is averaged for each 1-m interval). Letters A, B, C and C ' m a r k stratigraphic
levelsdocumented,
3
Although stratigraphic studies have allowed fairly detailed interpretations of the relationship between tectonism and basin sedimentation for some stratigraphic intervals within the basin (e.g. Visser and Johnson, 1978; Raynolds and Johnson, 1985; Johnson et al., 1985; Burbank and Raynolds, 1988; Cerveny et al., 1988; Burbank and Beck, 1989; Beck and Burbank, 1989), more detailed interpretation of depositional environments is needed in order to understand relationships between the varying paleoenvironments, evolution of organisms, basin physiography, and extrabasinal controls such as tectonics, eustasy and climate acting on various spatial scales (e.g. Behrensmeyer and Tauxe, 1982; Badgley and Behrensmeyer, 1982; Behrensmeyer, 1987). The ChinjiNagri Formation boundary (Fig. 3) marks a particularly important time of change in the Siwalik Group where deposition rates increase, the succession dramatically coarsens, a major faunal turnover occurs, and there is a dramatic decrease in the abundance of preserved faunal remains (Johnson et al., 1985; Barry et al., 1982, 1985). This paper focuses on documenting variation among deposits along specific stratigraphic intervals within the Chinji and Nagri Formations in order to understand the distribution of depositional environments, and the physical processes controlling deposition within paleochannels and on ancient floodplains. A companion paper (Willis, 1993b) examines larger-scale sediment variations across these formations in order to document the evolution of these fluvial systems during this important change in foredeep basin deposition. The stratigraphic section studied is located along the Gabhir River in the Punjab of northern Pakistan (Fig. 1). This area includes the stratotypes for both the Chinji and Nagri Formations (Fatmi, 1973). Strata here dip to the north at 3° to 18° but are otherwise generally undeformed. The deposits were documented along three different intervals within this 2-km-thick succession (Fig. 3). A map traced from air photos shows positions of measured sections across each stratigraphic interval documented (Fig. 4). Nearly complete east-west oriented exposure for several kilometers along strike at each interval allowed
4
B.WILLIS
construction of bedding diagrams from photo mosaics, vertical logs, and along-strike tracing of beds in the field. Bedding diagrams display the geometry and distribution of sandstone bodies, the geometry of bedsets within sandstone bodies, paleocurrent orientations relative to the outcrop plane, and traces of distinct paleosols. Sedimentary logs provide more detailed information on variations in sedimentary textures, structures, and paleocurrent orientations (relative to north) for specific sections within bedding diagrams. Logs were generally spaced closely enough to document along-strike variations within individual sandstone bodies and mudstone sequences; however, in places vertical cliffs limited access (particularly within the Nagri Formation). A key to symbols on bedding diagrams and sedimentary logs is presented in Fig. 5. No single bedding surface could be assumed perfectly horizontal along strike across an entire stratigraphic interval studied. A datum was established for successive logs measured along each interval by the correlation of several different paleosols, each of which appeared generally horizontal when traced along strike. Successive logs were aligned such that the average vertical dis-
tance of correlated paleosols from the datum was minimized. Thus it is assumed that only local floodplain topography caused paleosol horizons to vary significantly from paleo-horizontal and that down-basin slopes were negligible across distances studied. This method provides a better estimate of how logs align along strike than that based on arbitrarily assigning one traceable surface as horizontal because local relief reflected by any one bedding surface or paleosol is clearly much greater than regional floodplain slopes. Together the three stratigraphic intervals documented provide a sample of upsection sediment variations across the Chinji and Nagri Formations. The lower Chinji Formation has a low proportion of major sandstone bodies relative to mudstone-dominated beds (30%) and relatively low sediment accumulation rates (200 m / M a compacted thickness). The lower Chinji Formation was examined in detail along a 4 km by 60 m thick interval (Fig. 3, Level A, meters 347-407; bedding diagram, Fig. 6; associated sedimentary logs, Fig. 7; selected photos, Fig. 8). Strata near the Chinji-Nagri Formation boundary record a period of dramatically increased sediment accumulation rates (300-500 m / M a compacted thick-
:.2
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Fig. 4. Map of studyarea showingthe Gabhir Kas and tributaries (solid lines),positionsof stratigraphicintervalsstudied(Levels A, B, C, C'), and positionsof measuredsectionsused to constructbeddingdiagrams(numbereddots).
ANCIENT RIVER SYSTEMS IN THE HIMALAYAN FOREDEEP
ness) that mark the start of a pronounced upsection coarsening within this stratigraphic succession. These strata were also examined over a 4 km by 60 m thick interval (Fig. 3, Level B, meters 865-925; see bedding diagram, Fig. 10; associated sedimentary logs, Fig. 11; selected photos, Fig. 9). Finally the Nagri Formation is characterized by greater proportions of major sandstone bodies (75% of deposits), and continued high overall sediment accumulation rates (300-400 m / M a compacted thickness). Strata within the upper Nagri Formation were documented along a 3 km by 110 m thick interval (Fig. 3, Level C, meters
5
1512-1612; bedding diagram, Fig. 12; associated sedimentary logs, Fig. 13; and selected photos, Fig. 14), and along a second 750 m by 65 m thick interval located immediately below and to the west (left) of the first (Fig. 3, Level C', meters 1462-1517; bedding diagram, Fig. 15; associated sedimentary logs, Fig. 16). The strata can be divided into major gray sandstone bodies, and red to brown mudstonedominated intervals containing thin gray-brown sandstone beds. To facilitate descriptions and interpretations, these two types of strata are addressed separately below.
KEY TO SYMBOLS: BEDDING DIAGRAMS ~
Sandstone
~
~
Mudstone
Major erosion s u r f a c e
~
Paleosol
Minor erosion surface Paleocurrent arrow (up is into the outcrop)
~
A
25 Position of measured section
KEY TO SYMBOLS: SEDIMENTARY LOGS Carbonate-deficient mudstone
-~
Storey boundary
Laminated
~
Centimeter-scale bedding Decimeter-scale bedding
~/~
Paleocurrent arrow (north is up) Disruption (moderate, complete)
Symmeffi_calcross-lamination ~-"'~
Asymmetrical cross-lamination large-scale planar cross-strata Large-scale trough cross-sUata Planar strata
IQ toO'~ Q O o o o
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Intraformationalpebbles Extraformational pebbles Erosion surface
(O),(~jO0
Carbonate nodules (sparse,common, abundant)
(e) ,m,oo Sesquioxide nodules (sparse, common, abundant) "K'~ Mudcracks Root casts m Mottles P" Flutes 7" Tool marks G Gutter casts " Clay lined surfaces ~ Slickensided surfaces
1 ISandstone[ I Siltstone Claystone Fig. 5. Symbolkey for bedding diagrams and logs.
6
Major sandstone bodies
Description of major sandstone bodies Most major sandstone bodies continue along strike for the length of exposures (up to tens of kilometers) and vary from 10 to nearly 100 m in thickness. One major sandstone body was documented in the lower Chinji Fm. (Level A; Figs. 6, 7, meters y = 15-35) and at the Chinji-Nagri Fm. boundary (Level B; Figs. 10, 11; generally between meters y = 20-50 but locally thicker), whereas three major sandstone bodies were documented within the upper Nagri Formation (Level C; Figs. 12, 13; roughly between meters y = 0-40, 45-70, and 75-90, respectively). These sandstone bodies are underlain by a major erosion surface and are capped, in most locations, by a paleosol, Sandstone bodies are composed internally of many storeys stacked both along strike and vertically. Pronounced changes in the thickness of major sandstones along strike are generally associated with the termination of storeys aligned across different horizons within the body. Individual storeys vary between 4 and 30 m thick and extend along strike for a few 100 m to over 3 km. They are defined at their base by a major erosion surface and internally are composed of a set of inclined bedsets. Basal erosion surfaces of storeys display meter-scale relief and locally exhibit flute casts, gutter casts and tool marks. Bedset geometries, sedimentary structures, mean grainsize, paleocurrent orientations and the extent of sediment disruption vary in a systematic way both vertically and along strike within storeys (see Fig. 17). Bedding diagrams (Figs. 6, 10, 12) do not show traces of all bedset bounding surfaces observed (e.g. contrast Fig. 17, example 1 with Fig. 12, Storey J; and Fig. 17, example 2 with Fig. 6, Storey O), rather these diagrams present only general trends in bedset dip direction and steepness across each storey. In most storeys, bedsets dip dominantly in a single direction such that they concordantly overly one storey margin and are discordant with the opposite margin (e.g. Fig. 17). Bedsets within such storeys can be divided into .two groups. Those adjacent to the concordant margin are more
B. WILLIS
steeply inclined a n d / o r progressively steepen in dip, whereas those adjacent to the opposite margin progressively decrease in dip and become more concave upwards. Less commonly storeys contain convex upwards bedsets, and in these cases bedsets dip towards both margins of the storey (e.g. Fig. 6, Storey R; Fig. 10, Storey H). Bedsets (sensu Bridge and Diemer, 1983) within storeys are equivalent to lateral accretions of Allen (1983), inclined strata of Thomas et al. (1987), and stratum bounded by third-order surfaces of Miall (1988a). Beds within sets are strata containing similar lithofacies (i.e. similar sedimentary structures and textures) or are associated with individual sedimentary structures (e.g. an individual set of cross-stratification), and are equivalent to strata bounded by first- and second-order surfaces of Miall (1988a). Trends in bedset dip across storeys can be difficult to discern where bounding surface dips are low, and where adjacent bounding surfaces are separated by longer distances along strike (i.e. rarely, within the Nagri Formation, up to several tens of meters). In some locations, dips of bedset bounding surfaces could only be recognized by tracing individual bedding surfaces along strike over hundreds of meters. Furthermore, when outcrops are viewed from a distance, bedset bounding surfaces can occasionally be confused with the bounding surfaces of individual cross-sets, which rarely also extend along strike for several tens of meters (e.g. Fig. 14B). Despite local difficulties, variation of bedset geometry and internal characteristics can be confidentially mapped across most storeys (see Figs. 6, 10, 12). Bedsets are normally composed of coarse to very fine-grained sandstone, although locally decimeter-thick beds of extrabasinal pebbles occur at the base of some bedsets in the Nagri Formation. Bedsets are generally decimeters to several meters thick, and are bounded below by minor erosion surfaces that dip at 1° to 11° relative to the basal erosion surface of the storey. In some cases (particularly within the Nagri Formation) bedsets can be considerably thicker, and rarely they locally comprise the entire thickness of the storey (e.g. Fig. 12, Storey A; Fig. 14B). Bedsets typically fine upwards and show system-
ANCIENT RIVER SYSTEMS IN THE HIMALAYAN FOREDEEP
atic vertical variations in sedimentary structure, Grainsize fining and variation of sedimentary structures vertically across bedsets vary, however, depending on the vertical position within a storey (Figs. 7, 11, 13, 17). For example, bedsets measured lower in a storey generally have intraformational conglomerate at their base, grading upward into large-scale trough cross-stratified (i.e. dm to m thick cross-sets) a n d / o r planar stratified sandstone (Fig. 17, logs a, b and d). Bedsets measured higher in a storey generally decrease in grainsize upward and show an associated change in sedimentary structure from large-scale trough crossstrata (dm thick cross-sets) to cross-lamination (i.e. mm to cm thick cross-sets) or planar strata (Fig. 17, logs c, e and f). The thickness and mean grainsize of individual bedsets, the thickness of individual lithofacies beds within bedsets, and the thickness of individual sets of cross-stratification all tend to decrease upward within storeys. Near the top of storeys, bedsets are commonly dominated by planar stratified, cross-laminated or highly disrupted sandstone, Biogenic structures within major sandstone bodies are generally limited to either the upper parts of storeys, or the fine-grained tops of bedsets lower down within storeys (Figs. 7, 11, 13). Pronounced sediment disruption, obscuring all primary strata, occurs in the upper few meters of all storeys that are not truncated upwards. This represents over a quarter of the thickness of many storeys, but rarely more than half. Sandstone color commonly changes from light gray to pale reddish brown or yellowish gray where bioturbation/sediment disruption becomes pronounced. There are relatively few distinct types of trace fossils within these sandstone bodies, Nearly all are simple, unbranched, tube-shaped burrows of various circular diameters (1 mm to 1 dm, most 1-2 cm). Burrows have unlined walls, and meniscate, massive or mudclast breccia fills (Willis, 1992a). Lateral changes in bedset geometry, mean grainsize, sedimentary structures and paleocurrents across individual storeys vary greatly between different examples, and stacking patterns of storeys within these sandstone bodies is complex (see Figs. 6-14). Therefore, generalized
7
statements summarizing these variations are not useful. Examination of the following attributes in each sandstone body is integral to subsequent interpretations: (1) alignment of storeys across particular horizons within sandstone bodies; (2) occurrence of paleosols within some bodies, or a paleosol capping a body in some locations but truncated by upper storeys of the sandstone body in other locations; (3) truncation relationships between adjacent storeys; (4) along-strike changes in the thickness of individual storeys; (5) alongstrike changes in bedset inclination direction and steepness across each storey; (6) changes in lithofacies comprising bedsets in different parts of each storey; (7) changes in paleocurrent orientations relative to bedset dip trends observed within the plane of the outcrop across individual storeys; and (8) changes in paleocurrent orientations between adjacent storeys (see Figs. 6-14). In the interval documented at the Chinji-Nagri Fm. boundary (Fig. 3, Level B; Figs. 9-11) it is important to note the smaller size of storeys preserved to the west (left) side of the major sandstone body relative to those on the east (right) side. Note also the smaller overall size of storeys documented in the lower Chinji Fm. (Fig. 3, Level A; Figs. 6-8) relative to most examples documented within the upper Nagri (Fig. 3, Level C; Figs. 12-14).
