A sequence stratigraphic model for the organic-rich Upper Devonian Duvernay Formation, Alberta, Canada

Sedimentary Geology 387 (2019) 152–181

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A sequence stratigraphic model for the organic-rich Upper Devonian Duvernay Formation, Alberta, Canada Levi J. Knapp ⁎, Nicholas B. Harris, Julia M. McMillan Department of Earth and Atmospheric Sciences, 1-26 Earth Sciences Building, University of Alberta, Edmonton, Alberta T6G 2E3, Canada

a r t i c l e

i n f o

Article history: Received 20 February 2019 Received in revised form 21 April 2019 Accepted 22 April 2019 Available online 27 April 2019 Editor: Dr. B. Jones Keywords: Devonian Duvernay Sequence stratigraphy Mudstone Black shale Basin circulation

a b s t r a c t The Upper Devonian Duvernay Formation of western Canada is a prolific source rock that in recent years has become an important exploration target for shale gas and oil. We present a sequence stratigraphic model for the Duvernay Formation that characterizes relationships between sea level, basin circulation, lithofacies distribution, and mechanisms for sediment transport and organic-enrichment. The study demonstrates lateral and stratigraphic variability in the character of sequence stratigraphic surfaces and systems tracts and interprets this variability in the context of sea level cycles at varying time scales. Correlation of 24 core descriptions to a network of 759 wells with wireline logs led to the creation of a basin-scale sequence stratigraphic framework. Three 3rd order depositional sequences are identified; the middle sequence straddles a 2nd order maximum flooding surface. Each systems tract shows unique trends in lithofacies, rock properties, and wireline character. Transgressive systems tracts are characterized by the increasingly widespread deposition of clay-poor, siliceous, organic-rich mudstones in the basin, resulting from decreased clastic and carbonate input, concentration of pelagic organic matter and siliceous radiolaria, and low oxygen concentrations in basin center bottom waters. Highstand systems tracts are characterized by sustained radiolarian input and increases in carbonate sediment due to shedding of carbonate detritus during aggradation and progradation of platform margins and reefs. Lowstand systems tracts are more clastic-rich and organic-poor. Near reef complexes, coarse carbonates locally mark the base of the lowstand systems tract. Bottom currents reworked mud- and silt-sized material in toe-of-slope locations and contributed to slope-parallel progradation during regression. © 2019 Elsevier B.V. All rights reserved.

1. Introduction Sequence stratigraphy is a powerful tool for predicting the spatial distribution of sedimentary bodies and the nature of the contacts between them. Sequence stratigraphic methods are widely employed for shallow-water sedimentary units (e.g. Van Wagoner et al., 1990; Bhattacharya and Walker, 1991; Hunt and Tucker, 1993; Shanley and McCabe, 1994; Sprague et al., 2002; Zecchin and Catuneanu, 2013); however, their application to fine-grained, basinal sediments is much less developed. The challenges of applying sequence stratigraphy to basinal sediments stem from 1) the reduced effect of sea level change on basinal sediments compared to shallow water deposits, 2) the difficulty in identifying sequence stratigraphic surfaces and systems tracts in fine-grained deposits, and 3) a lack of understanding of depositional processes and controls for organic-rich mudstone deposition. Sea level change has more subtle effects on basinal lithofacies than on shallow water deposits (Herbin et al., 1995; Hemmesch et al., 2014). A sea level fall or rise represents a proportionately smaller ⁎ Corresponding author at: JOGMEC (Japan Oil, Gas and Metals National Corporation), Technology Research Center, 1-2-2, Hamada, Mihama-ku Chiba-city, Chiba 261-0025, Japan. E-mail address: [email protected] (L.J. Knapp).

https://doi.org/10.1016/j.sedgeo.2019.04.008 0037-0738/© 2019 Elsevier B.V. All rights reserved.

change in total water depth and distance to shoreline for basinal shales than coarser shallow water sediments, and as such, the effect on lithofacies is reduced. Subtle changes in the character of basinal lithofacies may result in minimal changes in wireline log signatures, which are commonly used to extrapolate core-based sequence stratigraphic observations. Identification of sequence stratigraphic surfaces and systems tracts in fine-grained sediments has been hindered by difficulties in observing important features, fabrics, and surfaces. Mudstone outcrops are often fissile due to weathering, making it challenging to observe sedimentary structures at a fine enough scale to characterize mudstone lithofacies. By carefully observing slabbed and polished drill core, critical changes in lithofacies are in some cases observable (e.g. Savrda et al., 1985; Egenhoff and Fishman, 2013; Knapp et al., 2017; Borcovsky et al., 2017), but in the absence of macroscopic grain size changes, the use of petrographic thin sections becomes vital (Macquaker et al., 2007). Diagenetic overprinting may hinder the identification of primary lithofacies but also provide stratigraphic markers associated with sea level changes (e.g. Raiswell, 1987; Lash and Blood, 2004, 2014). Integration with organic and inorganic geochemical analyses further enables the identification of important depositional or diagenetic changes (e.g. Harris et al., 2013; Turner et al., 2016).

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Application of sequence stratigraphy to organic-rich mudstone successions has potentially significant benefits for oil and gas exploration and production by improving predictions of critical rock properties such as organic richness, porosity, permeability, and brittleness. Petrophysical and geomechanical properties vary as functions of concentration and fabric of rock-forming components such as organic carbon (reported as total organic carbon - TOC), biogenic silica, and carbonate minerals (Schieber, 1996; Schröder-Adams et al., 1996; Loucks and Ruppel, 2007; Dong et al., 2015, 2017, 2018b). Inflow of nutrientrich bottom currents, which directly influence organic carbon accumulation, have been shown to vary predictably within a sequence stratigraphic context (Harris et al., 2018). Bottom currents (sometimes described as contour currents) differ from turbidity currents in that they are relatively persistent (varying strength over tens to thousands of years; Stow et al., 2008), and generally flow parallel to slope (Rebesco et al., 2014). These same currents transport, rework, sort, and concentrate silt- and sand-sized grains in mud-dominated environments (Knapp et al., 2017) and understanding how these processes vary relative to sea level change has significant implications for reservoir quality prediction. In order to shed light on these issues, this study investigated the geographic and stratigraphic relationships between lithofacies distribution and sea level change during deposition of the Upper Devonian (Frasnian) Duvernay Formation. Development of the Duvernay Formation as an unconventional gas and liquids reservoir has been underway since 2011 (Preston et al., 2016), but very little literature exists on its sedimentological and stratigraphic characteristics (Stoakes, 1980; Knapp et al., 2017; Wong et al., 2017). Sequence stratigraphic studies have been carried out on time-equivalent Leduc Formation reef strata (e.g. Chow et al., 1995; Van Buchem et al., 1996, 2000; Whalen et al., 2000a, 2000b; Potma et al., 2001; Whalen and Day, 2008; Wong et al., 2017), and a regional framework exists for the full Frasnian system of Alberta (Potma et al., 2001; Wong et al., 2017), but no detailed regional sequence stratigraphic framework for the Duvernay Formation has been published. The sequence stratigraphic framework presented here is built on detailed lithofacies analyses from drill core (Knapp et al., 2017) and correlation of surfaces through an extensive network of wireline logs. 2. Geological setting The Duvernay Formation was deposited in the Alberta Basin (part of the larger Western Canadian Sedimentary Basin) during the Late Devonian (Figs. 1, 2). During this time, global sea level was significantly higher than at present day (Johnson et al., 1985; Savoy and Mountjoy, 1995; Haq and Schutter, 2008), resulting in widespread flooding of continental margins and development of epicontinental seaways. Many organic-rich Devonian successions were deposited within these epicontinental settings, including the Duvernay, Horn River, Bakken, Marcellus, Chattanooga, Ohio, New Albany, and Woodford shales. During the Frasnian, the Western Canadian Sedimentary Basin was a passive margin at the western edge of North America. Sedimentation was dominated by open marine shales in British Columbia and the Northwest Territories, transitioning to shallow water carbonates in Alberta, and restricted dolomites and evaporites in Saskatchewan and Manitoba to the southeast (Ziegler, 1967; Switzer et al., 1994). In Alberta, the Frasnian section consists of thick, extensive reef complexes, with thick accumulations of basin-filling shales between the reefs (Switzer et al., 1994). Duvernay Formation sediments were deposited during the time of maximum flooding within a 2nd order depositional sequence (Potma et al., 2001; Wong et al., 2017) and encompass a 2nd order maximum flooding surface (Fig. 2). The late Givetian to Frasnian 2nd order depositional sequence is bounded by regional unconformities: a lower unconformity in the Givetian that is overlain by the Watt Mountain Formation's Gilwood Member sandstone, and an upper unconformity

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near the Frasnian – Famennian boundary that is overlain by the Graminia Formation siltstones (Potma et al., 2001; Wong et al., 2017) (Fig. 2). According to Wong et al. (2017), the base of the Duvernay Formation occurs within the conodont zones Montagne Noire 6 of Klapper (1989) and punctata of Ziegler and Sandberg (1990). The top of the Duvernay Formation occurs within the Montagne Noire 11 zone and at the contact of the jamieae and Early rhenana zones. Based on these ages, deposition of the Duvernay Formation represents approximately 3 million years. The Duvernay Formation conformably overlies the mostly organiclean mudstones of the Majeau Lake Formation throughout most of the basin (Switzer et al., 1994). In the far south and west, Majeau Lake Formation sediments were not deposited, and Duvernay Formation strata unconformably onlap Swan Hills Formation platform carbonates and older Devonian strata (Switzer et al., 1994). An informal lithostratigraphy, introduced by Andrichuk (1961) and used by industry, divides the Duvernay Formation into siliceous lower and upper members, and a calcareous middle member. Thick, basin-filling, organic-lean highstand shales of the Ireton Formation conformably to unconformably overlie the Duvernay Formation across much of the basin. Near reef complexes and the Grosmont Platform, Leduc Formation reef-margin strata and Grosmont Formation platform carbonates, respectively, overlie the Duvernay Formation. Tectonic features that influenced sedimentation during deposition of Duvernay Formation sediments include the Peace River Arch, West Alberta Ridge, Rimbey Arc (overlain by the Rimbey-Meadowbrook reef trend), and Meadow Lake Escarpment (overlain by the western edge of the Killiam Barrier Reef) (Fig. 1). The Peace River Arch was an emergent basement-cored landmass on the northwest side of the basin, fringed by Leduc Formation reefs (Dix, 1990; O'Connell et al., 1990). The West Alberta Ridge was flooded at the start of Woodbend Group deposition, but formed a base for extensive Leduc reef complexes (Switzer et al., 1994). The Rimbey Arc is a southwest-northeast trending basement lineament that exerted a strong control on accommodation space during Woodbend Group deposition (Ross and Stephenson, 1989). During deposition of Duvernay Formation sediments, the lineament was marked by a chain of Leduc Formation reefs called the Rimbey-Meadowbrook trend that divides the Alberta Basin into a West Shale Basin (WSB) and East Shale Basin (ESB). Accommodation space was limited in the ESB compared to the WSB as a result of slower subsidence on the eastern side of the Rimbey Arc (Switzer et al., 1994). This effect was compounded by differential compaction, due primarily to shale deposition in the WSB and carbonate deposition in the ESB. The Meadow Lake Escarpment is a pre-Devonian erosional and structural feature (Oldale and Munday, 1994) that approximately coincides with the western edge of the Killiam Barrier Reef during deposition of Duvernay Formation sediments, and marks the furthest eastward extent of the Duvernay Formation (Switzer et al., 1994). Knapp et al. (2017) described two phases of Duvernay deposition (Fig. 3). An initial platform construction phase, during which platformsourced sediment built out into the basin, was followed by a flooded platform phase marked by widespread deposition of organic-rich mudstones. 3. Methods Eight drill cores (Table 1) were selected for detailed description. Cores were described at a scale of 1:10, paying special attention to lithology, grain size, sedimentary structures, bioturbation, presence and type of cements, color, and weathering style. Core selection was primarily based on total thickness of Duvernay Formation cored, core quality, and geographic distribution across the basin. An additional 16 cores were described in less detail to observe facies variations and stratigraphic surfaces over a greater extent of the basin. Few of the 16 additional cores covered the entire Duvernay Formation interval. A total of 108 thin sections were cut from 4 cores, encompassing the major lithofacies and intervals of interest. Thin sections were cut to 20 μm

