Sep 24, 2016 - the 'Tuttle' and 'Yuma' sandstones in the uppermost part, which include ..... University, Baton Rouge. Miscellaneous Publication 71-1, p. 3-10.
49th Annual Meeting of AASP - The Palynological Society Post-Meeting Field Trip September 24, 2016
Paleocene - Eocene Stratigraphy and Paleontology of East-Central Texas: Wilcox Group - Claiborne Group - Jackson Group Guidebook
Chris Denison, Astra Stratigraphic
Thomas Demchuk, RPS and Jen O’Keefe, Morehead State University
North Geologic Map of Field Trip Area. Edited from the Geologic Atlas of Texas, Austin Sheet.
Paleocene - Eocene Stratigraphy and Paleontology of East-Central Texas: Wilcox Group - Claiborne Group - Jackson Group Strata SUMMARY The 4 locations we will visit today provide an opportunity to examine and discuss depositional environments and stratigraphic relationships in parts of the Wilcox Group, Claiborne and Jackson groups as developed in Central Texas. These stops are frequently visited by field parties from local colleges, societies and the petroleum industry, as reservoir analogs and as potential equivalents for subsurface units. Spatial and temporal relationships presented in this field guide have been developed over several years and numerous visits to the locations we will see today, combined with critical appraisal of the many publications concerning the Paleogene stratigraphy of Central Texas and beyond, but it remains a work-in-progress. We will see the succession in stratigraphic order – Stop 1 is at the boundary of the uppermost Calvert Bluff Formation and the lower part of the Riverside Formation (new informal name) of the Wilcox Group .Stop 2 is at the upper part of the Riverside Formation and the lower part of the Carrizo Formation, the lowest formation of the Claiborne Group. Stop 3, at the renowned Whiskey Bridge macrofossil collecting locality at the Brazos River, is the type locality of the Stone City Formation, upper part of Claiborne Group. Stop 4 is at Lake Somerville, where part of the Manning Formation of the Jackson Group, is exposed in the spillway INTRODUCTION In the lower part of the Wilcox Group, a marine origin for the Solomon Creek and Caldwell Knob formations is unequivocal. The Hooper Formation, still poorly known, has a diverse marine trace fossil assemblage at its base, but the upper part is tidal to possibly non-marine. Adams (1957) noted bi-directional current directions and glauconite in the Simsboro of Bastrop County and Boenig (1970) interpreted shoreface, tidal channels, barriers and back barrier for the Carrizo in Milam County, but these have been largely overlooked. Fisher & McGowan (1967) applied a Mississippi delta model, emphasizing swamp, lacustrine and overbank fresh-water environments for the lignite-bearing Calvert Bluff, and fluvial channel deposition for the Simsboro and Carrizo formations. The Mississippi delta model was later applied to the lignite-bearing Jackson Group (Fisher et al., 1970). This model strongly influenced subsequent interpretations (Bammel 1979; Hamlin 1983; Ayers, 1989; Ayers & Lewis, 1985; English, 1988). In the intervening Claiborne Group, consistent interpretations of marine depositional environments intercalated with non-marine or marginal marine units, are based primarily on macrofossils (e.g. Stop 3), with sporadic records of foraminifera (Feray, 1948; Gaskell, 1988; 1990; Sams, 1991; Sams and Gaskell, 1990)
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Today we will discuss the evidence from sedimentary structures and palynomorph assemblages that Wilcox and Jackson outcrops in Central Texas are much more marine, and tidal influences are more widespread along the outcrop belt, than the bulk of the literature suggests. In a freshwater-dominated, microtidal Mississippi delta scenario, which is still widely accepted, this widespread occurrence of marine palynomorphs, and sedimentary structures generated in marine environments is inexplicable: an abundance of tidal sedimentary structures is only preserved in mesotidal or macrotidal regimes, whereas in microtidal regimes fluvial, wave and storm processes obliterate most tidal sedimentary structures. Critical to understanding these Paleogene depositional environments is the separation of North and South America until final closing of the Panama isthmus in the Pliocene (Coates and Obando, 1996). Until that time, the proto-Caribbean was affected by the ocean tidal bulge as it moved across the continuous body of water linking the Atlantic and Pacific oceans. On the wide, shallow shelf of the proto-Central Texas coast, and north into the Mississippi Embayment, the tidal range was amplified to meso-tidal, resulting in widespread tidal depositional signatures through the Wilcox, Claiborne and Jackson groups. Preliminary