Interpretationof major sandstone bodies These major sandstone bodies record deposition within the migrating channel segments of a large river. Each storey within these bodies represents a channel bar deposit (set of inclined bedsets) and normally an adjacent coarse-grained channel fill (concave upwards set of bedsets). This distinction between bar deposits and channel filling deposits is somewhat subjective because successive bedsets vary gradually across storeys. Thus a storey is the fundamental architectural element comprising these sandstone bodies. Storey margins with relatively low dips and generally concordant overlying bedsets record channel incision and subsequent bar formation and migration. Migration of thalweg areas and cutbanks in the channel produced basal erosion
8
surfaces of storeys. Intraformational breccia above basal erosion surfaces represent eroded cut bank material. Steep storey margins with adjacent channel filling (concave-upward to horizontal) bedsets record channel cutbanks, Individual bedsets represent deposits of a single flood event, and bounding surfaces of bedsets record a cross-section through a channel bar or fill. Bedset thicknesses reflect distances the bar and channel migrated, a n d / o r the amount of channel filling that occurred during individual deposition events. Normally only a few meters to a few tens of meters of lateral accretion (i.e. accretion in horizontal directions) parallel to the outcrop plane is indicated. Rarely in the Nagri Fm., however, bedsets record over 100 m of lateral accretion during an individual flood event (e.g. Fig. 14B). Along-strike variation of sedimentary textures and structures within individual bedsets reflect the distribution of grainsize and bedforms across a traverse of a channel bar or fill. Dominance of large-scale trough cross-strata within bedsets lower in storeys indicate most bars were covered by sinuous-crested dunes in the deeper parts of channels. Abundant planar strata and crosslamination in bedsets near the tops of storeys (particularly in the Chinji Fm.) reflect the abundance of upper stage plane beds in shallow bartop areas and ripples within areas of slower flow. Vertical variation of grainsize and sedimentary structures within a bedset reflects varying flow velocities during individual floods. Minor erosion surfaces at the base of bedsets reflect local scouring and modification of bar surface geometry during initial or high flood stages. Upward fining within bedsets and the common transition from large-scale trough cross-strata to planar strata or cross-lamination indicate decreasing flow velocities a n d / o r depths as the flood waned. Bioturbated bedset tops record disruption of the sediment by burrowing organisms and, in the upper parts of storeys, plant roots. Intense disruption and root casts within the upper quarter or so of storeys reflect bioturbation of bar deposits when emergent during low flow periods or following channel abandonment (Stanley and Fagerstrom, 1974; Bridge et al., 1986). If associated with low
B. WILLIS
flow periods, a relatively large decrease in channel flow depth between major floods is suggested. Variations in grainsize and sedimentary structures within bedsets traced upwards across a storey are comparable to those inferred from bar migration within many sandy, low to moderate sinuosity, single-channel and braided rivers (cf. Harms et al., 1963; Sarkar and Basumallick, 1968; Bernard et al., 1970; Bluck, 1971; Singh, 1972; Shelton and Noble, 1974; Jackson, 1976; Schwartz, 1978; Bridge and Jarvis, 1982; Bridge et al., 1986; Singh and Bhardwaj, 1991; also cf. review of Bridge, 1985). Many storeys in the Chinji Fm., however, contain more planar strata than is reported in most of these modern rivers (an exception is Harms et al., 1963). Exceptionally largescale sets of planar cross-strata, associated in modern rivers with chute, tributary or alternate bars, were not observed within the sandstone bodies documented (cf. Smith 1971; Cant and Walker, 1976; Cant, 1978). However, a few major sandstone bodies observed in other locations within the lower Chinji Formation locally contain a thick (3-4 m) cross-set at their base that: constitutes a quarter of the thickness of the storey; contains abundant intraformational breccia; and continues along strike for at least 100 m. More detailed study of such bodies is needed to determine whether such cross-sets simply represent migration of a large dune-scale bedform or whether they record deposition on larger-scale channel bed features.
Lateral variations across storeys and storey stacking patterns Each bedset exposed in an outcrop plane shows the geometry of a channel bar as viewed along a specific two-dimensional cross-section. Alongstrike variations in bedset geometry, internal lithofacies, and paleocurrent orientations across individual storeys reflect the interaction between: (1) temporal and spatial variations of bar geometry and facies; (2) channel bar migration, which controls facies preservation and the three-dimensional geometry of the storey; and (3)the outcrop orientation (see also Bluck, 1971, 1976; Jackson, 1976, 1978; Cant, 1978; Collinson, 1978; Bridge
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ANCIENT RIVER SYSTEMS IN THE HIMALAYAN FOREDEEP
21
B
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Fig. 8. Photos associated with Level A bedding diagram. (A) Minor channel-form sandstone body (base marked by dashed line) overlying the major sandstone body (Fig. 6; right side of Storey A). Scale 10 m. (B) Inclined bedsets (two examples arrowed) within storey decrease in inclination toward a coarse-grained channel-fill (Fig. 6; Storey O). Scale 5 m. (C) Complex coarse-grained channel-fill (base marked by dashed line) consisting of smaller channel-forms with inclined bedding surfaces (small arrow marks base of one channel form within larger scale channel fill). A layer of mudstone locally splits the sandstone body where the thick storey terminates (arrow to the right; dotted line marks start of mudstone beds) (see Fig. 6; left side of Storey P). Scale is 5 m. (D) The right side of a storey with convex upward bedding surfaces. Three small arrows mark one bedset that is inclined to the right relative to the base of the storey (see Fig. 6; right side of Storey R). Paleocurrents are generally out of the outcrop. Scale is 5 m long. (E) Lateral truncation of storeys within the major sandstone body (see Fig. 6; truncation of Storey S by T). Bedsets to the left are generally horizontal and disconformable with the inclined erosion surface (one example marked by dotted line). Bedsets to the right are inclined at a low angle to the right and are generally conformable with the inclined erosion surface (small arrow marks one such inclined bedset).
22
B.WILLIS
O
Fig. 9. Photos associated with Level B bedding diagram. (A) Sandstone body displaying three, vertically stacked storeys (see Fig. 10, Storeys B-D). Storeys are defined by major erosion surfaces (arrows and dashed lines) which bound sets of low angle inclined beds (e.g. beds within the middle storey are inclined to the right). Scale is 5 m. (B) Two-storey sandstone body (dashed lines mark the base of each storey; see Fig. 10, Storeys I and O between meters 750-850). Lower storey displays beds inclined to the left which are truncated upward at the erosional base of the overlying storey (arrow). Only the lower part of the upper storey is exposed within the cliff face. Scale is 5 m. (C) Complex vertical and lateral stacking of Storeys H, J and K within the sandstone body (dashed lines and small arrows mark the base of each storey; see Fig. 10, meters 1300-1700). Lower storey to the left shows inclined bedsets dipping to the left (2 examples marked by the letter b). Scale is 5 m. (D) Lateral truncation of Storey H by the steep margin of Storey O (Fig. 10, meters 1800-2100). Note inclined bedding surfaces dipping to the right within the thinner storey (arrows) and beds disconformable with the storey, margin in the thicker storey. Also note small mudstone-filled channel cutting top of the smaller channel sandstone. Scale is 10 m.
A N C I E N T R I V E R S YS T E MS 1N T H E HIMALAYAN F O R E D E E P
and Jarvis, 1982; Bridge et al., 1986; Willis, 1989, 1992, 1993a). Interpretations of bedset geometry within storeys depend on relating the orientation of these two-dimensional cross-sectional views (i.e. recorded by successive bedsets within a storey) to the changing three-dimensional geometry and migration pattern of an ancient channel bar surface (see discussions in Willis, 1989, 1993a). Such interpretations are analogous to reconstructing the shape of a three-dimensional object from a set of serial cross-sections (Fig. 18). However, here such reconstructions are more complicated because: (1) the bar changes shape over time, and thus along-strike variations of successive bedsets exposed in an outcrop reflect both temporal changes in bar shape and the changing position of the bar relative to the outcrop plane; (2) distances of bar migration vary during different depositional events, and thus successive bedsets do not provide an evenly spaced set of crosssections along the channel bar; (3) deposits from some areas on the bar may be removed by erosion as the bar and channel migrates, and thus a complete record of changing bar shape and position may not be preserved, Although individual outcrop exposures of channel bar deposits do not provide the complete information needed to reconstruct changes in bar geometry and position over time, fairly specific interpretations can be advanced by comparing outcrop variations to modern analogues. Specifically, across storey changes in the geometry, lithofacies, and paleocurrents of successive bedsets need to be directly compared to: (1) variations of channel bed topography, facies, and flow orientations observed in planview across modern river channels; and (2) records of how such channel beds change as bars grow and channels migrate. There are relatively few field studies that report spatial variations of channel bed topography and facies in the detail that is required for such comparisons; however, examples of varying utility include both studies of single-channel rivers (e.g. Fisk, 1944; Smith, 1971; Bluck, 1971, 1976; Jackson, 1975, 1976; Bridge and Jarvis, 1 9 8 2 ; Crowley, 1983; Ferguson and Werritty, 1983; Dietrich and Smith, 1983, 1984) and of multichannel rivers (Cant and Walker, 1976; Cant,
23
1978; Bridge et al., 1986; Bridge and Gabel, 1992). Only Bridge et al. (1986) attempted to combine detailed surveys of the river bed with information about how the channel bed changed through time to predict the characteristics of bedding and facies in vertical cross-sections. Unfortunately there are no studies that document bedding geometry and lithofacies variations in deposits of modern river channels with known bed geometry and migration history. As a starting point, interpretations advanced in this study assume that planviews of ancient channel segments were consistent with sine-generated curves of Langbein and Leopold (1966), that paleocurrent orientations generally record down-stream orientations of the channel planform at the point measured, and that channel segments progressively increased in sinuosity and migrated downstream as deposits formed. The trace of the outcrop plane relative to planform traces, and amounts the channel segment increased in sinuosity and migrated downstream are estimated by considering: (1) length of the storey relative to channel width; (2) vertical trends of lithofacies measured at several locations along each storey; (3) trends in bedset dip direction and steepness within the outcrop plain; and (4) paleocurrent variations across the storey relative to the outcrop plain (cf. Willis, 1989, 1993a). Initial assumptions used here may not be useful for deposits formed in channels with highly asyrnmetrical planforms, or for deposits of bars with steep downstream dipping facies (e.g. "down-stream accreted" deposits of Miall, 1988a). However, the basic methodology of interpreting outcrop exposures in terms of sequential cross-sections recording channel migration and bar growth remains the same regardless of channel planform or bar geometry (e.g. see channel bar facies model of Bridge et al., 1986; paleochannel reconstructions of Todd and Went, 1991, and Willis, 1992, 1993a; and Kocurek et al., 1992 for similar reconstructions of eolian dunes). The basic assertion here is that variations in bedset dip, lithofacies and paleocurrents across a storey provide more information about bar geometry and migration than that provided by any one inclined bedset viewed alone. Vertical and along-strike stacking of storeys
24
within sandstone bodies may reflect: (1)stacking of bars and channels in a single or multichannel river associated with channel migration, switching and cutoff within a channel belt; or (2) superposition of different channel belts (Figs. 19-21). Understanding how bedset geometry and facies vary within storeys and how storeys are stacked is fundamental to interpretations of channel sandstone bodies (e.g. Allen, 1983; Friend, 1983; Ramos and Sopena, 1983; Bridge and Gordon, 1985; Miall 1985, 1988a, b; Ramos et al., 1986; and others; also see many excellent examples in Miall and Tyler, 1992). Several lines of evidence suggest that vertical stacking of storeys within the sandstone bodies documented is due mainly to superposition of different channel belt deposits (cf. Figs. 6, 10, 12-19). Nearly complete preservation of most lower storeys is indicated by similar thickness and facies of upper complete storeys and lower truncated storeys (Figs. 6, 10, 12). Complete preservation of lower storeys is evidence against superposition of deposits formed in channel segments with different erosional depths in a single channel belt (cf. Figs. 19B, 19C). Relatively long periods of time separating vertically stacked storeys is indicated by local preservation of pedogenically modified sandstone a n d / o r overbank mudstones overlying some lower storeys within sandstone bodies (e.g. Fig. 6, Storeys K and L; Fig. 8D), or a paleosol that caps a sandstone body in some locations but is truncated along strike by storeys higher within the body in other locations (e.g. Fig. 10, terminating margins of Storeys B and O). Basal erosion surfaces and top surfaces of storeys do not rise more than a few meters along strike (Figs. 6, 10, 12). This suggests only a few meters of vertical accretion within channel belts and is evidence against the high vertical aggradation rates that would be needed to preserve nearly 10 m thick portions of underlying storeys by lateral bar migration (cf. Fig. 19A). Because most major sandstone bodies within the Chinji and Nagri Formations appear to be composed of multiple channel belt deposits separated from vertically adjacent sandstone bodies by thick intervals dorainated by mudstone, they record the concentra-
B. WILLIS
tion in time of different channel belts on the floodplain (Fig. 3; Willis, 1993b). Bedset geometries, lithofacies and paleocurrent variations within storeys, and lateral stacking patterns of storeys within each sandstone body, vary greatly, thus individual examples are interpreted separately below.
Detailed interpretation of major sandstone body in LevelA Sedimentologic variations observed in this sandstone body of the lower Chinji Fm. (Fig. 3, meters 365-380) are presented in Figs. 6-8. Preservation of pedogenically disrupted mudstone beds within this sandstone body (e.g. Fig. 6, above Storeys K and L) suggests that the six storeys aligned across the top of the sandstone body (Storeys E-J, centered on meters y = 30) were deposited in a different channel belt than those lower within the sandstone body (cf. Fig. 19D). Several other storeys locally cut into the paleosol that caps this sandstone body (i.e. Storeys A, B and D) and these storeys undoubtedly postdate all others within this body. Because these storeys above the capping paleosol are not connected to others along strike, they are discussed along with minor channel sandstone bodies (below). Storey L underlies the paleosol in this sandstone body (Fig. 6, between meters x = 100 to 1200, y = 15 to 20), whereas the other storeys in the lower part of the sandstone body truncate this paleosol. The well developed paleosol capping Storey L reflects a period of minimal deposition a n d / o r erosion on bar top areas for on the order of 103-104 years (Behrensmeyer, 1989; Willis, 1993b). Because such long periods of channel bar stability within channel belts seem unlikely (cf. bars in modern Himalayan foredeep basin rivers described by Coleman, 1969 and Bristow, 1987) and because paleocurrent orientations within Storey L are generally outcrop-parallel in contrast to dominantly outcrop-normal paleocurrent orientations within all other storeys lower within the sandstone body, it is most probable that the storey (L) represents a different channel belt that
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ANCIENT
RIVER SYSTEMS
IN T H E H I M A L A Y A N
FOREDEEP
was deposited well before the other storeys within this sandstone body. The other storeys exposed within the lower part of this major sandstone body can be placed into three groups, each separated along strike by the paleosol, and each group interpreted as a different channel belt (i.e. K, M-O, and P-U). Given the similar stratigraphic position of these three channel belt deposits and their similar mean paleocurrent orientations (generally outcrop-normal), they may record local anastomosing of the river system as is locally observed within modern Himalayan foredeep rivers. However, this hypothesis can not be tested within this single outcrop exposure, The topmost channel belt deposit (Storeys E-J) is exposed generally along its downstream a x i s . The truncation of Storey F by E suggests bar superposition by downstream migration (cf. Fig. 20E). All other storey truncations suggest adjacent bar deposits superimposed by channel switching or cutoff (cf. Figs. 20F, 20G). Storey F was clearly deposited first, followed successively by those to the right. Storey E may also have been deposited after Storey F, or it may have been coeval. The along-strike extent of each storey exposed parallel to paleoflow direction provides a minimum estimate of downstream bar length (0.5-1.5 km; cf. Fig. 20). The dominance of downstream dipping bedsets within all storeys in this upper channel belt records the preservation of only downstream ends of channel bars and the dominance of downstream bar migration (cf. Fig. 20). The fining of successive bedsets upwards within storeys agrees with sediment variations observed on the downstream ends of modern river bars (e.g. Jackson, 1976, 1978; Bridge, 1 9 8 5 ; Bridge et al., 1986). Bedset dips (2-4 °) indicate bar surfaces were inclined at low angles in the downstream direction (evidence for the steepness of bar surface dips in the upstream direction has been removed by erosion). In some storeys, bedsets are initially somewhat steeper (5-7 °) and they gradually decrease in dip along strike toward the channel fill (e.g. Storey J). This suggests downstream dips of bar surfaces decreased slightly as the bar grew and the channel migrated. This may be related to
43
a gradual increase in the channel width-depth ratio and a gradual loss of channel discharge over time. The along-strike extent of individual bedsets (100-300 m) records bar and channel widths measured highly oblique (almost parallel) to mean paleochannel fow. Storey thicknesses (7-13 m) provide an estimate of maximum bankfull channel depth. Channel fills (which contain bedsets that progressively decrease in dip, decrease in mean grainsize, and contain a greater abundance of planar and cross-lamination) record decreased flow depth and velocities as the channel was being abandoned. Since channel filling deposits are generally concordant with and nearly as coarse-grained as bar deposits, and fine-grained (mudstone) channel fills do not occur, ends of abandoned channel segments probably did not become plugged with sediment as channels filled (cf. Bridge et al., 1986). Instead, filling of a channel segment must have been associated with a very gradual diversion of discharge into an adjacent, contemporaneous channel. Paleocurrent variations within storeys (30-90 °) suggest sinuosities of at least 1.05 to 1.3 for individual channel segments. Mean paleocurrent directions of different storeys suggests an average sinuosity for this paleo-river reach between 1.1 and 1.2. Storey L is also exposed generally parallel to paleoflow. The upstream dipping bedsets formed by deposition on the upstream end of a bar, and suggest that channel migration was dominated by increasing channel sinuosity with little downstream bar translation (Figs. 20a, 20d). The downstream length of this bar had to be greater than the 1.1 km length of the storey because much of the downstream accreted part of this storey was removed by erosion associated with deposition of the downstream adjacent storey. Bedset dips suggest that bar surfaces were inclined at low angles in the upstream direction (downstream dipping bar surfaces were not preserved). Thickening of the storey and drop of the basal erosion surface to the left can be related to increasing maximum channel depth (from about 5 to 8 m) as the channel bend increased in sinuosity (see Bridge et al., 1986; Willis, 1989). Similarly, slightly coarser-grained bedsets dominated by large-scale trough cross-strata lower in the storey
44
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ANCIENT R I V E R SYSTEMS IN THE HIMALAYAN FOREDEEP
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METERS Fig. 15. Bedding diagram documenting major sandstone bodies exposed in Level C' (stratigraphically just below Level C) and adjacent mudstone-dominated successions. See Fig. 5 for explanation of symbols. The bedding diagram is presented without vertical exaggeration (upper box), and with 4 × vertical exaggeration (lower box).