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Fig. 1. Map of study area. Duvernay Formation organic-rich mudstones are the basinal equivalent of Leduc Formation reefs and Grosmont platform carbonates. Numbers correspond to described-cores listed in Table 1. Cross sections are labeled A-A′, B-B′, and C-C′.

rather than the standard 30 μm so that more detail could be observed in fine-grained facies (e.g. Macquaker et al., 2007). Most thin sections were cut perpendicular to bedding, but a subset was selected for bed-parallel thin sections to further examine the nature of aggregates and trace fossils. Thin sections were scanned using a Nikon Super Coolscan 5000 ED scanner to observe centimeter- to millimeter-scale features. Millimeter- to micrometer-scale features were analyzed under transmitted and reflected white light using a Zeiss Axio Scope.A1 petrographic microscope. Samples for geochemical analysis were cut every 1 m from 5 of the 8 cores. These 5 cores were chosen for wide geographic coverage of the basin and to represent a range of thermal maturity from immature to dry gas. A 10 cm long by 2 cm thick slab was cut from the back of the core at each sample location. Vertical splits were cut along the length of the slabs for separate analyses. Weatherford Geochemical Services

Group in Shenandoah, TX performed Leco-TOC and Rock-Eval analysis. Total organic carbon values reported in Table 2 are averages for each sampled facies from the SCL Kaybob 02-22, ECA Cecilia 11-04, GuideX Gvillee 09-06, and EOG Cygnet 08-20 cores. Esso Redwater 16-28 was also sampled for TOC but was removed from calculations of average TOC because it is much less thermally mature than the other sampled wells (McMillan, 2016) and, as such, has higher TOC values for all facies. Major, minor, and trace element concentrations were determined by inductively coupled plasma mass spectrometry (ICP-MS) at Acme Analytical Laboratories. A detailed report and interpretation of the geochemical analyses of these samples is presented in McMillan (2016) and Harris et al. (2018). Surfaces identified in core were correlated regionally through an extensive network of wells with wireline logs. Correlations presented here represent sequence stratigraphic surfaces rather than lithostratigraphic

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boundaries. Correlations were made in Geoscout, primarily using gamma radiation and resistivity logs, but also sonic and density logs where necessary and available. A total of 759 wells with wireline logs were used in creation of cross sections and maps.

4. Lithofacies analysis A brief overview and interpretation of lithofacies character is presented here and in Table 2 and Fig. 4. The reader is directed to Knapp et al. (2017) for details of the lithofacies analysis. LF1 through LF3 are organic-rich siliceous (high ICP-MS SiO2) mudstones, siltstones, and wackestones. LF1 planar laminated siliceous mudstone (Fig. 4A) is the most organic-rich (average TOC = 3.4 wt%) and fine-grained lithofacies, deposited through a combination of suspension settling, bottom currents, and turbidity currents, based on a dominance of mud-supported fabric, presence of dispersed aggregates of organic matter and clay (e.g. “marine snow” of Macquaker et al., 2010a), and intermittent presence of traction structures. Thin planarparallel siltstone laminae with normal and inverse grading suggest occasional sediment transport by bottom currents and turbidity currents. LF1 is not bioturbated to weakly bioturbated (BI 0-1). LF2 contains organic-rich mudstones similar to LF1 but with interbedded silty-sandy laminae and thin beds (Figs. 4B, 5) that display planar-parallel lamination, low angle cross-lamination, lenticular- to wavy- bedding, starved ripples, gradational to sharp erosional bases, normal and inverse grading, sharp (non-erosional) upper contacts, flame structures, and soft sediment deformation. This array of sedimentary structures has been attributed to bottom currents (e.g. Shanmugam, 2000) and turbidity currents. Siliceous radiolaria are abundant in LF1 and the mudstone interbeds of LF2. Most silt-sized grains in LF1 and LF2 are carbonate rather than siliciclastic. LF2 is not bioturbated to sparsely bioturbated (BI 0–2). Rarely, large (N10 cm thickness), rounded, laterally discontinuous carbonate concretions are found within LF1 and LF2 (Fig. 6). LF3 (Fig. 4C) is a minor facies composed of styliolinid-rich mudstones and wackestones. It displays sharp-based and normally graded beds and abrupt contrast in grain size, fossil assemblage, and bioturbation intensity (BI 2–4) relative to under- and overlying beds. It is interpreted to have been deposited from weak turbidity currents. Based on the sparse bioturbation features, organic-rich mudstones of LF1–3 were deposited in strongly oxygen-limited basinal areas, with minor increases in bottom water oxygenation and bioturbation associated with turbidity currents and bottom currents. LF4, LF5, and LF6 are moderately- to strongly-bioturbated (BI 3–4) mudstones and wackestones. Relative to LF1–3, LF4 (Fig. 4D) is more calcareous and argillaceous, contains more detrital quartz and less siliceous radiolaria, more macroscopic pyrite, and abundant in-situ benthic macrofossils such as agglutinated foraminifera, bivalves, brachiopods, and gastropods, and fragments of arthropods and bony fish. Despite the presence of a benthic community, LF4 averages 2.1 wt% TOC. LF5 (Fig. 4E) is much more calcareous, with average TOC of 1.8 wt% and a benthic community mostly limited to agglutinated foraminifera, the presence of which has been used to infer dysoxic rather than anoxic bottom waters (Milliken et al., 2007; Schieber, 2009). LF6 is composed of nodular to burrow-mottled lime mudstone and wackestone (Fig. 4F) with average TOC equal to 1.1 wt%. Abundant fine calcareous fossil detritus in cemented beds is interpreted to have been deposited by turbidity currents; however, typical turbidite features such as normal grading was rarely observed due to significant bioturbation. Carbonate cement in LF6 is interpreted to have occurred early after deposition

Fig. 2. Late Devonian stratigraphy of the Western Canadian Sedimentary Basin in central Alberta (modified from Switzer et al., 1994). The Duvernay Formation encompasses a second order maximum flooding surface (Wong et al., 2017) and is informally divided into lower (siliceous), middle (calcareous), and upper (siliceous) lithostratigraphic members.

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Fig. 3. Duvernay Formation depositional model modified from Knapp et al., 2017. The model describes a platform construction phase during which platform-sourced sediment built out into the basin, followed by a flooded platform phase marked by widespread deposition of organic-rich mudstones.

based on deformation of uncemented mudstone beds around nodules and presence of mineral-filled brittle fractures in nodules. LF7 and LF8 are more argillaceous, dolomitic, organic-lean, and locally rich in anhydrite. Sediments of LF7 (Fig. 4G) are thinly-bedded

argillaceous-dolomitic mudstones with minor bed-bound and burrowfilling anhydrite and sparse to strong bioturbation (BI 2–4). The presence of sharp-based, normally-graded beds have been interpreted as sediment-gravity flows similar to deposits described by Bhattacharya

Table 1 Name and location of described cores. Length is described Duvernay section. Map #

Well Name

UWI

Sampled

Detailed Description

Length in Duvernay (m)

1 2 3 4 5 6 7 8 9 10 11 12 13 14 15 16 17 18 19 20 21 22 23 24

GuideX Gvillee 09-06 BPC et al. Smoky HT 04-36 Xerex SturLks 07-22 Enermax Panther SturLs 14-02 AOSC Grizzly 01-24 SCL HZ Kaybob 02-22 Chevron Chickadee 03-05 Imperial Virginia Hills 06-36 ECA Cecilia 11-04 CNRL HZ Edson 01-10 Imperial Cynthia No. 09-06 SCL HZ Ferrier 03-21 Penn West Pembina 10-17 Imp Cdn-Sup Norbuck 02-06 Imp Cdn-Sup Tomahawk 16-18 Forgotson Burk SGSpike 10-04 Sarcee et al. Pibroch 10-16 Imperial Deep Creek 04-33 Imperial Figure Lake 11-19 Tex et al. Lucky 09-09 Esso Redwater 10-27 Esso Redwater 16-28 Nexxtep 07-05 EOG Cygnet 08-20

100/09-06-076-23W5/00 100/05-36-072-01W6/00 100/07-22-069-21W5/00 100/14-02-069-21W5/00 100/01-24-061-23W5/00 100/02-22-063-20W5/00 100/03-05-062-16W5/00 100/06-36-063-12W5/00 100/11-04-058-23W5/00 100/01-10-052-17W5/00 100/09-06-052-11W5/00 100/03-21-040-07W5/00 100/10-17-045-06W5/00 100/02-06-047-04W5/00 100/16-18-052-05W5/00 100/10-04-051-27W4/00 100/10-16-061-26W4/00 100/04-33-068-22W4/00 100/11-19-062-18W4/00 100/09-09-061-18W4/00 102/10-27-057-21W4/00 102/16-28-057-21W4/00 102/07-05-050-25W4/00 100/08-20-038-28W4/00

Y N N N N Y N N Y N N N N N N N N N N N N Y N Y

Y N N N Y Y N N Y N N Y Y N N N N N N N N Y N Y

14.4 11.4 12.6 9.6 34 54 34 17 59.5 25 12.6 42.5 27 19.5 14 11.8 8.6 28.5 13.4 9.7 55.4 57.1 9.6 46.8 Total = 628

Table 2 Lithofacies characteristics for LF1–10. Lithofacies are described in detail in Knapp et al. (2017). Text in italics indicates that taxa may be present in-situ. LF1

LF2

LF3

Planar laminated siliceous mudstone

Wavy laminated silty Styliolinid mudstone wackestone

Bedding

Planar lamination

Wavy / lenticular bedding

Other Sed. Struct.