observations of the Miocene Oakville Sandstone at La Grange suggest that this is another tidally-dominated system. Tidal influences have yet to be reported from the lower part of the Calvert Bluff, but reevaluation is required. In the upper part, sedimentary structures show that tidal deposition was widespread along the outcrop belt (Bryer, 1987, 1989; Galloway, 2002; Klein, 2000; O’Keefe et al., 2005; Sturdy, 2006). The four localities visited today are all marine deposits. At Stop 1, the lower ‘Riverside Formation’ consists of distal shelf parasequences, whereas at Stop 2 the upper ‘Riverside Formation’ is more proximal muddy to sandy tidal flats. The overlying Carrizo is a tidal delta system. At Whiskey Bridge (Stop 3), the Stone City Formation is a shallow marine unit of glauconitic, fossiliferous sandstone, interbedded with marine siltstones. Finally, at Lake Somerville spillway, we will see shallow shelf marine parasequences, typically capped by lignite. In summary, for the up-dip outcrop belt of the Wilcox and Jackson groups, there is a growing body of evidence that non-marine deposits are a minor part of the succession; the majority of sedimentation is marine, in various tidal to shallow nearshore settings. In a more regional context, if the up-dip succession is more-or-less continuous, as is commonly portrayed, the predominantly shallow marine and tidal deposits present conceptual problems in having equally continuous coeval shelf and shelf edge deltaic systems – shoreface, tidal flats, tidal channels – some tens of kilometers outboard from the outcrop belt (Ambrose, 2015, Olariu, 2015). The up-dip succession is probably very fragmentary, representing only parts of highstands. During significant time gaps (Miall, 2014), on the order of several Myr, such as we will see at Stop 2, sediment would translate across the current outcrop belt and coastal plain, to supply shelf edge deltas and the deep offshore. Up-dip and down-dip deposits are not coeval: more time is probably represented in down-dip successions (see Chart 1). Given this, biozonal schemes developed from outcrops cannot be applied to down-dip deposits. Sediment provenance studies of outcrops have little or no bearing on the provenance of down-dip deposits.
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Implications for palynology Using the accepted environmental interpretations of the time for the Wilcox Group, Elsik (1968) considered that “Palynomorphs recovered from sediments deposited in closed environments such as bogs, swamps, marshes with no major drainage or stream pattern, and landlocked lakes often reflect only the local flora, and are, therefore, indicative of the environment of deposition”. Nichols (1970) focused on pollen assemblages from lignites in his pioneering palynological study along the Calvert Bluff outcrop belt in Texas, but did record dinocysts and acritarchs in ‘clay’ samples from central and south Texas. In the Calvert Bluff of northeastern Texas, Klein (2000; O’Keefe et al., 2005) concluded that tidal sedimentary structures and marine microplankton from enclosing sediments pointed to a dominance of brackish to marine depositional conditions, and Brooke et al. (2015) has demonstrated that salt-marsh indicator fungi occur in the tops of at two lignite seams in the Calvert Bluff. Gennett (1995; 1996) reported very rare dinoflagellates from the Manning Formation, and possible brackish influences in the upper part of a lignite seam indicated by abundant (55%) Cupuliferoipollenites. Sancay (2000) found greater numbers of dinoflagellates in the Manning, primarily Wetzeliella, as well as possible mangrove-indicator taxa (Raymond, pers. comm.). Given the emerging weight of evidence that sediments enclosing the lignites are mainly marine in origin, the environmental requirements of various plant species or groups based on pollen recovered from those sediments (e.g. Gennett et al., 1986; Klein, 2000; Sancay, 2000; O’Keefe et al., 2005), need to be carefully re-evaluated. A circular argument, using those ‘environmental requirements’ to identify other freshwater environments, should be avoided. Changes in spore/pollen assemblages in supposed swamp and lacustrine settings, interpreted as being driven by local environmental factors, need to be reviewed, as most assemblages are likely to be allochthonous, transported into marine settings. At best, changes in assemblages may reflect broader changes in the hinterland climate, albeit with unknown taphonomic modifications. Select palynomorphs from the Calvert Bluff Formation (Brooke, 2005 and Klein, 2000) and Manning Formation (Sancay, 2000; O'Keefe, personal collection) are shown in plates 1 and 2, respectively. Interpretations in the captions are for discussion purposes.