to the left are probably related to deeper flow depths. Fining-upward grainsize trends across this storey contrast with the coarsening-upward deposits expected in upstream ends of channel bars (cf. Jackson, 1976; Bridge, 1977, 1978; and others),
However, most of the fine-grained bedsets higher in the storey to the left are clearly channel filling, and thus the vertical log probably does not reflect the normal distribution of grainsizes across the channel bar. Small channel forms capping the
Fig. 14. Photos associated with Level C bedding diagram. (A) Two 20 meter-thick sandstone bodies (left side of photo) become connected to the right to form a single 40 meter-thick sandstone body (see Fig. 12, Storeys A, C and G between meters x = 0 to 50). The sandstone bodies become connected where a third (intervening) storey starts (see arrow; dashed lines mark the base of each storey, dotted lines mark the top of storeys that are not truncated upwards). Scale is 10 m. (B) Sandstone body displaying inclined bedsets dipping to the right (see Fig. 12, Storey A between meters 800 and 900). Bedset bounding surfaces (large arrows) are spaced 80 m apart along strike (apparent shortening of this distance resulted from telephoto lens). Within this thick bedset, cross-set bounding surfaces rise gradually to the right (two examples marked by small arrows). Scale is 10 m. (C) Two-storey sandstone body (bases marked by dashed lines; Fig. 12, Storeys I and K between meters 400 and 500). Basal erosion surface of upper storey (arrow) rises laterally, more fully exposing lower storey. Note single large cross-set at the base of the upper storey (just above arrow). Scale is 10 m. (D) Multi-storey sandstone body (base of each storey marked with dashed lines). Middle storey starts just to the left of this photo and thickens towards the right truncating the storey below (large arrow; see Fig. 12, Storeys K and J between meters 600 and 700). Inclined bedsets with the middle storey (small arrow marks one example) dipping to the right are generally conformable with downcutting basal erosion surface. Scale is 5 m. (E) Two-storey sandstone body (bases of each storey are marked with dashed lines, tops are marked by dotted lines where they are not truncated upwards; Fig. 12, Storeys H and K between meters 1200 and 1400). Inclined bedsets to the right within the u p p e r storey end at the basal erosion surface (arrow marks one example) and successive bedsets decrease in dip into the coarse-grained channel fill (to the right). T h e upper storey ends just to the right of this photograph, but the lower storey continues for about another kilometer. Scale is 10 m. (F) Termination of major sandstone body in a coarse-grained channel fill (major erosion surfaces marked by dashed lines, dotted line marks top of the sandstone body; see Fig. 12, Storey K between meters 2500 and 2700). Note the smaller-scale channel-form (arrow) capping the larger-scale channel-filling sequence. Scale is 5 m.
'
•
46
8. WILLIS
larger channel fill record deposition within smaller channels (1-2 dm deep, meters wide) during the final stages of filling. Along-strike paleocurrent variations (SO”) indicate a channel segment sinuosity of at least 1.2. All of the other storeys lower within the sandstone sheet are exposed generally perpendicular to paleoflow. Storey K (far left) was deposited by a 7 m deep, 120 m wide channel, based on its
I
2
Fig. 16. Sedimentological
3
logs associated
4
thickness and 5” bedset dips. Storeys M and N are dominantly channel filhng, suggesting they were relatively short-lived and did not migrate to any great extent. Storey thicknesses of M and N, and bedset dips suggest that channel segments had maximum depths of at least 3 m and widths of about 60 m. Opposite dip directions of bedsets within these two storeys suggest channel cutoff (with Storey N deposited first) or two coeval
6
5
with Level C’ bedding
diagram.
7
See Fig. 5 for explanation
8
of symbols
ANCIENTRIVERSYSTEMSIN THEHIMALAYANFOREDEEP
47
braid channels (cf. Figs. 21B, 21C). In either case, flow diverged into these two channels and they developed sinuosity in an opposite sense, Storey O contains bedsets that dip in a similar direction to those in Storey N. The truncation of Storey N by O may record deposits from different channels superimposed due to channel switching (Fig. 21D). Alternatively, this bedding truncation may record the migration of a single bar that changed shape between two successive depositional events (related to varying discharge within the adjacent channel). Steeply inclined bedsets (11 °) adjacent to the left margin of Storey O traverse a distance of about 50 m along strike indicating that the channel had a width of about 75 m and a maximum depth of at least 8 m. Decrease of successive bedset dips along strike indicate progressive decrease in channel depth due to gradual channel abandonment as the bar migrated, but also progressively more oblique cross-sections across the channel. Most channel filling bedsets are planar stratified and relatively fine grained; however, a few bedsets are coarser
Variations Example I
0
5'o
grained and large-scale trough cross-stratified (e.g. Fig. 7, log 15, meters 20-23), and probably record erosional scouring of the nearly abandoned channel segment during unusually high discharge floods. Paleocurrent variations within inclined bedsets suggest that the channel segment developed a sinuosity of at least 1.1. Within the channel belt defined by Storeys P - U , Storey Q was deposited first on a bar that migrated to the right within the outcrop plane. Storey R was then deposited. The convex-upward bedding surfaces within Storey R indicate two coeval channels separated by a mid-channel bar (cf. Fig. 21A). Diverging paleocurrent directions and vertical trends in grainsize indicate a crosssection through the upstream end of the bar (cf. Bridge et al., 1986). Right-dipping bedsets and storey thickness on the right suggest that the channel segment became at least 9 m deep and at least 150 m wide. Truncation of Storeys Q and R by Storey P indicates either channel switching (cf. Fig. 21C) or a mid-channel-bar with a somewhat complex history of across-stream accretion (cf.
in bedset geometry
1~)0
150
and lithofacies
2bo
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36o
35o
Example 2
0 Scale
25 50 75 in meters (2x v e r t i c a l a
b
c
100 125 exaggeration d
e
150
f
Fig. 17. Bedset geometry and lithofacieswithin two storeys (stippled). Thick lines are major erosion surfaces, thin lines are minor erosion surfaces, dotted lines are gradational contacts with overlyingmudstone beds. Characteristic vertical lithofacies variations across bedsets (each decimeters to meters thick) in different parts of each storey are shown by letters on bedding diagrams and sedimentary logs (a-jr).
48
B. WILLIS
Fig. 21B). The total vertical extent of Storey P and the lateral extent of individual bedsets within this storey suggest a maximum channel depth of about 12 m and a width of about 200 m. T h e r e is little evidence of lateral accretion prior to channel filling. Smaller-scale channel forms within the larger channel fill signal reduced discharge and deposition on smaller bars and channels that were several meters deep and a few tens of meters wide (cf. Bridge et al., 1986; Bristow, 1987). In Storeys S and T, bedset dips (initially 7 °) progressively decrease and bedsets become concave upward adjacent to right hand storey margins, indicating periods of lateral bar migration followed by periods of channel filling. Because left hand storey margins and internal inclined bedsets have similar dip directions, they may be related to changes in bar shape due to discharge variations between depositional events. Thus
Storeys S to T could record deposition on the same migrating bar. Given downstream length estimates of bars for other storeys within this sandstone body (generally 0.5 to 1 km), deposition of Storeys R - T by migration of the same bar would require a single bar deposit to extend in a direction transverse to paleoflow for over a bar wavelength• Such a relationship would indicate highly sinuous channel segments (greater than 1.8), which is unlikely given the restricted paleocurrent variations observed within this sandstone body. Rather, paleocurrent shifts of 30° across the storey margins separating R, S and T, and channel filling bedsets on the right side of truncated storeys, suggest that these storeys formed on different bars and that deposits were superimposed by channel switching (cf. Fig. 21D). Because most of Storey U was removed by erosion, the timing of its deposition relative to others within this sandstone body is ambiguous•
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Fig. 18. A channel bend shown in planview is contoured in equal increments of depth (darker stippling reflects deeper areas). Depth contours follow Hooke (1974), but no specific scale is implied here. A set of sequential cross-sections across the channel bend are shown for 11 parallel traverses superimposedon the same cross-sectionalplane. The set of cross-sectionstogether reflect spatial changes in channel bed geometry. Similarly,inclined bedsets within a storey reflect a sequential set of cross-sectionsthrough a migratingchannel bar.
ANCIENT RIVER SYSTEMS IN THE HIMALAYAN FOREDEEP
A
B
49
B and C, suggest channel maximum depths and widths of 10-12 meters and 150-225 m, respectively. Low paleocurrent variations within storeys (less than 30°), but high paleocurrent variation
D
Fig. 19. Schematic examples of storeys vertically stacked within sandstone bodies. Thin lines show bedding surfaces, thick lines show major erosion surfaces, arrows show paleocurrent orientations relative to the outcrop plane (up is out of the outcrop). (A) Vertical superposition of storeys within a channel belt produced by rapid vertical accretion and downstream bar migration. The maximum thickness of the lower truncated storey records the amount of vertical accretion that occurred as the channel migrated by one bar length (see also Bridge, 1975). (B) Vertical superposition of storeys within a channel belt related to channels of varying size and thus erosional depth. (C) Vertical superposition of storeys within a channel belt related to stacking deposits from equal-sized channels where the former channel, that eroded to a greater depth, had a higher sinuosity than the later channel. A similar case would occur for channel bends of equal sinuosity where the outcrop plane exposes the two channel bends across different parts of their planform (see also Willis, 1989). (D) Vertical superposition of storeys from two different channel belt deposits by avulsion and subsequent erosion of the older channel belt by channels in the younger channel belt.
Detailed interpretation of major sandstone body in Level B This sandstone body is located at the ChinjiNagri boundary (Fig. 3, meters 885-908). As in Level A, vertical stacking of storeys is interpreted as largely due to vertical superposition of differe n t channel belt deposits. Storeys A - C each record the edge of a different, vertically stacked, channel belt deposit. Storeys D, E and H define a
fourth channel belt deposit. Storeys I-K, along the base of the sandstone body o v e r m o s t of the left side, constitute a fifth channel belt deposit, The thick storeys on the right side (O-Q) comprise a sixth channel belt deposit. Finally, Storey R constitutes a seventh channel belt deposit, The thicknesses (10-12 m) and the lateral extent of inclined bedsets (100-150 m) in Storeys A,
~ P ' ~ ~ ~ ~ ~" ~ . , d
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Fig. 20. Schematic examples of storeys stacked laterally and parallel to mean paleoflow within a channel belt. Each exampie contains a map showing channel planforms preserved through successive migration steps, and a cross-section showing trends in bedding geometry. Paleocurrent directions within planform maps are from left to right (see small arrow). Paleocurrents within cross-sections are given by arrows (up is into the cross-section plane). The larger context of channel segments and bars is not shown in these diagrams. Thus, in each case diagrams might represent a side (point) bar growing into the displayed channel segment, or the half of a mid-stream (braid) bar that is growing into the displayed channel segment. There is assumed to be no net vertical aggradation within the channel belt. (A-C) Along-stream stacking of bars deposited on opposite sides of a channel segment. Bar growth is by sinuosity increase (A), sinuosity increase and downstream bar migration (B), or downstream bar migration alone (C). In each case adjacent bar deposits display similar trends in bedding geometry, but paleocurrents vary in an opposite sense. Note that in (B) and (C) the upstream channel-filling storey margin truncates the downstream adjacent storey. (D and E) Along-stream stacking of bars deposited on the same side of a channel segment. Bar growth is by sinuosity increase (D), or downstream bar migration (E). In each case adjacent deposits display similar trends in bedding geometry and paleocurrents vary in the same sense. Note that in (E) the upstream channel-filling storey margin truncates the down-
stream adjacent storey.(F-H) Along-stream stacking of bars produced by channel cutoff. After cutoff the channel segment can develop sinuosity in an opposite sense (F), or the same sense (G, H) compared with sinuosity development before the cutoff. Note that in both cases the depositional margin of the downstream storey truncates the upstream adjacent storey.
50
B. WILLIS
b e t w e e n d i f f e r e n t c h a n n e l b e l t deposits, suggest
d o w n s t r e a m d i p p i n g b e d s e t s suggest that both
that c h a n n e l s w i t h i n c h a n n e l belts had low sinu-
bars were d e p o s i t e d o n the same side of a chan-
osity, b u t d i f f e r e n t c h a n n e l belts could traverse
nel s e g m e n t (as in the c e n t e r part of Fig. 20D)
this position o n the f l o o d p l a i n f r o m very d i f f e r e n t
a n d that the c h a n n e l s e g m e n t h a d a sinuosity of
directions,
at least 1.17. Storey thickness a n d the along-strike
Storeys D a n d H r e p r e s e n t a c h a n n e l belt
extent of i n c l i n e d b e d s e t s i n d i c a t e c h a n n e l maxi-
deposit exposed generally p a r a l l e l to paleoflow,
m u m d e p t h s of at least 10 m a n d b a r d o w n s t r e a m
T h e p a t t e r n of b e d s e t dips in Storey D record
l e n g t h s of at least 1.5 km. T h e t h i n n e r c h a n n e l -
d o w n s t r e a m a c c r e t i o n o n a b a r toward the west (left side), a n d u p s t r e a m a c c r e t i o n o n a n alongs t r e a m a d j a c e n t b a r to the east (right side). P r e s e r v a t i o n of u p s t r e a m - a c c r e t e d b e d s e t s a n d the conformity of b e d s e t s b e t w e e n the two b a r deposits suggest that c h a n n e l b a r s did n o t migrate d o w n s t r e a m as they grew a n d the c h a n n e l
increased in sinuosity (cf. Fig. 20D). Nearly horizontal b e d s e t s n e a r the c e n t e r of the storey record filling of a wide c h a n n e l cross-over area as the c h a n n e l was gradually a b a n d o n e d . F i n i n g - u p w a r d deposits in d o w n s t r e a m accreting b a r areas b u t little vertical v a r i a t i o n in grainsize in u p s t r e a m accreted areas agrees with t r e n d s o b s e r v e d in m o d e r n rivers (e.g. Bridge et al., 1986). L a t e r a l paleocurrent variations between upstream and
~ al" a~ ~------~:~J~~z:'~'~'~'~'~'~
"
" A~'~a'°"
bl• ~..~.../.~1~.-~:j:~ ,~~~[~.~..~
Ae" ~-~e~ P---~:~:!~ •
" E
yel /~e2 ~ "
~ ~
" ~
4f"
~
Fig. 21. Schematic examples of storeys stacked laterally and perpendicular to mean paleoflow within a channel belt. Each exampledisplays a map showing channel planforms preserved through successive migration steps, and a cross-section showing trends in bedding geometry and paleocurrents. Paleocurrent directions within planform maps are from left to right (small arrow). Paleocurrents within cross-sections are given by arrows (up is into the cross-sectional plane). (A) Growth of a mid-stream (braid) bar dominantly by sinuosity increase. Bedding surfaces dip in both directions away from the center of the bar deposit into the adjacent channel segments. Rightdipping and left-dipping beds define broad convex-upward surfaces.Paleocurrents are generally out of the cross-sectional plane. Paleocurrents are directed toward the center of the bar deposit if the cross-section is through the downstream end of the bar (a2), and away from the center if the cross-section is through the upstream end of the bar (al). (B) Growth of a mid-stream (braid) bar by dominantly down-stream translation. Deposits generally resemble (A); however, some differences occur. Cross-sections through upstream parts of the bar will contain beds that are disconformable with adjacent channel deposits (bl). Right-dipping and left-dipping beds will be disconformable if the bar migrates oblique to paleoflow or discharge is transferred from one channel segment to the other over time (b2). (C) Across-stream stacking of two channel bars due to cutoff. The channel segment developed sinuosity in an opposite sense before and after cutoff. Deposits are quite similar to those formed by migration of a mid-stream bar (compare with b2). (D) Across-stream stacking of two channel bars due to channel switching where the channel segment developed sinuosity in the same sense before and after cutoff. Beds dip in the same direction within adjacent bar deposits; however, a disconformity of bedding and a shift in paleocurrent orientations are expected at the storey margin. Where this paleocurrent variation is slight, the local bedding disconformity could be mistaken as one produced by slight change in channel alignment, possibly initiated by discharge variations between depositional events. (E, F) Acrossstream stacking of bars formed on the opposite sides of a channel segment. In both cases beds within adjacent storeys dip toward each other. Paleocurrent trends within storeys and whether beds within storeys are conformable with their mutual channel fill depends on the orientation of the cross-section and migration of the bars (compare el, e2 and f).