Rare normal, inverse grading in siltstone laminae

Grain Size

Mudstone

Silt-Sand Grains

Calcite, dolomite silt, less common quartz, rare chert

Calcite, dolomite silt-sand, less common quartz, rare chert

Clay-Size Grains

Siliceous, calcareous, organic-rich

Bioturbation Index

Common Taxa

Cements

Avg. TOC Thin Section

LF5 Bioturbated calcareous mudstone Poorly bedded, wispy/wavy

LF6

LF7

LF8

LF9

Nodular wackestone

Argillaceous-dolomitic mudstone

Anhydrite-bearing dolowackestone

Intraclastic packstone

Nodular to burrow-mottled

Poorly laminated

Burrow-mottled to nodular

Massive to planar laminated

-

-

-

Normal grading, uncommon current ripples, cross lam

Normal grading, poor sorting

Poorly bedded to massive

Massive to wispy

Uncommon normal grading

-

-

-

Wackestone

Mudstone

Mudstone

Mudstone-wackestone Mudstone

Variable. Most commonly calcite, dolomite silt

Calcite, dolomite, quartz silt

Calcite, dolomite silt, uncommon quartz

Siliceous-calcareous, organic-rich

Calcareous, variably organic-rich

0-1

0-2

Radiolaria, styliolinids, tentaculitids Variable calcite, silica cement in laminae. Localized pyritic laminae (framboidal, carbonate grain replacement) 3.4 wt% Yes

Planar parallel lam, low angle cross lam, starved ripples, loading/flame structures Silty mudstone, siltstone

Mudstone-wackestone Packstone Calcite clasts, fossil frags. Less common quartz, dolomite silt, siliceous/pyritic/phosphatic clasts and fossil frags. Common mud rip ups

LF10 Limestone breccia

Pebble to boulder

Calcite, dolomite silt, less common quartz

Clay mineral aggregates? Calcite/dolo-mite silt?

Dolomite, calcite

Siliceous-calcareous, Calcareous, less organic increased clays. Variably organic-rich matter

Calcareous, low but variable organic-richness

Clay minerals, dolomite, calcite?

Dolomite, calcite, clays -

2-4

3-4

3-4

3-4

2-4

3-4

0-4

0

Radiolaria, styliolinids, tentaculitids

Radiolaria, styliolinids, tentaculitids, brachiopods, bivalves, crinoids

Styliolinids, tentaculitids, conodonts, brachiopods, bivalves, gastropods, agglutinated forams

Styliolinids, tentaculitids, arthropods, conodonts, agglutinated forams

Styliolinids, tentaculitids, brachiopods, bivalves, crinoids

-

Brachiopod, crinoid

Styliolinids, tentaculitids, brachiopods, bivalves, crinoids, gastropods, amphipora

-

Common calcite cement in laminae

Variable calcite cement

Pyrite concentrated in traces

Variable calcite cement

Heavily calcite-cemented nodules, spar-filled fractures

Minor-moderate anhydrite in laminae

Abundant anhydrite as Calcite, dolomite, minor burrow fills and evaporites nodules

2.4 wt% Yes

2.3 wt% Yes

2.1 wt% Yes

1.8 wt% Yes

1.1 wt% Yes

N.A. No

N.A. No

N.A. Yes

Calcite

Calcite

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Name

LF4 Bioturbated siliceous-calcareous mudstone

Calcite

N.A. No

157

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and MacEachern (2009), Macquaker et al. (2010b), and Aplin and Macquaker (2011). Other beds display both normal and inverse grading, suggestive of waxing and waning energy of bottom currents (Rebesco et al., 2014). Potential drivers for these currents include wind, runoff,

precipitation and evaporation rates, temperature gradients, and mixing of water bodies (e.g. Kump and Slingerland, 1999). LF8 consists of nodular to burrow-mottled (BI 3–4) dolomitic mudstone-wackestone (Fig. 4H). Anhydrite is very common as

Fig. 4. Photo plate summary of lithofacies. A) LF1 planar laminated siliceous mudstone with calcareous silty laminae. B) LF2 wavy-laminated silty mudstone. C) LF3 siliceous-calcareous styliolinid wackestone. D) LF4 bioturbated siliceous-calcareous mudstone. E) LF5 bioturbated calcareous mudstone. F) LF6 nodular mudstone-wackestone. G) LF7 thinly-bedded argillaceous-dolomitic mudstone H) LF8 anhydrite-bearing dolowackestone. I) LF9 intraclastic packstone. J) LF10 limestone breccia.

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burrow-filling cement and less commonly as anhydrite nodules. Evaporite minerals in LF7 and LF8 are interpreted to have an early diagenetic origin based on their occurrence as burrow-filling cements and replacements of dolomite nodules, and the lack of cross-cutting textures that would indicate late diagenetic replacement (Hardie et al., 1985). Coeval to Duvernay deposition, but not studied here, mixed dolomites and primary evaporites of the Duperow Formation were deposited in extensive shallow water platform areas approximately 200 to 300 km southeast of Duvernay mudstones of the East Shale Basin (Ziegler, 1967; Cutler, 1983; Switzer et al., 1994). LF9 and LF10 (Fig. 4I, J) are intraclastic packstones and limestone breccias. These lithofacies are interpreted to have been deposited by debris flows and turbidity currents, based on the presence of angular carbonate intraclasts and mudstone clasts, poor sorting, normal grading, and sharp erosional bases. Breccias are found only adjacent to reef complexes, while packstones extend further into the basin. 5. Bedsets 5.1. Results Duvernay Formation lithofacies form depositional packages that are typically 25–60 cm thick (herein “bedsets”; Campbell, 1967) and are

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typically bounded by abrupt shifts in lithofacies and TOC. Bedset thickness was typically below the typical 1 m sample spacing of geochemical measurements and therefore the abrupt shift in TOC was only directly measured where sample spacing was reduced to capture high resolution heterogeneity. Where not directly measured, abrupt TOC shifts at bedset boundaries can be inferred from observations of core color as well as average TOC values of observed lithofacies. Five common trends in stratigraphically ascending order characterize the bedsets (Fig. 7A): 1) decreasing quality (visibility and lateral continuity) of lamination and increasing bioturbation intensity; 2) siliceous lithofacies (LF1) are most abundant at the bases of bedsets, with more abundant calcareous lithofacies (LF2–6) upwards; 3) decreasing TOC; 4) increasing abundance of in-situ benthic macrofossils and transitions from agglutinated foraminifera to small bivalves to larger bivalves, brachiopods, and gastropods (although this full range is rarely expressed in a single bedset); and 5) increasing proportions of silt-sized grains. The tops of bedsets are locally heavily cemented and may be burrowed and/or scoured. This scale of lithofacies cyclicity is below the resolution of most well logs. The characteristics of bedsets vary geographically. Bedsets in central parts of the WSB typically consist of LF1 and/or LF2 with an increase in silt-sized grains towards the top of the bedset (Fig. 7B, C). In the Wild River Sub Basin, LF1–3 often grade upwards into LF4, associated with an increase in bioturbation intensity and abundance of in-situ benthic macrofossils at the top of the bedsets (Fig. 7D). In the carbonatedominated ESB, a typical bedset consists of thin LF3–5 deposits overlain by calcareous mudstones and wackestones of LF5 and LF6, typically associated with an increase in bioturbation intensity and degradation of primary lamination (Fig. 7E). In slope deposits adjacent to the Grosmont Platform, bedsets are often composed entirely of LF7 and LF8; these lack siliceous, organic-rich deposits, and bedsets grade upwards from argillaceous-dolomitic, variably laminated mudstones to burrowmottled, nodular, anhydrite-bearing dolowackestones. 5.2. Interpretation

Fig. 5. Coarse-grained beds within LF2 deposited by bottom water currents (A) and turbidity currents (B). Bottom water current deposits typically have sharp, nonerosional tops, and may have low angle cross-lamination (white arrow) and mud drapes (yellow arrow).

Trends in mineralogy, organic richness, bioturbation intensity, fossil assemblage, and proportion of silt-sized grains can be used to interpret changes in the depositional environment. An upward lithologic transition from siliceous and organic-rich to calcareous with lower TOC reflects changes in the influx of carbonate detritus to the basin. Increasing carbonate content is also associated with increasing proportion of silt-sized grains relative to clay-sized grains. Similar high frequency variation in carbonate and organic matter concentration in Upper Devonian-aged rocks has been described as a product of astronomical forcing associated with Milankovich cyclicity (e.g. Elrick and Hinnov, 2007, references therein). It is unlikely that bedsets represent large amplitude sea level fluctuations, as no evidence of significant sea level fall was observed (e.g. eroded material such as fossiliferous intraclasts or lithified phosphatic/pyritic clasts, or coarse carbonates associated with reef margin instability) within bedsets or at bedset boundaries. High frequency, minor sea level oscillations can result from thermal expansion of seawater during insolation maxima (Elder et al., 1994) and cause “highstand shedding” of carbonate detritus from reef and platform carbonates and an increase in carbonate sediment transported into the basin (Schlager et al., 1994; Vecsei and Sanders, 1997). Stronger storms during insolation maxima may also have increased carbonate transport into the basin and contributed to water column mixing and oxygenation (Eldrett et al., 2015), consistent with increased bioturbation intensity and concentration of in-situ benthic macrofossils observed in the carbonate-rich upper portions of bedsets. While the link between insolation and storm frequency and intensity is still debated, several studies suggest increased insolation or atmospheric temperatures may be associated with decreased overall frequency of cyclones and hurricanes but increased frequency and intensity of strong storms (Emanuel, 2005; Webster et al., 2005; Knutson et al., 2010). In coeval platform and reef

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Fig. 6. Laterally discontinuous carbonate concretions found within LF1 and LF2. Sediment draping of uncemented mudstone around the concretions and uncompacted laminae within concretions are taken as evidence for early diagenetic concretion growth. A) Outcrop photograph, Canadian quarter for scale. Perdrix Formation (outcrop nomenclature for Duvernay) at Roche Miette, Jasper National Park. B) Thin section scan. Core #6 (SCL Kaybob 02–22).