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FIELD TRIP STOPS STOP 1 SH 71 and Hwy 95/Jackson Street intersection (opposite Bucees). N 30º 06.285´, W 097º 18.389´ Riverside Formation – informal new name herein (Fig. 1.1). This is Stop 1 of Yancey et al. (2013), their figure 4. Best exposures are in the gullies: most of the outcrop has a thin veneer of sand washed down from the upper sandstone unit. There is only about 0.25m of bioturbated sandstone of the uppermost Calvert Bluff currently exposed. Above, there is a transgressive lag, the overlying deposits being siltstones and very fine grained sandstones, divided into several units by intervening iron-stained lags of varying thickness. At Stop 2, the upper part of this unit consists of tan laminated siltstones, becoming heterolithic upwards and eventually sand-dominated. None of these lithologies conform with the concept of the Sabinetown Formation, hence the new name ‘Riverside Formation’ is proposed. The upper surface of the ‘Riverside’ is the erosional base of the Carrizo Formation. This field guide is not intended to be the formal description of this unit.
Figure 1.1: Riverside Formation parasequences. Pale unit towards the top is very fine grained sandstone.
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Previous accounts either attribute this unit to the Sabinetown Formation or to the Calvert Bluff Formation. By including the Sabinetown at Stop 1 as the uppermost unit of the Calvert Bluff, Yancey et al. (2013) implied member status. As originally conceived by Plummer (1932) the Sabinetown was the uppermost formation of the Wilcox Group. At its type locality (now submerged in a reservoir) on the Sabine River at the Texas/Louisiana border, he described mainly glauconitic sand in the lower part (30 feet), overlain by brown clay (40 feet), sand (25 feet) and lignitic clay (15 feet). A molluscan fauna was compared with the Bashi Formation in Alabama, which implies an NP10 (upper P6) age, within the early Ypresian. Andersen (1954) summarized the type locality as ‘having a basal member consisting of a glauconitic sand, locally referred to as the "Pierson" glauconite, and an upper member consisting of a cross-bedded glauconitic sand, not named’. A 20 cm ledge of indurated iron-stained sandstone is the ironstone and vertebrate lag unit of Yancey et al. (2012, 2013). As the vertebrate-rich lower part of the lag is not present here, it is probably very localized. Here the lag is a fine-medium grained, sub-angular, poorly sorted sandstone. Iron-staining and cementation of the lag is from diagenetic alteration of the muddy matrix component.
Figure 1.2: Base of iron-stained lag with Ophiomorpha nodosa weathered out from underlying Calvert Bluff. Ophiomorpha are weathered out from the underlying sandstone; these are very fragile (Fig. 1.2). They are truncated at the lag, rather than being part of a Glossifungites assemblage. The 30 cm of underlying fine grained moderately sorted sandstone is probably uppermost Calvert Bluff, but this cannot be confirmed here. Ophiomorpha are abundant on a bedding plane on the flat ground near the main face.
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Above the lag are mainly siltstones in slightly coarsening upwards units, each separated by a thin lag of iron-stained sandstone. At the northern part of the outcrop a sand-dominated unit is present towards the top. The lowest unit is about 4.0 m of light brown to grey fine grained siltstone, coarsening up to medium grained siltstone. There are a few concretions in the upper part. An iron-stained finemedium grained sandstone, 5 cm thick forms the base of the next unit, which is a 2 m thick grey siltstone coarsening up to very fine grained sandstone. A 2 cm thick iron-stained fine-medium grained sandstone forms the base of the next unit, a 2.5 m thick very fine grained sandstone. The uppermost iron-stained fine-medium grained sandstone, again about 2 cm thick, forms the base of the next unit. Above, to the surface, are about 2 m of grey siltstone. Interpretation: distal shelf parasequences. Given the limitations of the outcrops, detailed interpretations are speculative. Thin laminations and paucity/lack of bioturbation suggest either a fluvial dominated delta-front, or, in view of the tidal heterolithics we will see at Stop 2, these could be sub-tidal deposits, distal from coastal tidal flats or a tidal delta.
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STOP 2 RED BLUFF/GOLF COURSE N 30º 04.416´, W 097º 16.820´ Upper part of Riverside Formation, truncated by erosional base of overlying Carrizo Formation. Gravel/conglomerate at the top is ‘Quaternary high gravel’. Control of floodwaters by dams upriver has resulted in the lower part of Red Bluff now being heavily vegetated. Plate 17 in Deussen (1924) shows a very steep bluff with sparse vegetation. At river level is a pale (?sandstone) unit, capped by a prominent ledge (?transgressive lag), now covered. Harris (1957, 1962) included all of the succession at Red Bluff below the Quaternary gravels in the Sabinetown, but prior and since, the Carrizo has been recognized in the upper part of the bluff, discussion being around the precise location of the base of the Carrizo.