ANCIENT RIVER SYSTEMS IN THE HIMALAYAN FOREDEEP
filling storey (Fig. 10, Storey E) was probably deposited later by a smaller channel that flowed along a depression associated with the larger, underlying, abandoned channel fill. Interpreting the truncation of Storey D by H is somewhat complex because the outcrop plane steps backwards in a direction perpendicular to that shown in the bedding diagram by about 200 m between logs 18 and 19 (Figs. 10, 11, between meters x = 1700 and 1800). Had the outcrop not stepped backwards, the downstream end of the bar just described would probably have been exposed. Convex-upward bedsets within Storey H, with paleocurrents out of the outcrop, indicate two channel segments that were separated by a mid-channel bar. Inclined bedsets and storey thickness suggest the right channel segment was about 8 m deep and 85 m wide. Here, successive bedsets coarsen upward or have constant mean grainsize. Bedsets dipping towards the channel segment on the left are not as well exposed, but channel depths appear to have been only about 4 m, widths are probably also somewhat less than for the channel segment to the right, and bedsets are finer grained and generally channel filling. Storeys I - K along the base of the sandstone body record a channel belt deposit exposed generally perpendicular to paleoflow, and provide a measure of flow transverse channel belt width of about 1.8 kin. Bedsets within Storey I suggest deposition on a mid-stream bar (cf. Figs. 21A, 21B). The thickness of the storey, and the lateral extent of inclined bedsets to the left, indicate a channel segment with a maximum depth of 13 m and a width of about 200 m. Increased bedset clips and storey thickness toward the channel fill to the left record bar growth and development of higher channel sinuosities (cf. Bridge et al., 1986). Local bedding truncations, reflecting slight variations in the dip of adjacent bedsets, record temporally varying discharge. Coarsening-upward trends suggest a cross-section through the upstream end of the bar (cf. Fig. 21A; Bridge et al., 1986). However, a few coarse-grained bedsets capping the channel fill to the left probably reflect short periods of erosion associated with channel segment reactivation following its nearly complete abandonment. Limited variation in pa-
51
leocurrents along strike within left-dipping bedsets suggest, alternatively, a cross-section through the channel bend apex and that the bar did not migrate to any great extent downstream but grew as the adjacent channel segment increased in sinuosity. Storey J truncates Storey I, and must have formed later. Storeys J and K contain bedsets that dip in opposite directions away from a central location, indicating bar(s) that grew into two different channel segments (cf. Figs. 21B, 21C). Both storeys record channel segments about 10 m deep and 100-150 m wide. Local vertical superposition of these two bar deposits dictate that Storey J formed before Storey K; however, a period when both channel segments coexisted can not be ruled out. Determining whether different channel segments were active at the same time or whether they represent deposits separated in time is a persistent problem within these deposits. In this channel belt (Storeys I-K) it appears that two (or briefly three) channel segments could have been active at one time, suggesting a braiding index up to 2 or 3. The channel belt deposit represented by Storeys O - Q is exposed generally perpendicular to paleoflow and contains storeys that are substantially larger than those exposed to the left. Minor along-strike paleocurrent variations within these storeys suggest low channel segment sinuosities, but individual bar deposits exposed in flow transverse cross-sections do not provide a reliable estimate of sinuosity. The thickness of Storey O, and the along-strike extent of its bedsets, indicate channel depths of 23 m and widths of up to 450 m. The base of this storey varies in elevation along strike. The initial drop in the basal erosion surface records channel incision. The gradual rise in the basal erosion surface along strike to the left over the next kilometer is an opposite trend to that predicted by Bridge et al. (1986) and Willis (1989). This may indicate that there was at least 4 m of vertical aggradation within the channel belt as the bar migrated and grew about a kilometer in a direction perpendicular to paleoflow, or that the channel gradually became shallower over time (i.e. reflecting gradually decreasing discharge). Bedsets with lower dips and coarse grainsizes (sandstone) within the
52
channel fill record gradual abandonment of the channel over many flood events, Storey Q is similar in scale to Storey O but contains bedsets that dip in the opposite direction (cf. Figs. 21B, 21C). Storey Q clearly formed after deposition of Storey O and it may reflect deposition within the same channel segment, that switched position and grew in the opposite direction (cf. Fig. 21C). In Storey P, bedset dips of only 1-3 ° suggest a low channel segment sinuosity or a cross-section through the bend cross-over area. Beds with finer grainsize, lower dips, and cross-lamination within the channel fill indicate relatively slow flows as this channel segment was being abandoned. Its smaller size (less than 17 m deep, 200 m wide) and its position on the underlying bar deposits suggest it records a major chute channel cutting across the bar top. The presence of such a chute channel, and the coarse-grained nature of channel fills within these large storeys, suggest that the river had a braided pattern in general. However, the braiding index can not be estimated because the full extent of this channel belt could not be documented, nor could the total number of storeys aligned across stream be assessed, Along-strike variation in thickness, bedding geometry, and paleocurrent directions of Storey R indicate a single bar deposit exposed parallel to the channel belt axis (cf. Fig. 20A). Such along-strike variations suggest a maximum channel depth of 12 m, a width of 200 m, a downstream bar length of a least 1 km, and a sinuosity of at least 1.18. Because this storey is not associated with others along strike, it is interpreted along with minor channel deposits (below).
Detailed interpretation of major sandstone bodies in Level C Three major sandstone bodies are documented within this interval. A fourth, its lowermost part documented in the topmost portion of the bedding diagram (Fig. 12, meters y = 100 to 110) will not be discussed. One sandstone terminates along strike (Fig. 12, Storeys H - L between meters y = 50 and 70), whereas the others continue for the length of the exposures (3-4 km). The lowest sandstone body within this interval (Fig. 12,
B. WILLIS
Storeys A - G between meters y = 0 and 40) is clearly composed of two different superimposed channel belt deposits that are connected locally by several stratigraphically intermediate channelfilling storeys. Both channel-belt deposits are exposed generally parallel to paleoflow. The truncation of Storey A by B suggests superimposition of downstream adjacent channel bar deposits by channel switching (cf. Figs. 20F, 20G); however, similar bedset dip directions and paleocurrents in these storeys suggest that truncation may alternatively represent simply a bedding disconformity within a single bar deposit. Minor along-strike paleocurrent variations (less than 45 °) suggest that channel sinuosities were low, but were at least 1.05. Convex upward bedsets, and local upstream dipping bedsets, within Storey A indicate preservation of both upstream and downstream ends of a channel bar. This suggests that the channel bar did not migrate significantly downstream, although it clearly grew preferentially in the downstream direction (cf. Fig. 20). The downstream length of the bar was greater than 3 km. Thickness and lateral extent of bedsets indicate channel segment depths of about 22 m and widths of at least several hundred meters. The alongstrike distance separating adjacent inclined bedset bounding surfaces (many tens of meters) indicates that bars migrated or grew in the downstream direction by tens of meters to 100 m during individual flood events (cf. Coleman, 1969; Bristow, 1987). Storey G indicates a channel less than 15 m deep that is exposed parallel to paleoflow. Very low bedset dips make it difficult to document along-strike bedding trends, and inhibit interpretation of this storey in terms of deposition on channel bar(s). However, low bedset dips and low along-strike variation in paleocurrents suggest deposition within a low sinuosity channel segment. Thin channel forms cutting into the top of Storey G (i.e. Fig. 12, meters x = 200 to 800; meters x = 1200 to 1800) record shallow chute channels, or minor overbank channels flowing along the major channel course after it had been abandoned. Inclined bedsets within such channel forms, and along-strike paleocurrent variations suggest that some of these minor channels had
53
ANCIENT RIVER SYSTEMS IN T H E HIMALAYAN FOREDEEP
high sinuosities (at least 1.5). Channel-form storeys that locally connect the upper and lower channel belt deposits (Fig. 12, Storeys C-E) record channels that were active while the larger channel system had avulsed to another area on the floodplain. Their restricted lateral extent suggests they were relatively short-lived, and did not significantly increase in sinuosity nor migrate laterally. Storey F is thicker and may record the edge of a major channel belt that continues further to the east (left). Storey H records a single-storey channel-belt deposit exposed about 45° oblique to paleoflow, The along-strike extent of this storey suggests that the channel belt was less than 1 km wide. Geometry of bedsets within the storey record a cross-section though a mid-channel bar separating two channel segments, the one to the right about 15 m deep and 150 m wide, and the one to the left about 6 m deep and 100 m wide. Low variability in paleocurrent directions suggests low channel segment sinuosities (less than 1.1). The braiding index of this channel belt was 2. The lower part of this sandstone body (represented by Storeys l - L ) records deposition within a different channel belt. Storeys I - K record lateral stacking of channel bar deposits, whereas the vertical truncation of Storey L records vertical stacking of bar deposits within the channel belt (cf. Figs. 19B, 19C). Paleocurrents between Storeys I and K vary by almost 70° suggesting that channel segment sinuosities were locally as high as 1.3. Downstream-dipping bedsets suggest that channel bars migrated some distance downstream as they increased in sinuosity. Truncation of Storey K by Storey I records superposition of adjacent channel bar deposits due to downstream bar migration (cf. Figs. 20A-20E). Because paleocurrents within these storeys progress in a similar way (i.e. outcrop-parallel to outcrop-normal), they must represent deposition on the same side of a channel segment (cf. Fig. 20E). Storey J truncates Storey K only locally, thus it records channel cutoff across the upstream part of Storey I. Gradual thinning of Storey K along strike to the right, and the associated progression of paleocurrents back to outcrop-parallel orientations, suggest that Storey K is exposed along a cross-
section that passes through a cross-over area before crossing the cutbank margin to the right. Thicknesses of Storeys M and N suggest channel maximum depths of about 15 m, and maximum bedset dips of 4-5 ° suggest that channel widths were less than 250 m. The dip of bedsets in opposite directions indicates a highly oblique cross-section through two channel segments that developed sinuosities in the opposite sense (cf. Figs. 21A-21C). It is not possible to assess whether channel segments were coeval braid channels or represent channel switching of a single channel segment. Preservation of upstreamdipping bedsets within the left-hand storey suggests that bar migration was dominantly by expansion (sinuosity increase). Along-strike variations in paleocurrents suggest channel sinuosities between 1.1 and 1.2.