deposits, Wong et al. (2017) describe meter-scale shoaling upward cycles that are likely to be the shallow water expression of the 25-60 cm thick bedsets observed in basinal mudstones. Increased bioturbation intensity and concentration of in-situ benthic macrofossils upwards through a bedset and the transition from lowoxygen tolerant agglutinated foraminifera (Milliken et al., 2007; Schieber, 2009) to small bivalves to larger bivalves, brachiopods, and gastropods also suggests that conditions at the sediment-water interface, particularly oxygen concentrations, became more conducive to life for aerobes. This trend is not observed where LF1–3 comprise the entirety of bedsets, as none of these lithofacies contain in-situ benthic fossils. In such settings, bioturbation is generally only associated with turbidite and contourite beds and is not associated with relative position in a bedset. Given the limited thickness of individual bedsets, it may be unlikely that evidence for increased oxygenation observed upwards through a bedset is simply a result of sediment build-up into more oxygenated waters. It is more likely that sediment buildup was accompanied by climatically-driven changes in water-column mixing. Variations in carbonate sediment influx and bottom water oxygen concentration both likely influenced bedset-scale TOC variation, however the controls on organic matter accumulation (e.g. productivity, preservation, and dilution) vary by area (Harris et al., 2018). In basinal areas (e.g. core #6 DS2 interval, Fig. 8), radiolarian-rich, minimally bioturbated LF1 and LF2 are dominant, which when combined with high concentrations of biogenic silica (Harris et al., 2018) suggest high bioproductivity. Organic enrichment also occurred in non-anoxic settings, particularly in the Wild River Subbasin where LF4 bioturbated (BI 3–4) siliceouscalcareous mudstones have TOC values locally exceeding 3% even at high thermal maturity (dry gas) yet show ample evidence of in-situ benthic organisms. Abundant recent literature supports non-anoxic conditions in other organic-rich mudstone successions (Beier and Hayes, 1989; Schieber, 1999; Sageman et al., 2003; Rimmer, 2004; Schieber, 2009, 2011; Kazmierczak et al., 2012; Wilson and Schieber, 2015; Li and Schieber, 2015) where mechanisms for organic enrichment may include enhanced flocculation/pelletization, settling, and burial rates (Logan et al., 1995; Grossart et al., 1997; Kennedy et al., 2002; Macquaker et al., 2010a).

6. Parasequences 6.1. Results Bedsets stack to form meter-scale cyclic units (“MSUs”) that are typically 2–6 m thick and represent the finest cyclicity observable on well logs. MSUs are bounded on top by abrupt or gradational shifts from relatively more carbonate-rich lithofacies to less calcareous, more organic-rich lithofacies (e.g. LF6 overlain by LF1; Fig. 9 – Interval 1). MSUs typically display trends similar to those described above at the bedset scale: they contain a greater proportion of siliceous, organicrich deposits in bedsets near the MSU base and increasing proportions of carbonate-rich sediments in overlying bedsets. MSU tops are chosen at maximum concentration of silt or carbonate content and maximum bioturbation intensity. Carbonate hardgrounds may also coincide with MSU tops. Additionally, MSU tops are surfaces displaying maximum basinward shifts in lithofacies (Figs. 10-12). In the lower and middle Duvernay members, MSU tops are typically abrupt and carbonate cementation is common (Fig. 9 – Interval 1). MSUs in the upper Duvernay commonly have gradational tops transitioning upward from coarser- to finer-grained, from calcareous to siliceous, and from low to high TOC (Fig. 9 – Interval 2). Hardgrounds are uncommon in the upper Duvernay member. Gamma log values (Fig. 9) and sonic transit time typically decrease and density log values increase upwards through an individual MSU, coincident with decreasing TOC and increasing carbonate content, although interpretation may be complicated by well log resolution, diagenetic effects and varying contributions of clay minerals and feldspars (potassium) and organic matter (uranium) to the natural gamma ray signal. 6.2. Interpretation MSUs are interpreted here as parasequences, since they are relatively conformable successions of genetically related beds, bounded by flooding surfaces (Van Wagoner et al., 1988, 1990; Bohacs, 1998; Bohacs et al., 2014). Whereas bedsets are b1 m thick and may not be regionally extensive, MSUs are several meters thick and can be mapped

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Fig. 7. A) Summary of bedset characteristics. B) LF1-dominated, typical of the central WSB. C) LF2- dominated, typical of the platform slope toe. D) LF4-dominated, typical of the Wild River Subbasin. E) LF5- and LF6-dominated, typical of the southern ESB.

over significant portions of the basin (Figs. 10–12), suggesting that MSUs are higher rank (lower frequency; sensu Catuneanu et al., 2009, 2011; Catuneanu, 2019) parasequences. Trends in mineralogy, organic-richness, bioturbation intensity, fossil assemblage, and proportion of silt-sized grains within an MSU can be

used to interpret changes in the depositional environment. Intervals of MSUs that display decreasing organic richness and increasing carbonate content, bioturbation intensity, proportions of silt-sized grains and concentration of in-situ benthic macrofossils are interpreted to represent regression (Section 5.2 and Knapp et al., 2017). The reverse of these

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7. Depositional sequences In order to understand lateral and stratigraphic relationships between time-equivalent sediments, 4 main surfaces (SB0–3) were correlated through an extensive network of wells with available logs in the Duvernay interval. As new core observations were made, the framework was updated in an iterative process to increase the accuracy of the framework away from core control. The three intervals of rock bound by these four surfaces are named DS1, DS2, and DS3, in stratigraphically ascending order. This study follows Catuneanu et al. (2009, 2011) and Catuneanu (2019) in defining a sequence as “a succession of strata deposited during a full cycle of change in accommodation or sediment supply”. Systems tracts were characterized based on type of bounding surface, stratal stacking pattern, and position within the sequence (Van Wagoner et al., 1987, 1988, 1990; Posamentier et al., 1988; Van Wagoner, 1995; Posamentier and Allen, 1999). The 4-tract depositional sequence model, with the sequence boundary at the base of the lowstand systems tract (Hunt and Tucker, 1992) is utilized here, however falling stage deposits were not observed. 7.1. Depositional sequence 1

Fig. 8. Gamma log, deep resistivity, lithofacies, TOC, and sequence stratigraphic interpretation for core #6 (SCL Kaybob 02–22). TST = transgressive systems tract; HST = highstand systems tract; LST = lowstand systems tract; MRS = maximum regressive surface; MFS = maximum flooding surface; SB = sequence boundary. Red triangle = regression; Green triangle = transgression.

trends is interpreted as transgression. The upper bounding surface of an MSU, picked at the point of reversal in these trends, represents the switch from regression to transgression and thus represents a flooding surface (FS). The presence or absence of hardgrounds at MSU boundaries has stratigraphic significance (Section 8.3).

7.1.1. Results The oldest surface, SB0, occurs near the Majeau Lake – Duvernay Formation contact and is the lower boundary for the DS1 interval. SB0 is a sharp, planar to irregular surface that is locally bioturbated (Fig. 13). The surface is overlain by a mm- to cm-scale packstone bed composed predominantly of fragmented fossils such as styliolinids, conodonts, brachiopods, bivalves, crinoids, and other unidentified bioclasts. Also observed, especially away from central basin areas (e.g. core #4, 9, 23) are undifferentiated carbonate intraclasts, and rounded mudstone clasts that are variably pyrite- or phosphatecemented. Euhedral, centimeter-scale gypsum crystals were observed in core #9 just below the SB0 surface (Fig. 13A). Across much of the WSB, argillaceous and generally organic-lean mudstones of the Majeau Lake Formation directly underlie the SB0 surface and contrast with the siliceous, laminated, organic-rich mudstones (LF1) of the Duvernay Formation that occur above the packstone bed. The SB0 surface is typically recognized on wireline logs by an abrupt increase in resistivity (e.g. core #6, Fig. 8). In the East Shale Basin (ESB), and along the western and southern margins of the WSB, the upper part of the Majeau Lake Formation is composed of argillaceous limestones, which reduce the resistivity contrast at the contact (e.g. core #24, Fig. 14). The lower part of DS1 contains more abundant organic-rich siliceous facies (Fig. 15A), while the upper portion contains more common organic-lean calcareous and dolomitic facies (Fig. 15B). These stratigraphic variations represent the overall trends for the DS1 interval, but DS1 is composed of approximately 4 parasequences that can be locally to regionally mapped. Lithofacies of DS1 show distinct transitions with increasing distance from the Grosmont Platform (Figs. 10–12, 15A, B), coinciding with changes in thickness of the DS1 section (Fig. 15C). Adjacent to the platform, on the upper platform slope, DS1 is dominantly composed of LF8 organic-lean, anhydrite-bearing dolowackestone. Further south and west from the platform edge, DS1 thins and LF8 grades into LF7 organic-lean argillaceous-dolomitic mudstones along the platform slope. Towards the toe of the platform slope, DS1 thins to b25 m and LF6 nodular wackestones and LF4 bioturbated siliceous-calcareous mudstones dominate most of the WSB from the toe of the platform slope basinward. Near the western margins of the WSB, DS1 thins to b10 m and most of the interval is composed of laminated, organicrich mudstones of LF1, while LF6 and LF5 are only present as thin beds at the top of the interval (e.g. core #9, Fig. 10). Local thickening, coupled with low gamma log values, occurs near reef complexes where carbonate lithofacies such as LF6 are dominant (e.g. left of core #6, Fig. 10).

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Fig. 9. Core description and cyclicity summary of core #9 (ECA Cecilia 11–04). Left) Gamma log, deep resistivity, lithofacies, TOC, and sequence stratigraphic interpretation. TST = transgressive systems tract; HST = highstand systems tract; LST = lowstand systems tract; MRS = maximum regressive surface; MFS = maximum flooding surface; SB = sequence boundary. Red triangle = regression; Green triangle = transgression. Middle) Higher resolution core description of interval 1. Transgressive intervals are relatively thin and hardground/scour surfaces are common at the tops of parasequences. Right) Higher resolution core description of interval 2. Both transgressive and regressive intervals are relatively thick and hardground surfaces are not observed. Relative carbonate concentration was estimated from the color of the core face (i.e. lighter grey = higher carbonate concentration), and strength of effervescence when hydrochloric acid was applied.

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Fig. 10. Gamma log cross section A-A- in the western WSB, from the Wild River Sub Basin towards the Grosmont Platform. Datum: top of Beaverhill Lake Group. Core descriptions are included for cores #9, 6, 7, and 8.

L.J. Knapp et al. / Sedimentary Geology 387 (2019) 152–181 Fig. 11. Gamma log cross section (B-B′) through the eastern West Shale Basin. Datum: top of Beaverhill Lake Group. Core descriptions are included for cores #12, 13, 14, 15, and 17. Core #15 lacks high quality well logs and as such is paired with the gamma log of an adjacent well (2.6 km separation).

