Interpretation of overbank–avulsion cyclicity
The heterolithic deposits with only weak soils are interpreted as avulsion deposits, which is consistent with previous interpretations (e.g. Kraus & Wells, 1999). The lateral extent of up to 1 km at the Deer Creek Amphitheater section (DCA) (up to 5 km in the Elk Creek area; Kraus, 1987) and the metre-scale thickness of the heterolithic intervals are characteristic of avulsion deposits rather than simple crevasse splay deposits. For example, along the Brazos River in Texas, Taha & Anderson (2008) found avulsion deposits that cover areas ca 12 km2, which is 25 to 50 times more extensive than non-avulsive splay deposits. Various authors have reported thicknesses of 3 to 4 m for avulsion deposits (Morozova & Smith, 2000; Aslan et al., 2005; Taha & Anderson, 2008).
The fine-grained deposits with more strongly developed palaeosols are interpreted as overbank deposits whose slow accumulation rates allowed the development of one or more (cumulative) palaeosol profiles. The vertical transition from overbank to avulsion deposits is mostly gradual, in line with the progradational style of avulsion in which a sediment wedge extends from the channel and grows in a downflow direction (Slingerland & Smith, 2004).
The metre-thick cycles (equal to simple pedofacies cycles of Kraus, 1987) are distinctive because their heterolithic components (avulsion deposits) are thick and laterally extensive (Figs 2 and 3). In contrast, the stratigraphic intervals that separate the individual palaeosols within the metre-scale cycles are thin and indistinct, and typically show evidence for some pedogenesis (Fig. 4). The present authors speculate that many of the thin intervals are also the result of channel avulsion; however, the difference between the thin intervals and the 3 to 4 m thick, laterally extensive heterolithic intervals is related to the scale of the avulsion. Various authors (e.g. Mackey & Bridge, 1995; Heller & Paola, 1996; Jerolmack & Paola, 2007) have distinguished between local avulsions, in which the new channel rejoins the channel from which it split a short distance downflow, and regional avulsions, which create an entirely new channel in a new floodplain location.
The thinner, less distinct stratigraphic intervals between individual palaeosols probably reflect local avulsion or a simple crevasse splay (Fig. 7; overbank phase). The resulting avulsion belt or splay deposit was small, and the coarser deposits, especially sands, did not extend to the distal floodplain where the more strongly developed soils were forming. Nonetheless, a rapid incursion of sediment onto the floodplain buried and interrupted development of floodplain soils, and that newest sediment became the lower part of the next soil. The local nature of this type of avulsion can explain the relative insignificance of the heterolithic interval it produces and the quick overprinting by pedogenesis. The fact that each pedofacies cycle generally contains several palaeosols is attributed to the shifting locus of avulsion over time. Because avulsion reduces local super-elevation, succeeding local avulsions can move up-channel, although it is observed that the locus of avulsion eventually reverts to a down-channel position and then again moves up-channel (e.g. Mackey & Bridge, 1995; Stouthamer & Berendsen, 2007). In the absence of robust evidence for a connection between local avulsion and climate, and because avulsion is the key autogenic process in avulsion-dominated fluvial systems, the present authors consider the local avulsions in the Bighorn Basin to have been autogenic in nature.
Figure 7. Schematic diagram showing the sedimentary response to a precession cycle in the Willwood Formation. The stratigraphic column shows an interval of avulsion deposits (light coloured) between two intervals of multiple, strongly developed palaeosols (dark coloured). The palaeosols represent the overbank phase of the precession cycle; the avulsion deposits represent the avulsion phase. The vertical scale of the stratigraphic column is thickness. A single precession cycle represents ca 21 kyr; however, the avulsion phase was relatively rapid with respect to the overbank phase.
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The thick, laterally extensive heterolithic intervals, on the other hand, are attributed to regional avulsion (Fig. 7; avulsion phase). The change from local to regional avulsion may reflect a period in time when, at larger scales in the basin, the channel belts became super-elevated relative to the floodplain, making them susceptible to avulsion. After this phase of channel reorganization, during which gradients between channel belts and floodplains were lowered, a period of relative channel belt stability started, and the floodplain was characterized by a dominance of overbank deposition and development of mature palaeosols (Fig. 7; overbank phase). The data used here were derived from a vertical one-dimensional section supplemented by field tracing of the cyclicity laterally over approximately 1 km. To quantify the regional significance of the overbank–avulsion cyclicity, a three-dimensional stratigraphic analysis is required.
The time-series analysis and bandpass filtering is performed here on depth (thickness) series rather than time series. Therefore, sedimentation rate changes at a scale larger than the overbank–avulsion cyclicity (but shorter than million-year time scale) are hard to detect without additional age constraints. The rate of sedimentation changes within individual overbank–avulsion cycles. Individual palaeosols represent short-term, slow sedimentation and the weak pedogenic modification of the avulsion package indicates a short episode of fairly rapid sedimentation. Spectral power at wavelengths shorter than the overbank–avulsion cyclicity of 4·5 to 8·5 m should, thus, be interpreted with caution because power will be derived from sediment intervals that, while short, may represent various durations in a time domain. This study focuses on the potential allogenic forcing of overbank–avulsion cyclicity, which has to be resolved before smaller time scales can be discussed.