Paleochannel reconstructions Qualitative reconstructions Generally braided channel patterns within channel belts are suggested by the relatively large number of storeys within individual channel belt deposits, limited along-strike extent of individual storeys exposed perpendicular to paleoflow relative to estimates of bar downstream length, dominance of coarse-grained channel fills (i.e. sandstone- as opposed to mudstone-dominated), low paleocurrent variation within individual storeys, local preservation of convex-upwards bedset surfaces recording a mid-stream bar, and abundance evidence for bar deposit superposition due to channel switching. Channel segments were low sinuosity (1.0-1.3), and channel bars dominantly migrated downstream (but not always). Channel cutoff and switching occurred frequently relative to channel migration rates, but individual channel segments were cutoff gradually and abandoned over many flood events. Bars all had similar geometries, characterized by low-angle channel bed dips (2-4 °) in upstream and downstream directions, and slightly steeper (7-11 °) channel bed dips in cross-stream directions. Lithofacies variations across storeys of the Chinji Formation suggest that during floods much of the channel bed was covered by plane beds
54
B. WILLIS
when channels were straighter (and shallower), but dunes dominated channel bed surfaces in deeper parts of more sinuous channel segments, Lithofacies variations across larger storeys within the Nagri Formation suggest that dunes covered a greater proportion of areas within these chartnels. Exposure of the full width of channel belts displayed perpendicular to paleoflow provide a more complete view of channel bed geometry and lithofacies measured across individual ancient rivers. For example, Storeys P - T (Figs. 6, 7) indicate a river that was about 2 km across and contained two active channel segments (i.e. those that deposited Storeys P and T), thus the river had a braiding index of 2 (as defined by Brice, 1964). Bedsets within these two storeys (P and T) suggest that most of this distance consisted of very shallow bar top areas covered dominantly by plane beds, but by ripples during falling flows. Patterns of sediment disruption suggest that much of this bar top area was exposed during low flows, There are significant differences in the size of storeys compared between different channel belts within some sandstone bodies (e.g. Fig. 10), and between sandstone bodies observed along different stratigraphic intervals (cf. Figs. 6 and 1 2 ) . Two distinct scales are recognized, larger ones (15-30 m thick) are restricted to the Nagri Formation (see east (left) side of Fig. 10, Storeys O-Q; and Fig. 12, Storeys A-B, G, I-N). These two distinct scales of channel deposits are interpreted to reflect deposition within two coeval rivers of different magnitude (Willis, 1993b). More detailed estimates of the dimensions and discharge of these two rivers are presented below,
Quantitative reconstructions Paleochannel geometry and hydraulics of several storeys within these sandstone bodies were reconstructed by comparison with a physical model that predicts the sedimentology of singlechannel bar deposits (Bridge, 1977, 1978, 1 9 8 2 ) . Full descriptions of paleochannel reconstruction methods are presented by Bridge (1978), Bridge and Diemer (1983) and Willis (1992, 1993a); thus only a brief summary will be given here. Given certain initial hydraulic and geometric conditions, the model predicts the local flow pattern, bar
topography, and variation in mean grainsize across the bed at channel forming (bankfull) discharge within channel bends of regular planform. Reconstruction of a paleochannel involves modifying the initial conditions of the model until it simulates the observed bar deposit as closely as possible. Many of these initial conditions can be determined directly from the deposits. For instance, the geometry of successive inclined bedset bounding surfaces, dimensions of channel fills, and lateral variations in storey thickness provide an estimate of the varying paleochannel geometry within a single two-dimensional cross-section. In past reconstructions such geometric parameters were transformed by a linear projection into a plane orthogonal to paleoflow before comparisons were made with simulated bar deposits (e.g. Bridge 1978; Gardner, 1983; Bridge and Diemer, 1983; Bridge and Gordon, 1985; and others). This method has been shown to provide unreliable estimates of channel width and bar shape (see Willis, 1989). Here a new computer program is used that allows comparison of outcrop information with simulated bars in any specified crosssection orientation (see also Willis, 1989, 1992, 1993a).Channel bend sinuosity and wavelength are estimated by considering the along-strike extent of the storey and paleocurrent variations within the storey. These estimates are checked against empirical equations relating channel width and depth to channel planform (as in Bridge, 1978). Hydraulic parameters such as mean water surface slope and friction coefficients are estimated from variations in grainsize, sedimentary structures and channel depth. Once initial estimates are obtained, repeated simulations are performed and estimates of model input parameters are refined to improve the fit of model output to data. Precise error limits associated with this methodology are difficult to estimate. Errors are associated with differences between real and modeled three-dimensional bar geometries, and inherent uncertainties associated with using depth/grainsize/sedimentary-structure relationships to estimate flow velocities. Reconstructions of the same channel deposit using maximum and
ANCIENTRIVERSYSTEMSIN THE HIMALAYANFOREDEEP
55
minimum reasonable input estimates generally produced discharge estimates that varied by less than 25% of the total value. This bar model was developed for single-channel meandering rivers (i.e. point bars) and it has been tested successfully against both modern rivers (Bridge and Jarvis, 1982; Bridge, 1984a) and ancient deposits (Bridge and Diemer, 1983; Bridge and Gordon, 1985; Willis, 1992, 1993a). The model also appears to provide adequate agreement with single curved channel segments in braided rivers where bars do not have steep down-stream dipping faces (Bridge and Gabel, 1992). Significantly, bedding geometries predicted
by this model are also broadly similar to those observed within the deposits documented here (cf. Figs. 6, 10 and 12 with Willis, 1989). It must be emphasized, however, that reconstructions of multi-channel rivers will provide only an estimate of conditions within individual channel segments. Therefore, dimensions of the entire river channel and details of its hydraulics can only be obtained if the channel braiding index can be determined. Tables 1-3 summarize estimates of paleochannel geometry and hydraulics reconstructed from storeys in each of the three stratigraphic intervals documented. Two distinct scales of channel deposits are recognized within these deposits (as
TABLE 1 Paleochannel reconstructions, Level A (obstacle sandstone) Storey number
1
2
3
Apparent bar width (m)
110
100
130
Estimated channel bankfull width (m)
160
140
Mean depth (m)
4.5
Maximum depth (m)
4
5
6
7
8
45
135
120
95
105
200
70
200
180
140
160
5.0
4.5
4.0
6.5
5.5
5.0
7.0
7.5
8.5
8.5
8.0
11
10.5
10
13
Centerline channel length in one bend (m)
1110
1010
1470
490
1300
1360
1160
1190
Centerline sinuosity
1.1
1.1
1.2
1.15
1.1
1.2
1.3
1.2
Mean Darcy Weisbach type friction coefficient
.04
.06
.06
.06
.06
.06
.06
.06
Dynamic friction coefficient
0.6
0.6
0.6
0.6
0.6
0.6
0.6
0.6
Mean friction slope (ram/m)
.059
.063
.055
.064
.051
.050
.052
.049
Mean velocity (m/s)
0.59
0.63
0.57
0.59
0.62
0.61
0.62
0.64
Bankfull discharge (m3/s)
430
440
510
170
810
610
430
710
Position of storeys within Fig. 6. 1: Storey E, meters x = 0-600, y = 20-30.2: Storey K, meters x = 0-250, y = 15-22.3: Storey L, meters x = 110-1200, y = 12-21.4: Storey O, meters x = 860-1140, y = 17-25.5: Storey P, meters x = 1170-1520, y = 17-29.6: Storey R, meters x = 1800-2140, y = 18.5-30. 7: Storey T, meters x = 2020-2550, y = 18-27.5.8: Storey J, meters x = 2550-4000, y = 18-31.
56
B. WILLIS
discussed above, and in Willis, 1993b), channel dimensions and flow reconstructed from these two different scale deposits are discussed separately below (reconstructions of larger channels are marked by an asterisk in Tables 2 and 3). Smaller channels were reconstructed from the lower Chinji Fm. (Fig. 3, Level A; Table 1), the west (left) side of the interval documented at the Chinji-Nagri boundary (Fig. 3, Level B; Table 2) and Storey H documented within the upper Nagri Formation (Fig. 3, Level C; Table 3). Estimated channel dimensions and flow parameters reconstructed from these smaller-scale deposits are: (1) depths 7.5-13 m; (2) widths 70 to 200 m; (3) downstream lengths of individual curved channel
segments 0.5 to 1.5 kin; (4) sinuosities of individual curved channel segments 1.0-1.3; (5) downstream flow velocities 0.60-0.80 m / s ; (6) downstream slopes 0.05 and 0.08 m m / m ; and (7) discharge of individual channel segments 150 and 1050 m 3 / s (although most channel segments had discharges between 500 and 800 m3/s). Full river discharges associated with these smaller-scale channel deposits can be estimated by adding discharges of all potentially coeval channels within an individual channel belt (this assumes only one channel belt was active, i.e. the system was not anastomosing). Three channel belts provide such estimates: (1) Storeys P - T in the Lower Chinji Formation (Fig. 6) indicate a flow-transverse
TABLE 2 Paleochannel reconstructions, Level B (GB 1 sandstone) Storey number
1
2
3
4
Apparent bar width (m)
140
130
90
100
Estimated channel bankfull width (m)
205
195
140
Mean depth (m)
6.0
7.0
M a x i m u m depth (m)
12
Centerline channel length in one bend (m)
5
6
7"
8
55
105
280
140
150
80
160
400
205
7.0
4.5
4.0
5.0
12
6.5
13
13
9
7.5
10
23
12
1430
1590
1360
1110
550
1220
2900
1480
Centerline sinuosity
1.1
1.3
1.25
1.15
1.1
1.25
1,1
1.18
Mean Darcy Weisbach type friction coefficient
.04
.06
.04
,04
.06
.06
.06
.04
Dynamic friction coefficient
0.6
0.6
0.6
0.6
0.6
0.6
0.6
0.6
Mean friction slope (ram/m)
.05
.05
.05
.06
.08
.06
.03
.04
Mean velocity (m/s)
0.59
0.76
0.70
0.65
0.60
0.62
0.72
0.71
Bankfull discharge (m3/s)
730
1040
680
440
200
480
3560
950
Position of storeys within Fig. 10. 1: Storey C, meters x = 0-500, y = 31-45.2: Storey I, meters x = 400-600, y = 15-27. 3: Storey I, meters x = 700-900, y = 16-25. 4: Storey J, meters x = 1300-1700, y = 26-35.5: Storey H, meters x = 1600-1950, y = 23-32. 6: Storey K, meters x = 1550-1960, y = 11-21.7: Storey O, meters x = 2000-3000, y = 20-43.8: Storey R, meters x = 2800-3900, y = 40-53.
57
A N C I E N T R I V E R SYSTEMS IN T H E H I M A L A Y A N F O R E D E E P
channel belt width of about 2 km, and two coeval channels with a combined discharge just under 1500 m 3 / s (reconstructions 5 and 7 in Table 1); (2) Storeys I - J at the Chinji-Nagri boundary (Fig. 10) indicate a flow-transverse channel belt width of about 1.8 km, channel braiding index of about 2 (but perhaps temporarily 3), and full channel belt discharge of about 1200 m 3 / s (reconstructions 3 and 4 in Table 2); (3) finally, Storey H in the u p p e r Nagri (Fig. 12) records a braid bar separating two channel segments with a combined width of just over a kilometer (measured 40 ° oblique to paleo-flow) with a combined discharge of just over 1350 m 3 / s (reconstructions 4 and 5 in Table 3).
Larger channel deposits were observed in the right side of the sandstone body documented at the Chinji-Nagri Formation boundary (Fig. 10, Storeys O - Q ) , and within the sandstone bodies documented in the u p p e r Nagri Formation (Fig. 12, Storeys A - B , G, H - N ) . Estimated channel dimensions and flow p a r a m e t e r s reconstructed from these larger-scale deposits are: (1) depths 8-23 m; (2) widths 180-400 m; (3) downstream lengths of individual curved channel segments 1.4 to 2.9 km; (4) sinuosities of individual curved channel segments 1.1-1.2; (5) downstream flow velocities 0.7-0.8 m / s ; (6) downstream slopes 0.03 and 0.06 r a m / m ; and (7) discharge of individual channel segments 1390 and 3560 m 3 / s
TABLE 3 Paleochannel reconstructions, Level C (GB 112 and 13 sandstones) Storey number
1"
2"
3"
4
5
6"
App~u'ent bar width (m)
180
120
230
110
150
230
Estimated channel bankfull width (m)
260
180
340
170
230
340
Mean depth (m)
9.5
10
8.0
3.5
7.5
11
Maximum channel depth
17
19
16
6
15
20
Centerline channel length in one bend (m)
1660
1390
2150
2800
2800
2150
Centerline sinuosity
1.15
1.2
1.15
1.15
1.2
1.1
Mean Darcy Weisbach type friction coefficient
.06
.06
.06
.06
.06
.06
Dynamic friction coefficient
0.6
0.6
0.6
0.6
0.6
0.6
Mean friction slope (mm/m)
.05
.047
.053
.06
.072
.045
Mean velocity (m/s)
0.76
0.76
0.74
0.52
0.62
0.79
Bankfull discharge (m3/s)
1920
1390
2180
310
1050
3390
Position of storeys within Fig. 12. 1: Storey I, meters x = 0-300, y = 54-70. 2: Storey J, meters x = 540-920, y = 50-69. 3: Storey K, meters x = 1300-1700, y = 52-65. 4: Storey H, meters x = 1000-1200, y = 60-75. 5: Storey H, meters x = 800-1400, y = 19-39. 6: Storey A, meters x = 500-700, y = 69-75.
58
B. WILLIS
O
ANCIENT RIVER SYSTEMSIN THE HIMALAYANFOREDEEP
(discharges between 2000 and 3000 m3/s are most common). Exposures are not extensive enough to estimate the full channel-belt widths nor braiding indexes for these large channel deposits, and thus full river discharges can not be estimated in detail. However, sedimentological evidence suggests that these larger channel belts had braided channel patterns (see above) and comparisons to modern large rivers of the Himalayan foredeep basin (Coleman, 1969; Bristow, 1987; Wells and Dorr, 1987; Singh and Bhardwaj, 1991) suggest braiding indexes less than 4 or 5. Thus full channel belt discharges recorded by these thicker channel belt deposits were probably on the order of 10,000 m3/s (i.e. at least a factor of four larger than reconstructed for thinner channel belts described above), Minor sandstone- and mudstone-dominated bodies
Strata separating major sandstone bodies are mudstone-dominated overall, but contain small laterally discontinuous sandstone bodies (Figs. 6-16). These strata can be divided into: (1) minor channel-form bodies, (2) lobate and wedge-shaped bodies, and (3)laminated mudstone bodies. These sediment bodies can be grouped into meters to ten meter thick sequences delineated by alternation of sediments containing primary stratifica-
59
tion with highly disrupted sediments defining paleosols. Sediment bodies and paleosol bounded sequences are described and interpreted separately below.
Minor channel-form bodies Description Most minor channel-form bodies are dominantly sandstone, are generally one to a few meters thick, and are continuous along strike for several tens of meters to less than a kilometer (e.g. Fig. 6, Storeys B, V and others; Fig. 10, Storeys M, N and others; Fig. 15 between meters 10 and 15). A few are thicker, however, and have dimensions similar to individual storeys within major sandstone bodies (e.g. Fig. 6, Storeys C and D; Fig. 10, Storey R). Most bodies are composed of a single storey in the plane of the outcrop, are flat topped with a concave-upward basal erosion surface, and have apparent width to thickness ratios between 10 and 50 (cf. ribbon bodies of Friend, 1983). However, some channel bodies are more extensive along strike and rarely contain two adjacent storeys (e.g. Fig. 6, Storey D). These have greater apparent width to thickness ratios (over 100) and are bounded below by an erosion surface that is generally horizontal for some distance between more steeply dipping margins (cf.
Fig. 22. Minor sandstone bodies. (A) Minor channel-form sandstone body (base marked by dashed line, top by dotted line) displaying inclined bedsets that grade upward to mudstone (arrows mark two examples). Scale is 2 m. Outcrop is located within the lower Chinji Formation (Fig. 3, stratigraphically adjacent to meters 498-502). (B) Minor channel-form sandstone body (base marked by dashed line) with inclined beds which pass into a mudstone-dominated channel fill (boundaries of channel fill marked by dotted lines). Scale is 10 m. Outcrop is located within the lower Nagri Formation (Fig. 3, stratigraphically adjacent to meters 1231-1235). (C) Minor sandstone body (base marked by dashed line) with inclined beds (e.g. arrow) that fill a channel-form preferentially from the right. Scale is 5 m. Outcrop is located within the upper most Chinji Formation (Fig. 3, stratigraphically adjacent to meters 863-866). (D) Minor channel-form body (base marked by dashed line) that contains left dipping bedsets and fines from dominantly sandstone to the right to dominantly mudstone to the left. Scale is 1 m. Outcrop is located within the lower Chinji Formation (Fig. 3, stratigraphically adjacent to meters 340-345). (E) Minor lobate sandstone body (base marked by dashed line, top marked by dotted line) with low angle inclined beds dipping in the direction the wedge thins (to the right). Small arrows mark one bedset that drops on elevation relative to the base of the wedge-shaped body about 2 m over 50 m. Scale is 5 m. Outcrop is located within the lower Chinji Formation (Fig. 3, stratigraphically adjacent to about meters 1497-1501). Figure 24 documents sediment variations along this sandstone body in more detail. (F) Wedge-shaped sandstone body that passes laterally into a major sandstone body (dashed lines mark the base of the major sandstone body and the base of the wedge-shaped body). Note inclined bedsets in the wedge-shaped body, dipping in the direction the wedge thins (to the right). Scale is 5 m. Outcrop is located within the upper Nagri Formation (Fig. 3, stratigraphically adjacent to meters 1497-1502). Sediment variations along this sandstone body are documented in Fig. 16, logs 2-6, meters 35-45.