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166 L.J. Knapp et al. / Sedimentary Geology 387 (2019) 152–181 Fig. 12. Gamma log cross section (C-C′) through the eastern West Shale Basin. The 3 logs on the right side of the section are resistivity logs due to the absence of gamma logs associated with cores #19, and 20. Datum: top of Beaverhill Lake Group. Core descriptions included for cores #24, 20, 19, and 22. Core #22 is projected onto the cross section. Core #22 is directly adjacent to the Redwater reef complex and as such contains lithofacies not found in basinal locations. Core #22 provides a link to the shallow water sequence stratigraphic and biostratigraphic analysis of Wong et al. (2017), among others. In the northern ESB a lowstand systems tract was identified above SB1 by the presence of greenshales and glauconitic siltstone (Fig. 17). These clastic facies are correlative to the greenshale clasts and burrow-fills at SB1 in core #22.

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The northern ESB is similar to the WSB in that LF8 grades southward, away from the platform, into LF7 (Fig. 12). Facies transitions through the central and eastern portions of the ESB are uncertain due to lack of core and good quality wireline data. Towards the southern part of the ESB, LF6 nodular wackestone is the dominant lithofacies (e.g. core #24, Fig. 14), and the ESB in general is much more calcareous than the WSB. DS1 thins towards 5 m in central parts of the ESB. Local thickening, coupled with low gamma-log values, occurs adjacent to reef complexes (e.g. SW of core #24, Fig. 12).

Fig. 13. Core photos of the SB0 surface (red dashed line). A) Core #9. B) Core #4. C) Core #24. Coarse-grained beds overlying SB0 include fragmented fossils such as styliolinids, conodonts, brachiopods, bivalves, crinoids, and other unidentified bioclasts, as well as undifferentiated carbonate intraclasts, and rounded mudstone clasts that are variably pyrite- or phosphate-cemented. Gypsum (yellow arrow) is locally observed below SB0.

7.1.2. Interpretation SB0 at the base of DS1 is interpreted to be erosional due to its crosscutting morphology and the presence of an overlying fossiliferous and intraclastic bed that is much coarser-grained than underlying and overlying deposits (Fig. 13). The presence of phosphatic and pyritic clasts also suggests erosion of lithified sediment. Similar features have been used to interpreted sequence boundaries in other Devonian shales (e.g. pyritized ooids of Schieber and Riciputi, 2004). No significant basinward shift in lithofacies was identified that would signify a mappable lowstand systems tract at the base of DS1. However, the lag at the SB0 surface may represent a very thin lowstand systems tract deposited as upslope erosion displaced sediment that was then transported downslope. Additionally, Wong et al. (2017) describe karsted surfaces at an equivalent surface (their WD2 basal sequence boundary; correlation based on their subsurface cross sections and our description of cores #21, 22 adjacent to Redwater reef complex) in Leduc Formation reef complexes in outcrop and subsurface, suggesting periods of reduced accommodation space and sea level lowstand. Those authors also note that lowstand deposits near reef complexes are limited in extent, particularly when the lowstand deposits are wedge-shaped. Rare centimeterscale gypsum crystals just below the SB0 surface (Fig. 13A) may indicate a change in geochemical conditions associated with decreased water depth. Oxidization of pyrite in water produces sulfuric acid, which then reacts with available calcium carbonate to form gypsum (Siesser and Rogers, 1976; Briskin and Schreiber, 1978; Hoover and Lehmann, 2009; Pirlet et al., 2010; Hoover et al., 2015; McCabe et al., 2015). Thin lowstand deposits may have been reworked and winnowed during transgression in which case the transgressive surface would lie below the lag. Interpretation of the lag as a falling-stage deposit is not favoured due to the absence of any overlying sediments that may be considered lowstand deposits. The abrupt basin-scale shift to siliceous, organic-rich facies above the SB0 surface and lag from generally organic-lean and argillaceous mudstones below the surface is interpreted as transgression. The DS1 maximum flooding surface occurs at the maximum landward position of lithofacies (Figs. 10–12). In some instances, this surface is marked by a spike in gamma radiation and TOC. Subsequent regression during DS1 is interpreted as highstand rather than lowstand regression because (1) no abrupt basinward shift in facies or regionally mappable erosion surface was observed, (2) abrupt increases in clastic influx or bioturbation intensity were not observed, and (3) proportions of carbonate silt gradually increase upwards through the section. Thick DS1 deposits adjacent to the Grosmont Platform and the progradation of carbonate-rich lithofacies through the DS1 highstand suggest platform aggradation and progradation. Many authors have studied time-equivalent Leduc Formation reef complexes, and document a period of overall backstepping, equivalent to early Duvernay Formation deposition, followed by significant aggradation and increasing reef to basin relief (e.g. Klovan, 1964; Chow et al., 1995; Potma et al., 2001; Van Buchem et al., 1996, 2000; and Wong et al., 2017). Dix (1990) also recognized a pattern of ramp- to platform-style evolution of Leduc Formation sediments that fringe the Peace River Arch. Cutler (1983) suggested that Grosmont Platform stages G1 and G2 correlated with the middle Duvernay carbonate, most of which was deposited during the DS1 highstand. DS1 corresponds to the platform construction phase of Knapp et al. (2017) (Fig. 3A).

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7.2. Depositional sequence 2 7.2.1. Results The contact between DS1 and DS2 is surface SB1, which occurs near the top of or within the middle Duvernay carbonate member depending on geographic location. Over most of the WSB and ESB, SB1 is a sharp, planar to undulatory, cross-cutting surface overlain by a mm- to cmscale fossiliferous packstone to rudstone bed (Fig. 16). This fossiliferous bed is composed of fragmented fossils such as styliolinids, conodonts, brachiopods, bivalves, crinoids, and other unidentified bioclasts, as well as undifferentiated carbonate intraclasts and rounded mudstone clasts that are variably pyrite- or phosphate-cemented (Fig. 16B). At core #22, adjacent to the Redwater reef complex in the northern ESB, green argillaceous mudstones occur overlying the SB1 surface and fill burrows in underlying organic-rich mudstone (Fig. 16D). Greenshale clasts occur above the SB1 surface. In the far northern ESB, greenshales and cross-bedded glauconitic siltstones were observed directly overlying SB1 in core #20 (Fig. 17). SB1 is typically underlain by nodular wackestones of LF6, which contrast sharply with organic-rich mudstones of LF1–2 above the contact. Over much of the basin, the contact is recognized as a sharp, significant increase on gamma logs (Figs. 10-11). The contrast in gamma ray values at the surface is reduced in areas such as the southern ESB where sediments both above and below the surface are relatively carbonate-rich (e.g. core #24, Figs. 12 and 14). Stratigraphically, carbonate-rich lithofacies such as LF6 decrease towards the middle of DS2 where planar-laminated siliceous mudstones of LF1 are most widespread (Fig. 18A). In the upper sections of DS2 (Fig. 18B) LF1 grades into silty mudstones and siltstones of LF2 in the central WSB (core #6), into bioturbated siliceous-calcareous mudstones of LF4 in the Wild River Subbasin (core #9), and calcareous mudstones and wackestones of LF5 and LF6 in the ESB (core #24). As in DS1, these vertical lithofacies changes are the overall trends resulting from several (~4–7) stacked parasequences. Adjacent to the Grosmont Platform in the WSB the DS2 interval is up to 25 m thick (Fig. 18C) and is composed of LF7 argillaceous-dolomitic mudstone (Figs. 10–12), based on well log character and isopach patterns. A belt of LF7 and LF6 extends southwards along the east side of the WSB, adjacent to the Rimbey-Meadowbrook reef trend. DS2 exceeds 10 m thickness on this eastern edge. In contrast to DS1, DS2 thickness approaches zero along the platform slope in the WSB. This phenomenon is most pronounced in the northern WSB (compare Fig. 10 versus Fig. 11). Towards the toe of the platform slope, DS2 begins to thicken and is dominated by LF6 nodular wackestones and LF2 siltstones and silty mudstones. Further south and west from the toe of the platform slope in the WSB, LF2 grades into LF1 planar-laminated, organic-rich, siliceous mudstones that dominate much of the WSB during DS2. DS2 thickens towards the Wild River Subbasin, where it exceeds 35 m in thickness and LF1 grades into LF4 bioturbated siliceous-calcareous mudstones. In the ESB, DS2 exceeds 25 m thickness near the Grosmont Platform, where peloidal packstones and LF6 nodular wackestones were deposited (presence interpreted from well logs not direct core observation). DS2 thins southward to b5 m where LF6 and LF1 planar laminated mudstones are the primary lithofacies. Localized thickening occurs near reef complexes (e.g. SW of core #24, Fig. 12), particularly in the southeast ESB (Fig. 18C). Core control in the central ESB is limited. 7.2.2. Interpretation Surface SB1, which divides DS1 from DS2, is interpreted to be an erosional sequence boundary based on its regionally mappable extent, cross-cutting morphology, and overlying lag of fossiliferous

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and intraclastic grains which are variably phosphatic and pyritic and much coarser than underlying sediments. Additionally, the presence of greenshale and glauconitic siltstone directly above the SB1 surface in the northern ESB (core #20, Fig. 17) along with greenshale, greenshale-filled burrows, and greenshale clasts near Redwater reef (core #22, Fig. 12; core #21, Fig. 16D) indicates an influx of terrigenous clastics and suggests lowstand conditions. Wong et al. (2017), at an equivalent surface (their WD3.1), observe greenshale-infilled karst surfaces in reef complexes and a basinward shift of facies on the Grosmont Platform. Lowstand clastics are limited in areal extent, and in most cores lowstand sediment is limited to thin fossiliferous/intraclastic lags. Directly above the SB1 surface and lag in all but the northern ESB, lower DS2 sediments are transgressive based on the abrupt, basinscale shift in lithofacies in the landward direction. The DS2 maximum flooding surface occurs within the middle to lower portion of DS2 at the maximum landward shift in lithofacies (Figs. 10–12, 18A). The DS2 MFS is interpreted to correlate with the second order MFS of Wong et al. (2017) (details in Section 8.1). The widespread dominance of siliceous lithofacies, rich in radiolaria and organic matter, indicates limited input of both calcareous and siliciclastic sediment in the lower DS2 interval, which further supports an interpretation of transgression. Regression during deposition of the upper DS2 interval is interpreted as highstand regression because no abrupt basinward shift in facies or regionally mappable erosion surface was observed, and there were no observations of abrupt increases in clastic influx or bioturbation intensity. In the Wild River Subbasin, the upper DS2 section becomes bioturbated upward and enriched in carbonate and siliciclastic silt and clay (i.e. dominance of LF4; Fig. 10, core #9). We propose that the increase in terrigenous sediments on the western side of the basin is the result of clastic input from a western clastic source such as the exposed Peace River Arch landmass. The decrease of terrigenous sediments from the Wild River Subbasin to the Kaybob area (core #6, Fig. 8), and thinning of DS2 strata towards cores #1 and #2 suggests that the main western clastic source area may be out of the map area to the west. In contrast to DS1, DS2 is generally thinnest near the toe or lower platform slope. In the northern WSB DS2 strata thin to zero and onlap the platform slope (Fig. 10). The minimal net-deposition of DS2 sediments overlying the platform slope suggests sediment bypass, erosion, and/or significantly reduced sediment supply from the platform and other shallow water areas north and east of the study area. Limitation of NE-derived sediment due to transgression is likely, given that an abrupt landward shift in lithofacies occurs at the base of DS2 and basinal deposits have minimal carbonate and clastic content. However, erosion is also evident at the SB1 surface and the erosional surface has greatest relief where DS2 strata thin to zero (Fig. 16C). The dominance of LF2 siltstones and silty mudstones at the toe of the platform and reef slopes may have been a result of erosion of carbonaterich and early-cemented slope sediments. A possible mechanism for erosion is bottom currents; sedimentary structures indicative of traction deposition are abundant in LF2 (see Section 4). LF2 siltstones and silty mudstones are less common in the eastern WSB where DS2 thinning is less pronounced, which may correspond to reduced strength of currents that erode and rework the sediment. Additionally, northeastderived sediments prograde more strongly to the SE, parallel to the platform rather than perpendicular to it. These progradational units are dominated by LF7 argillaceous-dolomitic mudstones containing both inverse and normal grading, which may have been a result of bottom currents oriented parallel to slope, although no directional data was obtained. In addition to the influence of slope-parallel currents, thickening of DS2 along the eastern margin of the WSB may have been aided by