Orbital climate forcing of river avulsion
The magnetochron C24r to C24n.3n boundary was found in the Upper Deer Creek section (Abels et al., 2012) which is located 3·2 km to the south. Stratigraphic correlations in the field place the top of the DCA section at the very base of the Upper Deer Creek section. The DCA section, thus, occurs within the upper part of magnetochron 24r. Chron 24r has an astronomically constrained duration of 3·118 ± 0·158 Myr (Westerhold et al., 2007) and is 1219 m thick in the part of the Bighorn Basin studied here (Clyde et al., 1994), giving a sedimentation rate of 391 ± 21 m/Myr; this rate indicates that a sedimentary cycle thickness of 7·1 m represents approximately 18·2 ± 1·0 kyr, assuming constant sedimentation rates (Table 1).
To reduce the stratigraphic interval over which average accumulation rates are calculated, a second sediment accumulation rate is derived from the estimated 0·8 Myr ± 0·04 duration of mammalian biochron Wa4 (Gingerich, 2010). Uncertainty here is calculated from the 5% uncertainty on the duration of C24r in Westerhold et al., 2007. This zone is approximately 225 to 235 m thick in the DCA study area, yielding an accumulation rate of 288 ± 20 m/Myr. At this rate, a sedimentary cycle thickness of 7·1 m represents 24·7 ± 1·9 kyr (Table 1).
A third sedimentation accumulation rate is derived from the carbon isotope excursions of the ETM2 and H2 hyperthermals, which are recorded in the Upper Deer Creek section (Abels et al., 2012). The peak excursions of both events are separated by ca 30 m stratigraphically and separated by four to five precession cycles, which means ca 84 to 105 kyr as estimated using a marine age model (Stap et al., 2009). The resultant sedimentation rate in the Upper Deer Creek section is 316 ± 41 m/Myr which, when extrapolated to the DCA section, gives a duration of approximately 22·5 kyr ± 3·4 for the 7·1 m cyclicity (Table 1).
The thickness and duration of the overbank–avulsion cycles in the Deer Creek Amphitheater section as recognized in colour records are similar to those found in the Polecat Bench and Red Butte sections (Abdul Aziz et al., 2008), and the Upper Deer Creek section (Abels et al., 2012). In the nearby, but older, Polecat Bench section that encompasses the Palaeocene–Eocene thermal maximum (PETM), cycle thickness is 7·6 m, resulting in a similar period of 19·4 kyr. In the Red Butte section, which is located to the south-east in the basin and is slightly younger, cyclicity was observed to be 8·7 m thick with an estimated period of 22·3 kyr (Abdul Aziz et al., 2008). In the only slightly younger, nearby Upper Deer Creek section, overbank/avulsion cycles are 7·0 m thick with a duration of 23·3 kyr per cycle. These estimates are all within the range of the duration of the orbital and climatic precession cycle with components between ca 18·8 kyr and 22·6 kyr in the early Palaeogene (Berger et al., 1992).
The amplitude of the precession cycle is modulated by the eccentricity of Earth's orbit (Berger et al., 1992) and alternations of differently expressed precession cycles are, thus, expected (Abels et al., 2009). Bundles of three to four clear precession cycles should be interrupted by one or two less developed precession cycles. The combination of lithology and filtered colour records indeed show more distinct cycles occurring in three groups of three or four cycles in the section (C–E, G–J and M–O; Fig. 5) separated by less distinct cycles (A–B, F, K–L and P). If precession is causing the overbank–avulsion cyclicity, this bundling matches the 100 kyr eccentricity cyclicity. The bundles of distinct cycles would relate to eccentricity maxima when precession amplitudes are largest and the climatic differences forcing sedimentation are strongest. The three observed 100 kyr eccentricity bundles (of three or four distinct precession cycles and one or two less distinct precession cycles) in the DCA section then mark a 405 kyr eccentricity maximum. Like the 100 kyr cycles, at the 405 kyr scale, the eccentricity maxima can be linked to well-developed simple pedofacies cycles with clear intervals showing mature palaeosol profiles.
On top of the section, a thick sheet sandstone is present, while, just below and above this sheet sandstone, the simple pedofacies cycles that are present are less developed. Also, just below the DCA section, the stratigraphy is characterized by a longer interval with poorly developed pedofacies cycles and the absence of mature palaeosols (Fig. 2). In line with this, a gradual increase of 0·5 points in the Soil Development Index occurs from the base of the section to the middle part, after which a gradual decrease of 0·5 points occurs towards the top of the section. These observations are consistent with deposition of the studied section during a 405 kyr eccentricity maximum.