60
B. WILLIS
0
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ANCIENT RIVER SYSTEMSIN THE HIMALAYANFOREDEEP
sheet bodies of Friend, 1983). Bodies with greater extent along strike relative to their thickness (i.e. sheets) tend to have higher proportions of inclined bedsets relative to channel-filling (concaveupward to horizontal) bedsets (e.g. Fig. 22A) and more commonly end in mudstone-dominated channel-fills (e.g. Fig. 22B). Minor channel sandstone bodies comprise less than 10% of deposits that separate major sandstone bodies, and across most stratigraphic horizons they are spaced many kilometers apart, Erosion surfaces at the base of these bodies have decimeter-scale relief, and are overlain by intraformational conglomerates of mud clasts, pedogenic nodules and (in places)vertebrate fossils, Oncolites (up to centimeters in diameter and concentrically laminated about a central clast) occur within such basal conglomerates locally in adjacent areas of the Chinji Formation. Cylindrical stromatolites (decimeters to a meter in diameter) were observed along the base of one minor channel body in the Nagri Formation (Behrensmeyer, 1987; Willis, 1992). The base of many channel-form bodies correspond to upper parts of paleosols or to nodular calcite layers within paleosols, Bedsets within minor channel sandstone bodies are fine- to very fine-grained, erosionally based, centimeters to decimeters in thickness, and fine upward (Figs. 7, 11, 13, 16). Lower in a body, bedsets grade upward from intraformational conglomerates into cross-laminated or less commonly large-scale trough cross-stratified
61
sandstone. Higher in a body, bedsets are mainly cross-laminated or intensely disrupted sandstone. Bedsets fine upward to mudstone in the upper parts of some bodies (e.g. Fig. 22A). Rarely, in the Chinji Formation, symmetrical ripple marks cap bedsets. In general, successive vertically superjacent bedsets fine and thin toward the top of the sandstone. Sediment disruption is extensive within most minor channel bodies, and in some cases stratification within the body is completely disrupted (particularly in the Nagri Formation). The disruption of sediment is complex and can not always be related to distinct structures; however, less than 10 mm diameter, tube-shaped burrows are common and are generally similar to those within upper parts of major sandstone bodies. Vertical cracks (up to 1 m in depth) and root casts are also common near the top of these minor bodies and throughout channel filling parts. In bodies with mainly channel filling bedsets (i.e. ribbon-shaped), a few bedsets are inclined concordant with one channel margin but most become concave upward, or horizontal and discordant with the other margin (e.g. Fig. 6, Storeys B and V; Fig. 22C). Paleocurrents in such bodies are generally perpendicular to the outcrop and vary little between bedsets (less than 30°). Sandstone bodies that extend for greater distances along strike, and are similarly exposed generally perpendicular to paleo-flow, display comparable lateral changes in bedset geometry to those just described but are dominated by inclined bedsets adjacent to the bedset-concordant margin that
Fig. 23. Details of a minor channel sandstone body exposed generally parallel to paleoflow. The sandstone body is exposed on a series of extensive bedding surfaces and within adjacent inclined rock faces. The outcrop is located within the upper Chinji Formation (Fig. 3, stratigraphically adjacent to meters 785-800). This sandstone body constitutes paleontology site Y76, which is one of the most prolific vertebrate fossil localities ever discovered in the Siwalik Group (for discussion of taphonomy see Behrensmeyer, 1987). (A) Map of paleocurrent orientations derived from large-scale trough cross-strata and parting lineations on planar strata measured along bedding surfaces near the base of sandstone body. Paleocurrents progress from southerly orientations on the east end of the exposure to westerly orientations on the west end of the exposure (arrows, south to top of page). Circled numbers indicate positions of measured sections. (B) Extensive bedding surfaces exposing edges of individual beds which curve around in a similar way as paleocurrent orientations (A). (C) Sedimentological logs measured at selected locations along exposure (see A). See Fig. 5 for explanation of symbols. Arrows represent paleocurrent orientations (up to the north). (D) Photo to the north of steep slope exposure between sections 4 and 5. Inclined bedding surfaces dip from top of the sandstone to its base. Beds grade to mudstone near the top of the body. Paleocurrents are toward the observer. (E) Schematic bedding diagram of sandstone body with two times vertical exaggeration. Beds dip in both directions away from the center of the sandstone toward laterally adjacent, fine-grained channel fills. Numbers mark positions of measured sections. Arrows display paleocurrent orientations relative to the orientation of the vertical exposure and a line connecting measured sections.
62
13.WILLIS
dip toward the channel filling margin. Individual inclined bedsets normally extend along strike for only a few tens of meters. These more extensive bodies also commonly display higher along-strike paleocurrent variations (up to 70°). Where channel sandstone bodies are exposed generally parallel to paleoflow, bedsets either (1) dip in both directions away from the center of the sandstone body toward marginal channel fills (e.g. Fig. 6, Storey C; Fig. 10, Storey R; see also unusually well exposed example documented in Fig. 23), or (2) are dominated by bedsets that dip toward a marginal channel fill at the down-stream end of the body and are truncated by a marginal channel fill at the upstream end of the body (e.g. Fig. 6, Storey D). Such bodies generally show large along-strike paleocurrent variations (commonly 100_150o). The character of sediments within minor channel bodies varies somewhat compared between examples observed within the Chinji and the Nagri Formations. In the Nagri Formation, most examples contain bedsets entirely of sandstone that are completely disrupted throughout the body. In the Chinji Formation, most contain some bedsets that fine upwards to mudstone; disruption of stratification becomes pervasive only in the upper half or so of the body; bodies showing extensive variation of paleocurrents (i.e. greater than 90 °) along strike are more common; some examples contain oncolites, and vertebrate fossil remains are more common,
Interpretation Most minor channel-form sandstone bodies are interpreted as deposits of crevasse, minor tributary, or floodbasin drainage channels (cf. Fisk, 1944, 1947; Smith et al., 1989; Tye and Coleman, 1989). However, where the thicknesses of these bodies are similar to those of individual storeys within major sandstone bodies (e.g. Fig. 6, Storeys B, C, D, V and others; Fig. 10, Storey R), interpretations are more problematical. Such exampies exposed parallel to paleoflow (e.g. Fig. 6, Storeys C and D; Fig. 10, Storey R) may record single-channel segments within major channel belts cut tangentially by the outcrop plane; however, this interpretation can not be tested within
single outcrop exposures. Thick channel bodies exposed perpendicular to paleoflow (e.g. Fig. 6, Storeys B and V; Fig. 10, Storeys M and N) that are not connected to a major channel belt deposit along strike, must be related to either large crevasse channels or to major tributary channels. The low proportion of minor channel sandstones within these deposits (less than 10%) suggests a paucity of overbank channels, and that these minor channels were commonly spaced at least kilometers apart on the floodplain. Nearly all channel bodies are composed of a single storey in the outcrop plane suggesting that they were deposited by single-channel river systerns. These channels were generally a few meters to 10 m deep, and their apparent channel widths varied from 10 m to less than 100 m. Most of these channels have generally SSE paleocurrent directions that are somewhat oblique to those within subjacent major sandstone bodies, suggesting that they may have been flowing gradually toward or away from a major channel belt. Many channel deposits are directly underlain by carbonate-rich paleosol horizons or laminated claystones suggesting that these sediments were more resistant to erosion (cf. Gibling and Rust, 1990). Bedsets (decimeters to centimeters thick) within bodies clearly indicate that deposition on bars and in adjacent fills occurred gradually over many flood events. Thin bedsets, dominance of cross-lamination, and fine to very fine grainsizes within most bodies suggest generally slower flows and lower deposition rates than within major channel systems. Mudstones capping some bedsets suggest very slow channel flows between flood events. Rare bedsets capped by symmetrical rippie marks in the Chinji Fm. indicate that sediments were locally reworked by waves between depositional events. Oncolites observed within the Chinji Fro. must have formed along basal erosion surfaces of channels (as they are never observed within adjacent deposits), and record the ponding of water and low deposition rates within channel fills between flood events (Schafer and Stapf, 1978; Nickel, 1983). The singular occurrence of cylindrical stromatolites at the base of a minor channel sandstone within the Nagri Formation similarly reflects ponding of water and the accre-
63
A N C I E N T R IVE R SYSTEMS IN T H E H I M A L A Y A N F O R E D E E P
tion of carbonate around submerged logs or tree trunks (cf. Kraus, 1987b). In general, sandstonedominated channel fills suggest that overbank channels continued to be conduits for flood waters until they where completely filled, whereas bodies with mudstone fills must have been abandoned completely and then gradually filled with suspended sediment deposited from slow moving flood waters. Pervasive sediment disruption by mudcracks and root casts within many bodies suggest that much of the channel area was emergent between flood events. However, restriction of completely disrupted sediments to upper parts of most bodies within the Chinji Fm., suggest that most of these channels contained some flow throughout the year until they began to be abandoned. Given the abundance of vertebrate fossils, much of the intense disruption within these channel deposits may have been due to trampling by large mammals (cf. Wells and Dorr, 1987). However, bioturbation by roots and burrowing organisms, and desiccation of the sediment clearly played an important role in sediment disruption, Trends of bedset inclination, lithofacies, and paleocurrents across bodies can be interpreted
and fill deposits exposed generally parallel to mean downstream direction. The proportion of upstream to downstream dipping bedsets within storeys exposed parallel to paleoflow reflects the distance the bar migrated downstream as it grew (cf. Fig. 6, Storeys C and D, and Fig. 20) and along-strike variations in paleocurrent directions reflect the channel sinuosity (e.g. about 1.3 for Storeys C and D). In many cases differences in minor channel body geometry (i.e. ribbon versus sheets) probably reflect only the orientation of the outcrop plane across the channel deposit. However, sheet-like bodies occur more commonly directly overlying major sandstone bodies, suggesting that moderately higher sinuosity overbank channels may occur along the abandoned courses of major rivers. Perhaps topography along abandoned major channel belts is associated with local areas of very low relief. Because many of the sheet-like minor channel bodies have dimensions similar to individual storeys within major channel belt deposits, some examples may reflect deposits associated with the avulsion of the major river channel (i.e. avulsion belt deposits of Smith et al., 1989).
employing the same methods used to interpret individual storeys in major sandstone bodies (see above; Fig. 18; Willis, 1989). Minor channel sandstone bodies with paleocurrent orientations generally perpendicular to the outcrop, and inclined bedsets filling to one side, record minor channel deposits exposed perpendicular to mean downstream direction. Apparent width to thickness ratios of such bodies (less than 50, i.e. ribbons), low proportion of lateral-accretion bedsets relative to channel-filling bedsets and limited variation in paleocurrent directions, suggest that these channels did not migrate laterally nor increase in sinuosity to any great extent before they became abandoned and filled. This suggests that most channels were relatively short-lived, or that channel banks were stabilized by vegetation or erosion-resistant overbank sediment. Bodies with inclined bedsets dipping in both directions away from the center of the body, have paleocurrent orientations generally parallel to the outcrop (cf. Fig. 20), and are interpreted as single-channel bar
Lobate and wedge-shaped bodies
Description Most deposits between major sandstone bodies occur in lobate or wedge-shaped bodies, several decimeters to several meters in thickness, defined by vertical variations in bedset thickness, mean grainsize and extent of sediment disruption (Figs. 7, 11, 13, 16). These bodies commonly appear to be horizontal beds in outcrop, but they gradually vary in thickness and internal lithofacies when traced along strike. These bodies continue along strike for only a few hundred meters to rarely a few kilometers. Some are dominated by sandstone where they are thickest, whereas many are dominated by mudstone throughout. In general, those dominated by sandstone are thicker and extend for greater distances along strike. Bodies contain decimeter to centimeter thick bedsets internally. Bedset inclinations and along-strike changes in bedset geometry could be recognized
64
B. WILLIS
only in the thickest and coarsest-grained examples (see below), Although correlation of individual lobate/ wedge-shaped bodies between adjacent measured logs is possible in many cases (Figs. 7, 11, 13, 16), individual bodies are not included on bedding diagrams (Figs. 6, 10, 12, 15). This is because the limits of bodies can become difficult to define where they thin and become highly disrupted, and correlations are inaccurate unless individual bodies were traced across the numerous gullies cutting exposures. Time limitations precluded such detailed tracing. Scale limitations allowed only generalized trends in bedset thickness and grainsize to be presented on sedimentary logs (Figs. 7, 11, 13, 16). A well exposed and unusually thick example of a lobate body was documented several kilometers to the west of the deposits presented in Fig. 6, at a similar stratigraphic level within the Chinji Formation (Figs. 22E, 24). This body is dominated by fine- to very fine-grained sandstone. It has a horizontal basal erosion surface and a convex-upward top that grades upward into overlying mudstone beds. The body is nearly 10 m thick (where thickest) and gradually thins to only a few meters thick in both directions over distances of about a kilometer. Bedsets are underlain by minor erosion surfaces, and fine upward. Nearly all bedsets are dominated by cross-laminated sandstone, although planar stratified sandstone or disrupted mudstone is common in their fining tops. Bur-
rows, root casts, and mudcracks disrupt all bedsets to some extent throughout the body. Individual bedsets thin, fine, contain less planar strata and become more disrupted in the direction the body thins. Where the body is thickest, bedsets dip at 1-2 ° in the direction the body thins (e.g. Fig. 22E). In places, bedsets vary slightly in dip, producing very low-angle truncations of adjacent bedsets. Bedset dips gradually decrease to horizontal toward the thinned edges of the body. Successive bedsets also tend to thin, fine, become dominated by cross-lamination, and become more disrupted vertically within the body. However, in locations where the body is thinner, successive bedsets coarsen and thicken upward in the lowest few decimeters (Fig. 24, log 2). Further along strike in the direction it thins, the body divides into multiple, thinner, fining-upward sets of bedsets that feather into adjacent mudstonedominated deposits (Fig. 24, log 1). Paleocurrents progressively vary by at least 30° along strike (Fig. 24). Wedge-shaped bodies display similar alongstrike and vertical variations in bedset geometry, grainsize, sedimentary structures, and disruption as lobate examples. However, wedge-shaped sandstone bodies thin away from an adjacent channel sandstone. Two types of wedge-shaped bodies are recognized: those with bedsets that pass conformably into the adjacent channel sandstone, and those that are truncated by the adjacent channel sandstone. Relatively thick and
-~70 meters~-~50 m e t e r s ~ 3 5 meters-4-~35 meters-~-
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Fig. 24, Sedimentological logs measured at several locations along a thick lobate sandstone body. Figure 22E presents a photo of this same sandstone body taken where the body begins to thin in the opposite direction (about 100 m to the right of the area covered by these logs).
ANCIENT RIVER SYSTEMS IN THE HIMALAYAN FOREDEEP
coarse-grained examples of both types are documented in Figs. 15 and 16. In the upper example (Fig. 15, adjacent to Storey A), the wedge-shaped body thins and fines to dominantly mudstone within several hundred meters along strike (Fig. 16, logs 1-6, meters 37-42). Bedsets within this wedge-shaped body are horizontal and appear conformable with those in the top of the channel sandstone body to the left, inclined at about 3° just to the right of the major sandstone body margin, and gradually decrease in dip to the right as the wedge-shaped body thins (Fig. 22F). Where the wedge-shaped body is thickest, bedsets are dominated by planar strata, but contain a few decimeter-thick sets of large-scale cross-strata lower and cross-lamination higher. Bedsets gradually fine, become more disrupted and are increasing dominated by crosslamination to the east (right) as the wedge-shaped body thins. Paleocurrent orientations within the wedge-shaped body are generally parallel to the outcrop and nearly perpendicular to those within the adjacent major sandstone body to the left, but they gradually shift to become more outcrop-perpendicular along strike in the direction the body thins. Small 1-3 m thick and several meters wide channel-forms cut into the top of this wedgeshaped body where it is thickest to the left (e.g. Fig. 16, log 3, meters 40-42). Bedsets within these small channel forms are slightly coarser than those within the underlying wedge-shaped body, and contain large-scale cross-strata lower down but dominantly planar strata and crosslamination higher up. Paleocurrent orientations within these small channel forms are generally parallel to those within the adjacent channel sandstone body and thus are at a high angle to those within the underlying wedge-shaped body. In the second example (adjacent to Storey C in Fig. 15; Fig. 16, log 8, meters 21-28) bedsets within the wedge-shaped body are truncated by the upper part of the channel-filling margin of the adjacent major sandstone body. In this second example bedsets also dip to the right at a very low angle away from the channel margin, and paleocurrents are oriented generally perpendicular to those of the adjacent major sandstone body. Bedsets are finer grained and thinner (1 to
65
a few cm thick) than those adjacent to the channel in the first example, and are dominated by cross-lamination throughout. Most lobate and wedge-shaped bodies are thinner (1 to a few meters thick) and finer grained than the examples just described, and thinning along strike can only be noted when tracing the body over several hundred meters (Figs. 7, 11, 13, 16). Some examples do not appear to have erosional bases, and the base of these bodies are defined only by an abrupt vertical change in bedset thickness, grainsize and the level of disruption. Many bodies fine upward, but some coarsen upward or initially coarsen and then fine upward (e.g. Fig. 5, logs 21-23, meters 39-45, and many others). Bedsets within these thinner bodies appear horizontal in the field and sandstone-dominated bedsets occur only where the body is thickest. Coarser bedsets can have erosional bases with locally developed mud-clast breccia, but finer bedsets are not clearly erosional. Internally, coarser bedsets are similar to those described above, whereas finer-grained bedsets commonly contain horizontally laminated mudstones. Some finer-grained bedsets also contain symmetrical cross-lamination. Individual bedsets commonly can be traced along strike for many tens of meters, but in places they also thin laterally. Shallow channel forms, decimeters deep and meters wide, occur rarely within these bodies. These shallow channel forms are bounded below by a minor erosion surface, overlain by a thin layer of intraformational breccia, that truncates several underlying beds. Such channel forms are normally filled by horizontal mudstone bedsets. The thickness of individual bedsets within lobate and wedge-shaped bodies normally decreases vertically and in the direction the body fines, but the degree of disruption always increases upward. Therefore, in fining-upward bodies, thicker and coarser bedsets lower down contain relatively undisrupted cross-laminated siltstone to very fine sandstone, but thinner and finer bedsets higher up have highly disrupted mudstone internally and are commonly cut by abundant mudcracks (e.g. Fig. 7, log 5, meters 41-47). Conversely, in coarsening-upward bodies,
66
B. WILLIS
thin and fine bedsets lower down display well preserved horizontal lamination and only moderate disruption, but thicker and coarser bedsets higher up are completely disrupted internally (e.g. Fig. 7, log 5, meters 4-9).