Fig. 14. Gamma log, deep resistivity, lithofacies, TOC, and sequence stratigraphic interpretation for core #24 (EOG Cygnet 08–28). TST = transgressive systems tract; HST = highstand systems tract; LST = lowstand systems tract; MRS = maximum regressive surface; MFS = maximum flooding surface; SB = sequence boundary. Red triangle = regression; Green triangle = transgression.

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Fig. 15. Lithofacies distribution at A) DS1 maximum flooding surface, and B) DS1 top (immediately below SB1). C) Isopach of the DS1 interval. Note: Platform edge is from Switzer et al. (1994) and is fixed for all maps.

sediment influx from the structurally-higher ESB and reef complexes of the Rimbey Meadowbrook trend. Thickening of DS2 strata towards the Wild River Subbasin is likely due to an influx of carbonate sediment from nearby reefs and clastic sediment from the exposed Peace River Arch landmass. Thickening is associated with shallowing based on a significant increase in the abundance of in-situ benthic macrofossils in LF4 bioturbated siliceouscalcareous mudstones. 7.3. Depositional sequence 3 7.3.1. Results SB2 is within the upper Duvernay member. This surface records an abrupt, basin-wide shift to more calcareous, argillaceous, bioturbated

facies (commonly LF4). Over most of the WSB, the surface is directly overlain by a mm- to cm-scale fossiliferous wackestone to rudstone bed (Fig. 19A, B, D). The SB2 surface is typically irregular as bioturbation draws fossiliferous, calcareous sediment down into underlying siliceous mudstones. Near reef complexes, the surface is typically overlain by sharp-based limestone breccias or fossiliferous to intraclastic packstone beds (Fig. 18C). Directly overlying the SB2 surface near the Redwater reef complex in the northern ESB (core #22), minor green argillaceous clays are mixed with dolomitic mudstone (Fig. 19E). Across much of the WSB and ESB, the surface is picked on well logs near the base of a significant drop in resistivity (e.g. cores #6 and 9, Figs. 8 and 9). There is some deviation from this well log character near basin margins where coarse-grained, cemented carbonates overlying the surface are more than a few centimeters thick, causing high resistivity values.

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Fig. 16. Core and thin section photos of the SB1 surface (red dashed line). A) Core #6 core photo. B) Core #6 thin section photomicrograph. C) Core #8 core photo. D) Core #21 core photo. Mixed carbonates and green argillaceous clays overlie SB1 and fill burrows extending down into DS1. Greenshale clasts occur at the top of the photo.

Adjacent to the Grosmont Platform in the WSB, the DS3 interval is up to 45 m thick (Fig. 20D) and is composed of LF8 anhydrite-bearing dolowackestone which grades basinwards into LF7 argillaceousdolomitic mudstone. LF8 and LF7 extend southward along the eastern edge of the WSB (Fig. 20A–C), considerably further than during DS2. DS3 exceeds 15 m thickness on this eastern edge of the WSB. Similar to DS2, DS3 thickness approaches zero along the platform slope in the WSB. This phenomenon is most pronounced in the northern WSB (i.e. Fig. 10 versus Fig. 11). Towards the toe of the platform slope, DS3 begins to thicken but is b5 m thick over much of the WSB. LF2 siltstones and silty mudstones are dominant at the toe of the platform slope and extend further southward and westward than during DS2, but eventually grade into LF1 planar laminated siliceous mudstones.

LF1 is present over a considerably smaller area in DS3 compared to DS2. DS3 thickens towards the Wild River Subbasin where it exceeds 20 m in thickness and LF4 bioturbated siliceous-calcareous mudstones are dominant. In the ESB. DS3 exceeds 20 m thickness near the Grosmont Platform where LF7 argillaceous-dolomitic mudstones were deposited. DS3 thins southward to b15 m in basin center areas. In the southern ESB (core #24), LF5 bioturbated calcareous mudstones and LF6 nodular wackestones are dominant, and LF1 planar laminated mudstone is minor. Significant thickening, exceeding 40 m, occurs in the south (e.g. SW of core #24, Fig. 12) and southeast ESB towards the Killiam Barrier Reef and other major reef complexes. Core control in the central ESB is limited.

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Fig. 17. Greenshale and glauconitic siltstone above SB1 in the northern ESB (core #20).

Stratigraphically, the base of DS3 is marked by an abrupt basinward shift in facies on a basin-wide scale (Figs. 10–12, 20A). LF1 at the top of DS2 is overlain by LF7 and LF8 along the eastern side of the WSB (e.g. core #14 Fig. 11), and by LF4 over much of the central and western areas of the WSB (e.g. core#6, Fig. 10). Core control is limited in the ESB but in the southern ESB (core #24), LF4 bioturbated siliceous-calcareous mudstones dominate the lower DS3 interval (Figs. 12, 14). In-situ benthic macrofossils, such as brachiopods, bivalves, gastropods and agglutinated foraminifera become abundant in the Wild River Subbasin and the southern ESB. Above the lower DS3 interval, a landward shift in lithofacies occurred. Organic-rich basinal lithofacies were deposited over a greater areal extent (Figs. 10-12, 20B) relative to the lower DS3 section but not as extensively as during DS2. Towards the top of DS3, the areal extent of organic-rich facies decreases (Fig. 20C). In the southern ESB, LF5 bioturbated calcareous mudstones are dominant above the basal LF4dominated section and transition to LF6 nodular wackestones towards the top of DS3 (core #24, Figs. 12, 14). The uppermost major bounding surface, SB3 (Fig. 21), is located near the top of the Duvernay Formation over much of the central WSB and ESB where the contact is abrupt and unconformable. Siliceous to calcareous Duvernay Formation mudstones (typically LF1–6) below the surface are typically overlain by a cm-scale fossiliferous wackestone to packstone bed (Fig. 21B, C). The surface is irregular and strongly bioturbated. The fossiliferous bed is overlain by moderately bioturbated to churned, organic-lean, argillaceous-calcareous Ireton Formation mudstones. In non-reef-proximal areas of the ESB (e.g. well #24) the SB3 surface is typically demarcated by the top of a carbonate hardground which is sharp, planar to irregular, and in places bored (Fig. 21A). Also observed were coated grains within burrows/borings and centimeterscale euhderal gypsum crystals just below the SB3 surface (Fig. 21A). Closer to basin margins, in the Wild River Subbasin (core #9) and near the Peace River Arch (core #1, 2), upper Duvernay member mudstones are more argillaceous and calcareous than siliceous basinal equivalents. In these areas the lithological contrast with overlying argillaceous Ireton Formation mudstones is reduced, and the SB3 surface is several meters above the uppermost occurrence of siliceous lithofacies (LF1–2). Near reef complexes and the Grosmont Platform (e.g. core #21, 22), SB3 also occurs well above the uppermost occurrence of siliceous lithofacies, within reef and platform carbonates. In central areas of the WSB and ESB where SB3 closely overlies organic-rich siliceous lithofacies, it is associated with an abrupt decrease in resistivity (e.g. core #6, Fig. 8). The change in resistivity is more gradual in the Wild River Subbasin (e.g. core #9, Fig. 9) and near the Peace River Arch.