Allogenic mechanism forcing overbank–avulsion cyclicity
Channel avulsion is of considerable interest as a control on stratigraphic patterns in alluvial deposits (e.g. Mackey & Bridge, 1995; Heller & Paola, 1996; Stouthamer & Berendsen, 2007; Hajek et al., 2010). Recent work, for example, has focused on channel sandstone clusters in both laboratory and field settings, and the fact that such clusters can be the result of autogenically controlled avulsion alone (Hajek et al., 2010, 2012). Other workers have ascribed avulsion-generated alluvial stratigraphies to a mix of autogenic and allogenic factors depending on the particular situation (e.g. Stouthamer & Berendsen, 2007; Phillips, 2011). Channel avulsion is principally related to super-elevation of the channel above its floodplain because of differential deposition (e.g. Bryant et al., 1995; Mohrig et al., 2000; Jerolmack & Paola, 2007). Levées directly adjacent to the channel aggrade more rapidly than the more distal floodplain and cause super-elevation, which renders the channel susceptible to avulsion. Eventually, some form of trigger initiates crevassing, and sands and muds are deposited rapidly on the floodplain above the interval with more intense soil development. Triggers include phenomena, such as ice jams or unusual flood discharges, which are spatially or temporally sporadic (Jones & Schumm, 1999).
Although the distinct cyclicity in the Willwood stratigraphy could simply be the result of autogenically controlled regional avulsion, the fact that the cyclicity shows distinct Milankovitch periodicities indicates that avulsion has been influenced by orbital climate change. For example, Nádor et al. (2003) noted that, in dominantly autogenic depositional systems, such as fluvial systems, changes in climate can influence the recurrence time of autocyclicity. Thus, if Milankovitch periodicity is recognized in a dominantly autogenically influenced lithological record, it can be attributed to allogenic modulation of the autocyclicity. In the case of the Willwood Formation, the rate of channel belt build up, and so the time required for the channel to become super-elevated above its floodplain, depended on sediment and water input, but ultimately on the creation of accommodation space by subsidence of the basin. Local accommodation space was generated during the overbank phase of the precessional cycle through a combination of continued subsidence and slower aggradation on the floodplain than on the channel belt (Fig. 7). That space appears to have been created at a rate that was paced by astronomical climate change, such that avulsion was triggered by climate change at the time super-elevation was reached. In this scenario, one would expect avulsion to be instigated during the orbital extremes or somewhere in the transitional stages towards the orbital extreme (Mann & Meltzer, 2007).
Jones & Schumm (1999) suggested that increases in flood magnitude and frequency can cause more rapid floodplain aggradation and, thus, increase the susceptibility of a channel to avulsion and also enhance the probability of a triggering event. Similarly, Sinha et al. (2005) concluded that avulsion of the Baghmati River in India appears to have been instigated by major floods. Consequently, the present authors hypothesize that regional avulsions were triggered during periods dominated by high magnitude floods, whereas more typical floodplain deposition took place during periods of lesser floods (Fig. 7).
Orbital climate change causing avulsion–overbank cyclicity
In the previous paragraphs, a mechanistic pathway of allogenic climate change was synthesized causing Willwood overbank–avulsion cyclicity. The regular cyclicity in four distinct stratigraphic sections and the match in time scales make a link with the climatic precession cycle tentative. However, the impact of the precession cycle on the Eocene Bighorn Basin is not well-known. In the following, the orbital cycle to climate pathway is discussed.
Orbital cyclicity of Earth's axis and orbit produces significant insolation differences (Berger et al., 1992). Precession extremes can generate insolation differences up to 20%, depending on latitude and season. The resulting climate impact on Earth's surface, or in the depositional environment, depends on local climate and, thus, land–sea configuration has a great influence. No temperature or precipitation data were available covering a complete precession cycle from the early Eocene Bighorn Basin. In the studied section, palaeosol proxies are hard to apply across a precession cycle because one part of the cycle is characterized by sandy mudrocks showing no or very incipient soil formation.
Lawrence et al. (2003) modelled realistic precession extremes with Eocene-like boundary conditions. These authors found that mean annual temperature and precipitation in the western interior of North America were not significantly affected by the precession cycle, but their model results show a summer temperature increase of 5°C, a ca 50% decrease in autumn precipitation, and ca 40% lower annual net moisture (precipitation – evaporation) availability in this region during precession minima relative to precession maxima.
Other studies show that precession cycles do affect the strength of monsoonal systems through their impact on summer insolation. For example, Tuenter et al. (2007) found that, in the Asian and African monsoon systems, the summer monsoonal system strengthens during precession minima because summer insolation is increased. If the modern North American monsoon becomes stronger, summer thunderstorms are more frequent and generate larger floods (e.g. Mann & Meltzer, 2007), ideal for triggering river avulsion. In contrast, if the system weakens, the magnitude and frequency of summer floods decline. Increased floods are related to warmer periods, such as the last precession minimum (Mann & Meltzer, 2007).
The precession modelling of Lawrence et al. (2003) and monsoonal studies (Tuenter et al., 2007; Mann & Meltzer, 2007) produce results that are contradictory, and a direct phase relation between overbank–avulsion cyclicity and orbital parameters cannot be readily deduced. For now, the higher seasonality leading to extreme summer and winter seasons during precession minima is the most likely candidate for triggering the regional avulsion phase. Another unknown is whether the trigger for regional avulsion occurred in exact phase with precession extremes or if triggering occurred somewhere in the transitional stages between extremes.