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Interpretation Meters-thick, lobate and wedge-shaped bodies are similar to modern levee or crevasse-splay deposits (Fisk, 1947; Kruit, 1955; Coleman, 1969; Singh, 1972; Ray, 1976; Farrell, 1987; Tye and Coleman, 1989; Smith et al., 1989). Lobate bodies record a cross-section through crevasse splays that were hundreds of meters to kilometers across, Wedge-shaped bodies, thinning away from an adjacent channel deposit, may record cross-sections through splay lobes or through levees extending along the adjacent channel. Where bedsets within the wedge appear conformable with those in the top of the adjacent channel body (e.g. Fig. 15, adjacent to Storey A) the s p l a y / l e v e e must have developed adjacent to the depositional channel mar'-in on the inside of a channel bend (Fig. 25a). Where the wedge body is truncated by the channel filling margin of the adjacent channel body (e.g. Fig. 15, adjacent to Storey C), the river channel must have continued to migrate laterally in the direction the splay/levee was prograding, eroding earlier formed channel proximal splay/ levee deposits (Fig. 25b). Along-strike thinning of bedsets, decrease of mean grainsize, increase in sediment disruption and change in sedimentary structures record a decrease in flow strength and deposition rate laterally in the direction the body thins. Crevasse channels were relatively rare as demonstrated by the scarcity of small channel-fills cutting into the top of lobate and wedge bodies. Thus much of the s p l a y / l e v e e deposition must have been associated with shallow sheet flows. Thicker, coarsergrained bodies record deposition by more rapid flows, presumably in crevasse channel-proximal locations. Horizontal basal erosion surfaces with scour structures record local erosion prior to deposition of the splay/levee. The high proportion of overbank deposits composed of these metersthick bodies (perhaps 70%) suggest that most of the sediments on the floodplain were deposited
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Fig. 25. Bedding relationships between channel bodies and adjacent wedge shaped bodies. (a) A crevasse splay deposit formed on the inside of the channel bend will overly (and
perhapstruncate) the bedset concordant storey margin of the underlying channel sandstone body. (b) A crevasse splay/levee deposit formed on the outside of the channel bend will be truncated by the channel filling margin of the adjacent chan-
nel sandstone body, and channel proximal deposits may be removed as the channel continues to migrate.
by progradation of crevasse splays and levees, and not by gradual vertical accretion in floodbasins. Bedsets within these bodies record individual flood events. Minor erosion surfaces underlying bedsets record slight erosion of the splay/levee surface during initial or high flood stages. Inclined bedsets, which dip down to meet the basal erosion surface of the body at up to 3 °, indicate gradual progradation of the s p l a y / l e v e e onto the floodplain over many flood events. Bedset slopes suggest that prograding splays/levees could produce a local floodplain topography of nearly 10 m and local relief of nearly 1 : 20; however, such slopes were quite unusual. More usual examples (see Figs. 7, 11, 13, 16) suggest a floodplain topography of only a few meters and local relief less than 1: 100. Decreased bedset dips from channel-proximal to distal locations reflect decreased slopes away from the channel. The dominance of cross-lamination within most bedsets suggests most of these splays/levees were cov-
67
ANCIENT RIVER SYSTEMS IN T H E HIMALAYAN FOREDEEP
ered by ripples during episodes of deposition, Planar strata within fine-grained bedset tops may record rapid shallow flows over upper stage plane beds or ripple migration under conditions of a very low deposition rate. Upward decrease in grainsize within bedsets records waning overbank flows and deposition of mud from suspended load. Bioturbation, mudcracks and root casts record post-flood disruption of sediment, Paleocurrents within wedge-shaped bodies suggest flows parallel to directions of bedset dip (contrast with bedsets in channel deposits). Paleocurrents within wedge-shaped bodies, which are initially perpendicular to those within the adjacent channel body but become more parallel to those within the channel where the wedge thins along strike, record overbank flow away from the channel where splay/levee slopes are steep, but gradual diversion of flow to down-basin directions further onto the floodplain where splay/ levee slopes decrease (e.g. Fig. 16, adjacent to Storey A; cf. Smith et al., 1989). Paleocurrents within channel-belt-proximal crevasse channels suggest that they tend to flow obliquely away from the major channel belt, preferentially feeding new areas in the down-basin direction (e.g. Fig. 15, wedge deposit adjacent to Storey A). Thinner, mudstone-dominated bodies record deposition by slower flows, presumably in crevasse channel distal locations. Nearly horizontal bedsets suggest low floodplain relief in such locations. Rare, shallow channel-forms and local mud-clast breccia at the base of bedsets record local scouring of the floodplain surface during flood events. However, the absence of erosion surfaces underlying some bodies, and underlying many bedsets within bodies, indicate that scouring by flood waters was less extensive than within channel proximal locations. Rare symmetrical cross-lamination within bedsets records locally ponded water and the formation of wave ripples, Fining upward bodies record a progressive decrease in deposition rate and flow velocity over successive flood events. Coarsening-upward bodies probably record gradual progradation of crevasse splays or levees into crevasse channel distal areas over many flood events. Bedsets lower in bodies generally have well preserved internal
stratification, suggesting relatively rapid burial or deposition in areas of ponded water. The progressive upwards increase in bioturbation/disruption observed in most bodies may record gradual decrease in deposition rate with time. However, because bedsets become more disrupted upward even in coarsening-upward bodies, sediment lobes were probably deposited rapidly and then were followed by a period of non-deposition during which more deeply buried bedsets were protected from disruption.
Laminated mudstone bodies Description Laminated mudstone bodies (composed dominantly of claystone) are several decimeters to meters thick, have non-erosional bases, and have horizontal tops (e.g. see relatively thick examples in Fig. 7, logs 5-27, meters 30-40; Fig. 16, logs 2-8, meters 0-10). The base of these bodies conform to the topography of underlying deposits (most commonly the top surface of a paleosol). Some bodies grade upwards into an overlying horizontal paleosol, but more commonly these bodies are truncated upwards by overlying channel deposits or lobate bodies. Laminated mudstone bodies can continue along strike for kilometers. A few examples thin along strike and gradually become more disrupted until they are indistinguishable from the surrounding deposits, However, most examples continue until they are truncated along strike by a channel body. Laminated mudstone bodies can not be divided into bedsets, but show changes upwards. Lower down in bodies, sediments are relatively undisrupted (but contain sparse vertical burrows and root casts), most sediments are finely laminated (but centimeter-thick beds can be massive and calcified), and isolated lenses of symmetrical cross-laminated siltstone can occur. Root casts and mudcracks become abundant upwards in the body. Initially thin centimeter-deep mudcracks are restricted to thin horizons separated by beds of laminated mudstone, but near the top of the body, wider decimeters-deep cracks filled with clasts of laminated claystone completely disrupt the sediments.
68
B. WILLIS
Rarely, light gray calcareous laminated mudstones also occur, most commonly above and adjacent to channel filling deposits (e.g. Fig. 5, logs 3-7, meters 39-41). These mudstones occur in decimeters-thick bodies that can be traced along strike for only a few hundreds of meters before they become completely disrupted. Within such bodies there is a vertical alternation between centimeter-thick layers containing fine laminae and layers that are highly disrupted by burrows, roots, and mudcracks. These layers can be horizontal but also can be contorted and recumbently folded along irregular underlying disrupted surfaces, or where cut by large mudcracks. Fine, equant sparry to micritic carbonate concentrations occur along the walls of some mudcracks, burrows and root casts,
Interpretation Bodies of horizontally laminated claystone record deposition within shallow ponds on the floodplain that were at least kilometers across (e.g. Tye and Coleman, 1989). The irregular nonerosional bases and horizontal tops of bodies, and the horizontal laminae within, suggest lake sediments gradually filled topographic low areas. Root casts and lenses of symmetrical cross-laminated siltstone lower down within bodies suggest that ponds were shallow enough for bottom sediments to support plants and become modified by water surface waves, respectively. Sparse vertical burrows imply that ponds filled rapidly or were biologically restricted. Thin beds that are massive and highly calcified resemble hardgrounds, and suggest periods of low suspended sediment input and low wave reworking. Upward increase in disruption by bioturbation and desiccation cracks records shallowing of the ponds and periods of subaerial exposure of sediments as ponds filled, Calcareous laminated mudstone beds record ponded areas where deposition rates where low. High carbonate content, fine laminae conformably covering irregular surfaces and recumbent folding of some beds suggest they may have been covered by cyano-bacterial mats. Local concentration of calcite along mudcracks, burrows and root casts record periods of subaerial exposure and sediment desiccation,
Paleosols Nearly all mudstone beds show evidence of disruption by roots, and thus must have been disrupted to some degree by pedogenesis. However, many meters-thick layers in these strata are distinct from superjacent deposits due to their heightened red hues, intense disruption, and complex vertical variations of structure and fabric (Figs. 7, 11, 13, 16). Such disrupted layers cornprise nearly a third of mudstone-dominated deposits within this stratigraphic succession and are widely recognized to be distinct paleosols (Johnson et al., 1982, 1985; Retallack 1985, 1986, 1992; Behrensmeyer 1987, 1989; and others). Paleosols are not sediment bodies in the same sense as others described here because they are not defined by variations in primary sedimentary structures and textures (cf. sediment bodies described above). Rather, they are mappable layers of strata defined by distinct patterns of sediment disruption superimposed on underlying sediment bodies. Paleosols continue across entire documented intervals, except where they are truncated by overlying sediment bodies. Many paleosols stay nearly parallel along strike for kilometers (Figs. 6, 10, 12, 15). However, in some cases, paleosols gradually converge along strike, and some join laterally to form composite paleosols. In most cases, paleosols can be divided into two distinct horizons: an upper horizon that lacks disseminated groundmass carbonate, and a lower horizon that contains abundant carbonate concentrations. The thickness of these horizons and the abundance of various pedogenic structures within paleosols can vary greatly, both between different paleosols and along strike within individual paleosols (Figs. 7, 11, 13, 16). Detailed studies of along-strike variations in paleosol profiles and micromorphology in relation to variations in lithofacies and floodplain physiography are in progress (Behrensmeyer et al., 1992; Behrensmeyer and Willis, 1992; Willis and Behrensmeyer, 1992a, b; Behrensmeyer and Willis, unpubl, data), thus only a very generalized description of paleosols will be presented here. Higher within a paleosol, deposits are generally characterized by: (1) bright red-brown to
69
A N C I E N T R IVE R S YS T E MS IN T H E H I M A L A Y A N F O R E D E E P
strong-brown hues; (2) groundmass that lacks elfervescence in dilute hydrochloric acid (interpreted to reflect the leaching of carbonate from the upper part of the soil); (3) gray mottles (cornmonly reflecting local reduction along root casts and burrows); (3) coarse-grained concentrations (interpreted to reflect eluviation of fines from the upper part of the soil); (4) slickensided clay films (interpreted to record clay particles aligned along surfaces faulted due to clay swelling); (5) small carbonate or sesquioxide concentrations. Carbonate concentrations can occur along cracks but more commonly occur as micritic or fine crystalline calcite nodules, normally clustered along root casts, burrows and mudcracks. Sesquioxide concentrations occur as millimeter-diameter concretions and coatings on carbonate nodules. Ped structure is observed only rarely within these paleosols, and the structure of most paleosols is dominated by pervasive burrows, Lower in paleosols, abundant carbonate concentrations typically occur. Carbonate concentrations normally occur as large, composite nodules (several centimeters across), but also occur as smaller isolated nodules aligned along mudcracks, burrows and root casts. In some layers, composite nodules are elongate vertically along large mudcracks (reflecting preferential carbonate precipitation along preexisting mudcracks; cf. Blodgett, 1988; Joeckel, 1991). Sesquioxide nodules can be broadly disseminated within the paleosol, but also can occur densely packed directly above intervals dominated by carbonate nodules, Rarely argillic concentrations are observed along burrows and root casts, particularly in the top parts of layers containing abundant carbonate concentrations. Complex textural relationships within carbonate concentrations, and superposition of carbonate and sesquioxidic concentrations, suggest complicated histories of accretion and changing pedogenic conditions over time. Siwalik paleosols display vertical variations in structure and fabric similar to that observed within modern alluvial soils of the Indo-Gangetic basin (Sehgal et al., 1968; Sehgal and Stoops, 1972; Ahmad et al., 1977; Brammer and Brinkman, 1977; Brinkman, 1977; Sidhu and Gilkes, 1977; Sidhu et al., 1977). Estimates of time sepa-
rating paleosols based on paleomagnetic dating (order of 104 years; see Behrensmeyer, 1987; Willis, 1993b) are comparable to estimates for development of well differentiated soils formed within the modern Indus Basin (Ahmad et al., 1977). Sesquioxide concentrations along carbonate nodules and within small concretions appear most commonly in modern soils where the soil profile becomes seasonally waterlogged then desiccated causing repetitive cycles of reduction and reoxidation (Sehgal and Stoops, 1972; Brinkman, 1977; Sobecki and Wilding, 1983; Wright et al., 1991). Paleosols of the Chinji Formation contain sesquioxidic concentrations more commonly then paleosols of the Nagri Formation (80% versus 35% of paleosols, respectively), suggesting that seasonal waterlogging of floodplain surfaces was more common during deposition of the Chinji Formation. Siwalik Group paleosols have been interpreted as oxisols to alfisols underlying tropical, thickly wooded to savanna vegetation because of their red coloration and the presence of sesquioxide nodules (Johnson et al., 1981; Retallack, 1985, 1986, 1992). These interpretations also reflect fossil flora evidence extrapolated from the Siwaliks of India, and fossil faunal evidence from the Siwaliks of Pakistan. Paleosol-bounded sequences
Strata between well defined paleosols are composed of multiple lobate/wedge-shaped bodies, minor channel-form bodies, and bodies of |aminated claystone (Figs. 7, 11, 13, 16). Whereas lobate/wedge-shaped bodies and minor channelform bodies extend along strike for only a few tens of meters to a few kilometers (as described above), most paleosols extend for at least tens of kilometers along strike. Although some strata between well developed paleosols are intensively disrupted by roots, burrows and mudcracks, most contain primary stratification and are comparatively undisrupted relative to paleosols. This alternation between completely disrupted paleosols separating strata with preserved primary bedding defines larger-scale (1-10 m thick) sequences within these mudstone-dominated deposits. Understanding the geometry of paleosol-bounded