7.3.2. Interpretation SB2 is regionally extensive and typically erosional, although it has a different character than SB0 and SB1. Whereas SB0 and SB1 scour and cross cut early-lithified sediments in many locations, SB2 (away from reefs) is typically marked by a bioturbated fossiliferous wackestone to rudstone bed overlying sediments that do not appear to have been lithified at the time of SB2. Burrows filled with calcareous sediment extend downwards into organic-rich mudstones. Burrows are compacted and display “mantle and swirl” fabrics (e.g. Schieber, 2003), suggesting they were formed in highly water-saturated mud before compaction and lithification. In contrast, the SB2 surface near reef complexes is sharp, unburrowed and directly overlain by packstones and breccias (Fig. 19C), indicating erosion into more competent sediment. The presence of abundant benthic fossil material overlying the SB2 surface indicates that bottom waters were conducive for aerobes, at least in upslope areas. Bioturbation at the surface, even in basinal locations indicates increased bottom water oxygen concentration. The abrupt, regional, basinward shift in lithofacies above the SB2 surface and fossil bed suggests lowstand deposition, as does the presence of mixed carbonates and green argillaceous mudstones directly overlying SB2 in the northern ESB near Redwater reef (core #22). Sediments across the basin abruptly become more argillaceous, calcareous, and bioturbated relative to sediments below the SB2 surface, indicating that clastic sediments and reef/platform-derived carbonates were more effectively transported into the basin. Additionally, there was significant southward progradation of organic-lean LF8 anhydritebearing dolowackestones and LF7 argillaceous-dolomitic mudstones along the eastern margin of the WSB indicating significantly shallower water than underlying organic-rich DS2 sediments. LF4 in western areas such as the Wild River Subbasin (core #9) is moderately to strongly bioturbated with common in-situ benthic macrofossils, suggesting that water depths were shallower during the DS3 lowstand in the Wild River Subbasin than to the northeast at the toe of the platform slope where LF2 is dominant. At a possibly equivalent surface (their WD3.2 sequence boundary), Wong et al. (2017) describe greenshaleinfilling of karst surfaces at reef complexes and quartz siltstone and sandstone on the Grosmont Platform. The top of the DS3 lowstand is interpreted as the point of maximum basinward progradation of lithofacies (maximum regressive surface, MRS; Figs. 10-12). Only a gradual and moderate landward shift in lithofacies was observed above the MRS, suggesting that transgression during DS3 was not as sudden or significant as DS1 and DS2 transgressions. Regression during deposition of the upper DS3 interval is interpreted as highstand regression because no abrupt basinward shift in facies or regionally mappable erosion surface was observed, and there were no observations of abrupt increases in clastic influx or bioturbation intensity. Similar to DS2, platform-parallel progradation and thickening on the east side of the WSB was observed for the DS3 interval, probably influenced by a similar set of factors, including slope-parallel currents, slope erosion, and reef- and ESB-sourced sediment. Compared to DS2, areas of near-zero net thickness extend further to the southeast, along with the belt of LF2 siltstones and silty mudstones. While basin circulation patterns likely contributed to minimizing sediment influx into basin center areas, increasing basin relief through the second order highstand systems tract may also have been a factor. The top bounding surface of DS3 (SB3) is interpreted as a sequence boundary based on its regional, erosional nature (Fig. 21), and overlying abrupt basinward shift in lithofacies. In the carbonate-rich southern ESB, at well #24, relative sea level fall is suggested by erosion of a carbonate hardground, filling of burrows and borings with phosphatic coated grains, and growth of cm-scale gypsum crystals, possibly associated with oxidization of previously formed pyrite (Siesser and Rogers, 1976; Briskin and Schreiber, 1978; Hoover and Lehmann, 2009; Pirlet et al., 2010; Hoover et al., 2015; McCabe et al., 2015). Other evidence for relative sea level fall at SB3 includes widespread erosion and abrupt facies shifts in reef complexes, and megabreccias (Wong et al., 2017),

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Fig. 18. Lithofacies distribution at A) DS2 maximum flooding surface, and B) DS2 top (immediately below SB2). C) Isopach of the DS2 interval.

channelized conglomerates, and turbidites (Whalen et al., 2000b) adjacent to reef margins. Quartz siltstone and green shale has been observed on the Grosmont Platform (Wong et al., 2017). 8. Discussion 8.1. Sequence hierarchy Biostratigraphic ages from Wong et al. (2017) and references therein indicate that the Duvernay Formation is bounded by their WD2 and WD4 sequence boundaries and represents a depositional time span of approximately 3 million years. Based on cross sections in Wong et al. (2017) and description of cores #21, 22 adjacent to the Redwater reef complex, WD2 is equivalent to SB0 at the base of DS1 in this study and WD4 is equivalent to SB3 at the top of DS3. Wong et al. (2017)

assign the Duvernay Formation to two 3rd order depositional sequences, designated WD2 and WD3 and locate the 2nd order MFS in their WD2 3rd order sequence; this MFS appears to be placed below the middle Duvernay carbonate member but no Duvernay Formation internal stratigraphy is mentioned. We conclude that Duvernay strata in fact represent three 3rd order sequences, designated DS1, DS2 and DS3 and are bounded by 4 sequence boundaries (SB0 – SB3). Based on the lithofacies analysis of Knapp et al. (2017) and the sequence stratigraphic analysis presented here, we have placed the 2nd order MFS in DS2 above the middle Duvernay carbonate member, at the point of maximum landward shift in lithofacies. The 2nd order MFS identified by Wong et al. (2017) is interpreted to be equivalent to our DS1 3rd order MFS. Our SB1 sequence boundary is likely equivalent to the WD 3.1 sequence boundary of Wong et al. (2017). Our SB2 sequence boundary is not

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Fig. 19. Core photos of the SB2 surface (dashed red line). A) Core #6. B) Core #9. C) Core #1. D) Core #17. E) Core #20. Mixed carbonates and green argillaceous mudstone directly overlie the SB2 surface at core #20. SB2 surface not shown.

identified as a 3rd order sequence stratigraphic surface in the scheme proposed by Wong et al. (2017) but may correspond to their WD 3.2 surface. Based on an approximate total duration of 3 million years, each 3rd order depositional sequence identified here would represent an average

duration of 1 million years. Parasequences, at a frequency of 4 to 12 per depositional sequence, would represent 4th order cyclicity with a period of approximately 83 kyr to 250 kyr. Bedsets represent 5th order cyclicity, potentially on the timescale of axial precession of Earth's rotational axis.

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Fig. 20. Lithofacies distribution at A) DS3 maximum regressive surface (top of lowstand), B) maximum flooding surface, and C) DS3 top (immediately below SB3). D) Isopach of the DS3 interval.

8.2. Sequence boundaries and lowstand systems tracts The rock properties and well log character of lowstand sediments and sequence boundaries in mudstone successions are highly variable, depending on paleogeography, change in accommodation space, rates of sedimentation, and proximity to sediment sources, among other factors. In basin-center areas, sequence boundaries typically appear as subtle thin, erosional or bioturbated beds. Phosphatic and pyritic clasts are common. Other black shale sequence boundaries have been interpreted on the basis of thin beds of pyritized ooids in Upper Devonian shales of the N.E. United States (Schieber and Riciputi, 2004) and redeposited phosphatic material in the Barnett shale (Abouelresh and Slatt, 2012). Gypsum crystals, formed by oxidization of pyrite and interaction of

sulfuric acid with calcite (Siesser and Rogers, 1976; Briskin and Schreiber, 1978; Hoover and Lehmann, 2009; Pirlet et al., 2010; Hoover et al., 2015; McCabe et al., 2015), may also indicate a sequence boundary but typically occur several centimeters below the surface. In any case, correlation of the surface on a regional scale is critical in proving that the above mentioned features are not related to local phenomena. Clastic sediments are more efficiently transported into the basin during lowstands (Catuneanu et al., 2011), although a shallowing trend may be very difficult to recognize from grain size variation if the clastic source is very fine-grained. Clastic enrichment, although potentially difficult to recognize visually in core, can be distinguished using spectral gamma and resistivity logs (Passey et al., 2010) and major

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Fig. 21. Core photos of the SB3 surface (dashed red line). A) Core #9. B) Core #10. C) Core #17. Euhedral gypsum crystals (yellow arrows) are locally observed below the SB3 surface. Millimeter-scale rounded phosphatic clasts (blue arrow) are locally observed in burrows/borings in the hardground underlying SB3.

element geochemistry. While total gamma may not vary significantly across a sequence boundary (e.g. SB2, Figs. 8 and 9), other authors have demonstrated that spectral gamma ray logs can be more responsive; organic-rich mudstones below the sequence boundary typically have high U values (Schmoker, 1981; Meyer and Nederlof, 1984; Zelt, 1985; Mann et al., 1986; Wignall and Myers, 1988; Stocks and Lawrence, 1990; Arthur and Sageman, 1994; Lüning and Kolonic, 2003; Dean, 2007; Hemmesch et al., 2014), whereas clay- and feldspar-rich mudstones in the lowstand above a sequence boundary have higher K and Th values. This is likely a common phenomenon in the Duvernay Formation at SB2, based on the petrographically observed increase in clays, and increase in ICP-MS Al2O3, K2O, and Th (McMillan, 2016; Dong et al., 2018b). Thus, due to variation in the sources of gamma radiation, variation in the standard gamma log is not always correlative to sea level changes (Bohacs, 1998; Hemmesch et al., 2014) or TOC (Lüning and Kolonic, 2003) in organic-rich mudstones. An upwards change from organic-rich to clay-rich mudstones may also cause a considerable decrease in the resistivity log at the sequence boundary (e.g. SB2, Fig. 9) as the positive influence of organic matter and hydrocarbons on resistivity (Meyer and Nederlof, 1984; Passey et al., 1990; Creaney and Passey, 1993; Passey et al., 2010) is replaced by the negative influence of clay minerals and clay-bound water (Archie, 1942; Waxman and Smits, 1968; Clavier et al., 1984). Relying on idealized well log signatures for well log correlation can lead to errors, however. Basin center gamma and resistivity response is likely to be different from the well log response nearer basin margins, particularly in mixed carbonate-clastic systems, where localized coarse carbonate detritus dominates the lowstand systems tract, or in systems where redeposited phosphatic material causes increased gamma radiation in lowstand deposits (Abouelresh and Slatt, 2012). In contrast to the Duvernay Formation, deviation from the above-described

gamma and resistivity trends can also occur in basins where organicenrichment primarily occurs in the lowstand systems tract due to a reduced connection to the open-ocean and restricted circulation (Röhl et al., 2001). In this case, gamma radiation should be highest in the lowstand systems tract as gamma radiation is produced from U-rich organic matter and K- and Th- rich terrigenous sediment. 8.3. Diagenetic expressions of stratigraphic surfaces Carbonate cementation and hardground formation is common in the Duvernay Formation and is interpreted to have occurred shortly after burial, based on deformation of uncemented mudstone around carbonate nodules (Fig. 4F), differential compaction of cemented versus uncemented laminae (Fig. 6), and cross-cutting erosion at hardground tops that pre-dates overlying mudstone deposition (e.g Fig. 16C). Hardground formation in Duvernay Formation and overlying Ireton Formation deposits was ascribed by Stoakes (1980) to reduced sedimentation rates during transgression, particularly for clastic minerals. Early carbonate cementation in organic‑carbon-rich sediments is described abundantly in the literature (e.g. Kennedy and Garrison, 1975; Mullins et al., 1980a; Curtis et al., 1986; Raiswell, 1987; Mozley and Burns, 1993; Raiswell and Fisher, 2000, 2004) and may be aided by the breakdown of organic matter in the sediments during sulfate reduction and methanogenesis, which produce bicarbonate ions and carbonate minerals respectively (Claypool and Kaplan, 1974; Whiticar et al., 1986). The contrast between abrupt carbonate-cemented MSU tops typical below the DS2 MFS, and gradational uncemented MSU tops typical above the DS2 MFS, is consistent with a change from transgression to highstand regression in the second order depositional sequence (Potma et al., 2001; Wong et al., 2017) and reflects fundamental

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differences in accommodation space and sediment supply. During the 2nd order transgression, up to the 2nd order MFS (DS2 MFS), the influx of clastic minerals into the basin progressively decreased, and carbonate platforms transitioned into more isolated reef complexes (Switzer et al., 1994; Potma et al., 2001; Wong et al., 2017). Reduced sedimentation rates, particularly of clastic sediment, at flooding surfaces (MSU tops) resulted in hiatal surfaces that were prone to carbonate cementation and hardground formation. The 2nd order highstand above the DS2 MFS marks the beginning of a gradual increase in terrigenous sediment flux (and then rapid increase in the overlying Ireton Formation) that eventually filled the basin (Wong et al., 2017). This greater flux of terrigenous sediment to the basin inhibited the development of hiatal surfaces where sharp flooding surfaces and hardgrounds could occur. Enhanced sedimentation rates acted to bury the flooding surface (MSU top) more quickly, moving it out of the early diagenetic window where carbonate precipitation is common (Curtis et al., 1986). MSUs above the DS2 MFS are thicker relative to MSUs below the DS2 MFS (Fig. 9 Interval 2 vs Interval 1), and in particular have thicker transgressive systems tracts. This suggests overall higher sedimentation rates, even during transgression, which is the time during which hiatal surfaces and condensed sections are expected to form (Mitchum, 1977; Vail et al., 1984). This phenomenon is subtly expressed within DS2, as that interval corresponds to the 2nd order transgressive maximum, but becomes increasingly apparent in DS3 (Fig. 9 – Interval 2). Carbonate cementation also occurs as discrete spheroidal concretions that are typically concentrated along specific horizons. Uncompacted laminae within concretions relative to adjacent mudstone laminae confirm that carbonate concretions formed at shallow burial depth before significant compaction had occurred. While carbonate concretion growth can be triggered by chemistry changes, basin overturns, or influx of calcareous and organic material to the sea floor (Bondioli et al., 2015; Yoshida et al., 2015), concretion growth in mudstone depositional environments is most probable when sedimentation rates are low (Brett, 2003; Dattilo et al., 2008), allowing more time for concretion growth at shallow burial depth. This relationship between concretion growth and sedimentation rate suggests that concretions are more likely to occur near maximum flooding surfaces when detrital sediment influx is low (Raiswell, 1987; Lash and Blood, 2004, 2014). Because discrete concretions are laterally discontinuous (Fig. 6), they may not be intersected by a well bore and as such are difficult to use as stratigraphic markers in a subsurface analysis. Conversely, large concretions that are penetrated by the well bore may easily be misinterpreted as laterally-continuous carbonate beds in core and well logs. The distinction of continuous and discontinuous carbonate cementation has implications for expected fracture propagation during hydraulic stimulation. 8.4. Bottom currents in mudstone-dominated successions This study presents an argument for bottom currents based on three lines of evidence. Firstly, sedimentary structures such as inverse and normal grading, sharp non-erosional tops of siltstone beds interbedded with mudstones, traction-derived ripples and cross lamination, among others suggest that bottom currents reworked sediments in environments ranging from the platform slope (LF7) to toe-of-slope and basin (LF2) (Knapp et al., 2017). Secondly, slope erosion was observed in association with the above described sedimentary structures. Erosion by bottom currents and creation of condensed sections and unconformities in carbonate environments is well-documented (Mullins and Neumann, 1979; Mullins et al., 1980b; Brunner, 1984a, 1984b; Anselmetti et al., 2000; Rendle and Reijmer, 2002; Dorschel et al., 2005; Eberli et al., 2010), although many of these studies have examined isolated platforms and associated deep water environments which are considerably different from the epicontinental waters of the Alberta Basin

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during the Late Devonian. However, minor erosion by bottom currents in the Devonian organic-rich, epicontinental Middlesex shale is documented by Schieber (1999). Thirdly, slope-parallel progradation (parallel to the platform and eastern edge of the WSB, rather than perpendicular to it) was observed, which may have been influenced by bottom currents, although no directional data was obtained from sedimentary structures. Similar patterns of current-driven progradation of sediments and platform margins are well-documented in the Bahama Banks (e.g. Mullins et al., 1980b, 1987; Wilber et al., 1990; Anselmetti et al., 2000; Betzler et al., 2014). Progradation patterns in the Alberta Basin may have also been influenced by changes in sediment source. For example, around the time of maximum flooding in DS2, northeast-derived carbonate and clastic sediment influx may have been significantly reduced due to the creation of upslope accommodation space. At this same time, and throughout DS2 and DS3, locally-sourced sediment from reefs of the Rimbey-Meadowbrook trend and the ESB may have accumulated along the eastern margin of the WSB. Increasing reef to basin relief through the second order highstand may have limited progradation further into the basin. The suggestion that circulation patterns affected sediment distribution in the Alberta Basin is not new. Stoakes (1980) inferred slope-parallel currents entering the basin from the north during Duvernay and Ireton deposition to explain the clockwise, east to west progradation direction of mudstone clinoforms – a basin-filling pattern which seems to have begun during Duvernay time at the start of the second order highstand in DS2. Similarly, Andrichuk and Wonfor (1954) suggested currents from the NW based on the clastic content of Ireton Formation mudstones. Other studies, suggesting current direction generally ranging from NW to NE, base their interpretations in distribution of reef-derived carbonate sediment (Newland, 1954; McCrossan, 1961; Andrichuk, 1961), stromatoporoids (Wendte, 1974), wave-resistant reef rims (Andrichuk, 1958; Klovan, 1964), and distribution of acid-resistant microfossils (Staplin, 1961). The effects of bottom currents in organic-rich mudstone successions have significance for reservoir rock properties in shale gas and oil production. Concentration and sorting of silt- and sand-sized material within an organic-rich mudstone succession has the potential to locally enhance porosity and permeability, and create dual permeability systems and hydrocarbon carrier beds (Schneider et al., 2011; Cronin et al., 2016). This effect may be diminished in carbonate-dominated systems such as the Duvernay, where coarse-grained beds are typically cemented. Bottom currents may also indirectly contribute positively to reservoir rock properties if they help to deflect clastic sediment away from basinal areas, allowing the concentration of pelagic organic matter and siliceous organisms. Bottom currents are also considered to have been significant to organic enrichment. Harris et al. (2018) ascribed development of elevated TOC content in the Duvernay Formation to penetration of the basin by cold nutrient-rich bottom currents from the north that enhanced bioproductivity and indirectly led to reduced oxygen levels, similar to a model proposed for the Cretaceous Western Interior Seaway (Slingerland et al., 1996). This effect was particularly significant during sea level highstands, when cold bottom waters could most effectively breach a bathymetric barrier between the Peace River Arch and the Grosmont Platform. Bioproductivity and organic enrichment have significant impacts on both petrophysical and geomechanical properties. In the DS2 and DS3 intervals of the Duvernay Formation, where evidence for bottom currents is most abundant, basinal mudstones are rich in biogenic silica sourced from radiolaria, which has been demonstrated to increase brittleness (Dong et al., 2018b) and potentially improve fracture propagation during reservoir stimulation. Similarly, porosity development in organic matter is significant, particularly at high thermal maturities (Dong et al., 2019); thus sea levelcontrolled enrichment in organic matter strongly affects porosity distribution in the Duvernay.

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The presence of bottom currents can be identified from sedimentary structures in drill core; however without a basin-scale sedimentological and stratigraphic study, the full effect of bottom currents on sediment character and distribution may be unappreciated. The impact of bottom currents on organic-rich mudstone successions is likely more common than reported. 9. Conclusion 1. Cyclicity in Duvernay Formation mudstones is recognized at three scales: 25-60 cm thick coarsening-upwards bedsets, 2–6 m thick 4th order parasequences, and three 3rd order depositional sequences that comprise the Duvernay Formation and several meters of overlying strata near basin margins and reef complexes. 2. Depositional sequence 1 records a period of aggradation and progradation of platform margin strata in the northeast part of the basin. Stratal thinning southward and westward into the basin is associated with increased proportions of siliceous, organic-rich mudstones. Depositional sequence 2 encompasses a 2nd order MFS and marks the most laterally extensive deposition of siliceous, organicrich mudstones. Depositional sequence 3 contains the only regionally extensive lowstand systems tract, within which TOC is reduced, and bioturbation intensity and clastic input is increased. 3. Below the 2nd order MFS, transgressive deposits are thin and flooding surfaces are sharp and cemented. Above the MFS transgressive deposits thicken and flooding surfaces become gradational as sedimentation rates increase in the 2nd order early highstand systems tract 4. Slope-parallel progradation and concentration of silt- and sand-sized grains was influenced by slope-parallel currents, consistent with previously published observations of sedimentary structures and sediment distribution patterns in the Alberta Basin. 5. With adequate core coverage, sequence boundaries and systems tracts can be identified and correlated in well logs. Recognition of lateral changes in the depositional environment and sediment source is critical. Acknowledgements This work was supported by Imperial Oil, Shell Canada, ConocoPhillips, Nexen, Devon Energy, and the Natural Sciences and Engineering Research Council of Canada (all on grant number CRD 445064-12), as well as Husky Energy (grant number RES0023336). Throughout the project, technical discussions with a number of individuals helped form and evaluate the interpretations presented here, including: Korhan Ayranci, Tian Dong, Noga Vaisblat, Chris Schneider, Murray Gingras, Matthew Fay, Hugh Alley, and John Weissenberger. Thanks to the editor and reviewers for their feedback which significantly improved the quality of this manuscript. References Abouelresh, M.O., Slatt, R.M., 2012. Lithofacies and sequence stratigraphy of the Barnett Shale in east-central Fort Worth Basin, Texas. Am. Assoc. Pet. Geol. Bull. 96 (1), 1–22. Andrichuk, J.M., 1958. Cooking Lake and Duvernay (Late Devonian) sedimentation in Edmonton area of Central Alberta, Canada. Am. Assoc. Pet. Geol. Bull. 42, 2189–2222. Andrichuk, J.M., 1961. Stratigraphic evidence for tectonic and current control of Upper Devonian reef sedimentation Duhamel area, Alberta, Canada. Am. Assoc. Pet. Geol. Bull. 45, 612–632. Andrichuk, J.M., Wonfor, J.S., 1954. Late Devonian geologic history in Stettler area, Alberta, Canada. Am. Assoc. Pet. Geol. Bull. 38, 2500–2536. Anselmetti, F.S., Eberli, G.P., Ding, Z.D., 2000. From the Great Bahama Bank into the Straits of Florida: a margin architecture controlled by sea-level fluctuations and ocean currents. Geol. Soc. Am. Bull. 112 (6), 829–844. Aplin, A.C., Macquaker, J.H.S., 2011. Mudstone diversity: origin and implications for source, seal, and reservoir properties in petroleum systems. AAPG Bull. 95 (12), 2031–2059. Archie, G.E., 1942. The electrical resistivity log as an aid in determining some reservoir characteristics. Paper SPE-942054-G, Petroleum Transactions, AIME 146, pp. 54–62. Arthur, M.A., Sageman, B.B., 1994. Marine black shales: depositional mechanisms and environments of ancient deposits. Annu. Rev. Earth Planet. Sci. 22, 499–551.

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