70 sequences relative to lithologic variations within, and relative to terminal margins of major channel belt deposits, is fundamental to interpretations of overbank deposition, Paleosols record ancient floodplain surfaces that received minimal deposition for prolonged periods of time (see above). Conversely, stratified sediments between paleosols record periods when overbank deposition rates were generally high and long hiatuses in deposition did not occur, Paleosol-bounded sequences thus reflect an alternation between periods when overbank deposition rate was generally high and continuous, and periods when overbank deposition rates were minimal. Because such sequences contain many bedsets, they can not be related to individual deposition events. Further, because sequences contain multiple minor channel-form and lobate/ wedge-shaped bodies, they can not be related to progradation of individual splays or levees onto the floodplain. Individual crevasse splays can be deposited within just several tens of flood events, and complex overbank sequences including overlapping crevasse lobes and channel deposits can form in on the order of a hundred years (e.g. Smith et al., 1989; Tye and Coleman, 1989). Similarly high deposition rates within these deposits are supported by relatively undisrupted sediments lower within paleosol-bounded sequences, Paleosol-bounded sequences within the Chinji and Nagri Formations represent on the order of 104 years on average (Willis, 1993b). Evidence for rapid deposition rates of sediments between paleosols suggests that most of the time recorded within these strata is associated with development of paleosols. Rapid deposition of overbank sequences have been reported in modern river systems associated with: (1) build-up of multiple, overlapping crevasse splays directly adjacent to the major channel belt (e.g. Fisk, 1947; Coleman, 1969; Farrell, 1987; see also Bown and Kraus, 1987; Smith, 1990); (2) progradation of sediment into low areas on the floodplain further away from the major channel belt associated with tributary, distributary or crevasse channels (e.g. Parkash et al., 1983; Wells and Dorr, 1987; Tye and Coleman, 1989); (3) rapid deposition of overbank deposits
B.WILLIS following river avulsion but directly preceding development of a new channel belt (e.g. Smith et al., 1989). Each case involves the deposition of lobate overbank deposits as flows pass to areas of lower gradient and decelerate. In the first case, overbank deposition would largely be associated with the build up of an alluvial ridge and deposition of coarse-grained sediments close to the major channel belt. In the latter two cases, overbank deposition would be associated with the filling of low areas on the floodplain, and thus the reduction of floodplain relief. In the latter cases, grainsize and thickness variations within sequences would not necessarily reflect the proximity of the major channel belt. In these deposits, most paleosol-bounded sequences vary in thickness in a complex way along strike (Figs. 6, 10, 12, 16). Some contain minor channel and coarser-grained crevasse-splay deposits where they are thicker, and they thin and fine along strike (e.g. Fig. 7, logs 6-23, meters y = 45 to 50; Fig. 11, logs 11-19, meters 3-9; amongst others). In most cases, however, the thickness of paleosol-bounded sequences is not related to internal lithologic variations. For exampie, the paleosol-bounded sequence to the left of Storey V (Fig. 6; Fig. 7, logs 1-12, meters 5-10) does not vary in thickness in a systematic way with distance from the minor channel body, and the sequence adjacent to Storey M (Fig. 10; Fig. 11, logs 13-24, meters 10-15) dramatically fines to the right but changes little in thickness. Laminated mudstone beds occur most commonly lower within finer-grained parts of sequences (e.g. Fig. 7, meters 30-40; Fig. 11, logs 18-24, meters 1015). In places, paleosols converge defining the margins of paleosol-bounded sequences (e.g. Fig. 6, between meters 0-10; Fig. 11, logs 15-19, meters 5-10). In most cases, however, paleosolbounded sequences continue across entire documented intervals despite local variations in thickness and internal lithology. Strata directly overlying the major sandstone body documented in Figs. 6 and 7 show variations along a paleosol-bounded sequence. This sequence is discussed in more detail below to emphasize relationships between sediment bodies within paleosol-bounded sequences. This se-
A N C I E N T R I V E R SYSTEMS IN T H E HIMALAYAN F O R E D E E P
quence is defined by two paleosol horizons that extend across the entire area studied: a lower one capping the major sandstone body (Fig. 6, about meters y = 30 to 35), and an upper one about 10 m higher (Fig. 6, meters y = 40 to 45). The lower paleosol records soil formation on an elevated channel belt following its abandonment by avulsion. Laminated mudstones across the base of this sequence suggest that this area later became a floodplain low that was initially ponded. Thickening of the sequence to the left can be related to filling of the depositional topography on the submerged channel belt deposit. This low ponded area was then locally invaded and filled by multipie levees or crevasse splays associated with minor channels. Multiple splay/levee deposits adjacent to Storey C (Fig. 6), and thinning of these deposits to the left, record decreased deposition rates away from the minor channel deposit (Fig. 7, logs 25-32, meters 35-45). However, because these splay/levee deposits are truncated to the right by Storey C, it is not possible to distinguish whether the channel eroded into a previously deposited splay lobe, or whether the splays/ levees formed away from the channel and lower beds were truncated as the channel enlarged or migrated. Storey B (to the left) truncates mainly laminated mudstones, thus clearly not all minor channels are associated with splay deposits (see also numerous other examples in Figs. 7, 11, 13, 16). Calcareous mudstones capping this channel sandstone and adjacent laminated mudstones to the right (Fig. 5, logs 3-8, about meter 40), suggest lower deposition rates occurred after this channel (B) was filled. Storey A also truncates mainly laminated mudstones; however, a metersthick levee deposit is associated with its top. Following deposition of Storey A, deposition rates of this sequence slowed and the deposits became capped by the upper paleosol. The deposits within this sequence suggest a rather complex set of depositional environments: including areas of ponded water that received only suspended muds, the development of overlapping crevasse splay lobes formed where flows decreased in velocity into ponded areas, and minor channels that locally had adjacent levees. Similarly complex histories of deposition are suggested by sediment vari-
71
ations across other paleosol-bounded sequences within the Chinji and Nagri Formations. Termination of storeys in major sandstone bodies against overbank deposits show how overbank deposits vary away from the edges of major channel belts. Most major channel margins truncate several paleosol-bounded sequences; however, only the topmost truncated sequence can be related to overbank deposition coeval with the existence of, and active deposition within, the adjacent channel belt. In some cases sequences adjacent to the top of a major channel margin are coarser-grained levee/splay deposits (e.g. Fig. 16, meters 40-45); however, such cases are rare and such splay/levee deposits normally thin and fine to mudstone along strike over less than a kilometer. More commonly, topmost sequences truncated by a major channel margin are n o t unusually thick nor coarse-grained, and they do not vary systematically away from the channel deposit (e.g. Fig. 11, logs 5-25, meters 40-55; logs 21-25, meters 15-20; and many others). In fact, sequences truncated by major channel margins can be quite fine grained (e.g. sequence dominated by claystone directly adjacent to the terminating margins of Storeys O and P in Fig. 11). The thickness of paleosol-bounded sequences adjacent to channel belt margins suggests that only a few meters of sediment was deposited on the floodplain adjacent to channel belts before they became abandoned, and thus channel belts were associated with relatively subdued alluvial ridges (less than a few meters in elevation). The absence of thick, coarse-grained overbank sequences adjacent to most major channel belt margins, and pronounced variations in the thickness and grainsize of both splay/levee deposits and paleosol-bounded sequences over distances of only a few hundred meters, do not support overbank deposition models that relate the thickness and grainsize of meter-scale overbank sequences to the proximity of coeval major river channels (cf. Johnson et al., 1981; Bridge, 1984b; Kraus, 1 9 8 7 a ;Smith, 1990). There is also little evidence that paleosols vary systematically away from channel-belt margins (cf. Bown and Kraus, 1987; Smith, 1990). Paleosol-bounded sequences clearly record a
72
complex distribution of environments including local levees, crevasse splays and ponded areas that do not vary in a simple way across the floodplain. Finely laminated mudstones lower within many paleosol-bounded sequences suggest that high overbank deposition rates were associated with the filling of low ponded areas on the floodplain. Abundance of paleosols with iron nodules (particularly in the Chinji Formation) suggest that local ponding of water in low areas on the floodplain was common for at least part of the year. It is not possible to tell whether there was continuous filling of different localized low areas on the floodplain (e.g. Tye and Coleman, 1989), versus wide spread (but short-lived) periods of rapid deposition in low areas associated with avulsion of major channel belts (e.g. Smith et al., 1989). Relatively thick sequences, that contain minor channel-form bodies of similar thickhess to individual storeys in major sandstone bodies, are the best candidates for avulsion-related deposits (e.g. Fig. 7, meters 30-40). Conclusions (1) The Miocene Chinji and Nagri Formations of northern Pakistan record fluvial deposition adjacent to the rising Himalayan Mountains. Rivers were predominantly braided with low channel sinuosities (1-1.3). Individual channel segments of the Chinji Formation were generally 100-200 m wide, 5-10 m deep and carried discharges of about 400-800 m3/s. Full channel-belt widths were less than 2 km, braiding index was about 2, and full river discharges were on the order of 103 m3/s. Deposits of a larger river are recorded within the Nagri Formation, where individual channel segments were 200-400 m wide, 15-30 m deep and carried discharges of about 3000-5000 m3/s. Full channel widths and braiding index of these larger rivers are difficult to estimate; however, there was clearly more than one coeval Channel and thus full channel discharges appear to have been on the order of 10,000 m3/s. (2) Channel bars were characterized by low angle (2-5 °) along-stream dipping surfaces and slightly steeper (5-11 °) cross-stream dipping sur-
B. WILLIS
faces. Bars had along-stream lengths of 0.5 to 1.5 km within Chinji Formation rivers, but 1.5-3 km within the larger Nagri Formation rivers. Most bars migrated downstream as they grew and increased in sinuosity, but not all. Channel switching was clearly an important process within channel belts as indicated by the downstream stacking of storeys and the dominance of coarse-grained channel fills. In a few locations, unambiguous evidence for braid bars is recorded by diverging paleocurrents around concave-upward bedding surfaces. (3) Floodplains had low relief, that was generally less than meters per kilometer, but higher local relief was associated with abandoned channel belts, small overbank channels, crevasse splays and levees. Most floodplain surfaces were covered by well developed soils. Overbank deposits are characterized by meters-thick stratified sequences separated by paleosol horizons. These sequences are dominated by overlapping crevasse splay and levee deposits. Finely laminated mudstones were deposited in ponded areas away from active splays/levees. The thicknesses and grainsize of these overbank sequences are not typically related to the proximity of coeval major river channels, in contrast to recently published overbank deposition models proposed for other basins (e.g. Bridge, 1984b; Kraus, 1987a; Smith, 1990). Instead most overbank deposition appears to be related to localized crevasse channels that carried sediment to low areas on the floodplain, or to rapid periods of floodplain deposition associated with channel belt avulsion. Acknowledgements This paper is part of the authors Ph.D. dissertation supervised by John S. Bridge at the State University of New York at Binghamton; it was greatly improved by his instructive comments throughout the research. I am grateful to A.K. Behrensmeyer of the Smithsonian Institution for introducing me to the Siwalik exposures. This research was supported by grants to J.S. Bridge and A.K. Behrensmeyer from N.S.F. and the Smithsonian Institution, and a grant to J.S. Bridge from the Donors of the Petroleum Research Fund
73
ANCIENT RIVER SYSTEMS IN THE HIMALAYAN FOREDEEP
(administered by the American Chemical Society).
J.S. Bridge, A.K. Behrensmeyer, S. Gabel, B. Demicco, T. Lowenstein, and J. Kradyna are thanked for their comments on the initial manuscript. A. Miall, P. Friend, D. Burbank, and an
anonymous
reviewer provided helpful
corn-
ments via reviews for Sedimentary Geology. This research benefitted from discussions in the field with J. Barry, A.K. Behrensmeyer, P. Friend, M. Raza, and M. Zaleha (amongst others). This research could not have been completed without the logistical support of the Geological Survey of Pakistan (Stratigraphy and Paleontology Branch, Islamabad), with specific thanks to officers M. Raza, K. Sheikh, I. Khan, and B. Jimal for their generous efforts that allowed me to stay in the field for extended periods. References Abroad, M., Ryan, J. and Paeth, R.C., 1977. Soil Development as a function of time in the Punjab River plains of Pakistan. Soil Sci. Soc. Am. J., 41: 1162-1165. Allen, J.R.L., 1983. Sedimentary Structures: their Character and Physical Basis. Elsevier, New York, N.Y., 1160 pp. Badgley, C. and Behrensmeyer, A.K., 1982. Paleoecology of Middle Siwalik sediments and faunas, northern Pakistan. Palaeogeogr., Palaeoclimatol., Palaeoecol., 30: 133-155. Barndt, J., Johnson, N.M., Johnson, G.D., Opdyke, N.D., Lindsay, E.H., Pilbeam, D.R. and Tahirkheli, R.A.K., 1978. The magnetic polarity stratigraphy and age of the Siwalik Group near Dhok Pathan village, Potwar Plateau, Pakistan. Earth Planet. Sci. Lett., 44: 355-364. Barry, J.C., Lindsay, E.H. and Jacobs, L.L., 1982. A biostratigraphic zonation of the Middle and Upper Siwaliks of the Potwar Plateau of northern Pakistan. Palaeogeogr., Palaeoclimatol., Palaeoecol., 37: 95-130. Barry, J.C., Johnson, N.M., Raza, S.M. and Jacobs, L.L., 1985. Neogene Mammalian faunal change in southern Asia: correlations with climatic, tectonic and eustatic events, Geology, 13: 637-640. Beaumont, C., 1981. Foreland basins. Geophys. J.R. Astron. Soc., 65: 291-329. Beck, R.A. and Burbank, D.W., 1987. Basin-scale isochronous paleocurrent study of Chron 9 interval, Middle Siwaliks, Potwar Plateau, northern Pakistan. Soc. Econ. Paleontol. Mineral., Annu. Midyear Mtg., Abstr., 4: 6. Beck, R.A. and Burbank, D.W., 1989. Regional fluvial facies variability within the C5N magnetic polarity interval, Siwalik Group, northwest Himalayan foredeep, Pakistan. 4th Int. Conf. Fluvial Sedimentology. International Association of Sedimentologists, Barcelona-Sitges, Progr. Abstr., p. 70.
Behrensmeyer, A.K., 1987. Miocene fluvial facies and vertebrate taphonomy in northern Pakistan. In: F.G. Ethridge, R.M. Flores and M.D. Harvey (Editors), Recent Developments in Fluvial Sedimentology. Soc. Econ. Paleontol. Mineral., Spec. Publ., 39: 169-176.
Behrensmeyer,A.K., 1989. Overbank facies and paleosols of the Chinji Formation, Northern Pakistan. 4th Int. Conf. Fluvial Sedimentology. International Association of Sedimentolgists, Barcelona-Sitges, Progr. Abstr., p. 77. Behrensmeyer, A.K. and Tauxe, L., 1982. Isochronous fluvial systems in Miocene deposits of Northern Pakistan. Sedi-
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Am., Abstr. Progr., 24: 229. Behrensmeyer, A.K., Quade, J. and Willis, B.J., 1992. Floodplains and paleosols in the siwalik Miocene sequence:
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