Figure 2.1 (left): Riverside tan siltstones coarsening up to tidal heterolithics. Figure 2.2 (right): Upward coarsening heterolithics, becoming sand-dominated. The lower part of the outcrop is Riverside. Tan finely laminated siltstones initially coarsen up slightly (Figure 2.1). As the bluff face becomes more vertical, deposits become increasingly heterolithic, with the appearance of lenticular and wavy bedded sandstones. Siltstone beds become dark grey to black, suggesting more organic content. Tidal silt/sand doublets become common. Sand net/gross increases upwards, the upper part of the Riverside being sanddominated, with some small-scale cross-bedding (Figure 2.2). Trace fossils are extremely rare in
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the siltstones, with the appearance of sparse, small Planolites and Skolithus in the upper heterolithics. Prior interpretations place the formation boundary subjectively at an increase in sand content, effectively putting the upper part of the Riverside tidal heterolithics in the base of the Carrizo. Yancey et al. (2010) concur with this view, despite noting that the lithologic boundary, i.e. the subjective increase in sand net/gross, is separated by several meters from their ‘subaerial unconformity with paleosol’, placing their sequence boundary within the basal Carrizo. Herein, the subjective boundary is rejected. Instead, a prominent Glossifungites surface marks the boundary between the Riverside and the Carrizo. This is high on the cliff face and cannot be accessed at the river, but can be readily seen in gullies adjacent to Riverside Road. Discussion: a time gap of some 4 Myr is interpreted at Glossifungites surface. The Riverside is earliest Eocene, based on the age of vertebrates from the basal lag (Yancey et al., 2012), approximately 56 Ma. The next age control is from the Marquez Member of the overlying Reklaw Formation. Deposition of the Reklaw appears to be essentially continuous with the Carrizo (Sams, 1991). The lower Newby Member is silty in the lower part and has cross-bedded sands very similar to the Carrizo in the upper part (Sams, 1991), so a case can be made for it being a backstepping upper unit of the Carrizo. Transgression with eventual deposition of shelfal marine shales does not occur until the Marquez Member, where planktonic foraminifera indicate zones P8 to P9 (Sams, 1991), an age span somewhere within 49.0 to 51.8 Ma. Taking an arbitrary 51.0 Ma for the Marquez, and a relatively brief interval of time, less than 1 Myr, span for deposition of the Carrizo tidal delta, the time gap between the Riverside and the Carrizo is roughly 4 Myr. Brown & Loucks (2009) indicate a 2.5 Myr gap. If Carrizo deposition was continuous over this 4 Myr, large volumes of sediment would have to translate across a tidal delta to deposit shoreface and tidal deposits in more distal shelf locations. Instead, it is proposed that down-dip shelfal Upper Wilcox deposits occupy this time gap. Similarly, down-dip Lower and Middle Wilcox deposits may be represented by up-dip hiatuses, as yet unrecognized. On the left is a steep, south facing bluff, formed in part by the road cut. Cemented ‘Quaternary high gravel’ caps the outcrop, protecting the less indurated Carrizo (Fig. 2.3). The Carrizo is fine grained cross-bedded, locally sigmoidal cross-bedded sandstone. Dominant current orientation is to the east. Where the Carrizo is at road level, near the top of the hill, robust Ophiomorpha nodosa occur (Fig. 2.4), some 3-4 m above the base Carrizo.
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Figure 2.3: Carrizo overlain by ‘Quaternary high gravel’. Robust Ophiomorpha nodosa occur to the right of the person for scale. The basal Glossifungites surface appears at road level as we proceed downhill, and can be traced laterally Initially the uppermost part of the Riverside is a structureless grey siltstone; this is probably completely bioturbated. Vertical to sub-vertical sand-filled Thalassinoides (not roots) penetrate up to 1.0m into the Riverside (Fig. 2.5). More of the Riverside is exposed as the road descends.
Figure 2.4: Robust Ophiomorpha nodosa in the lower part of the Carrizo.
Figure 2.5: Glossifungites surface at base Carrizo, with Thalassinoides penetrating into the underlying Riverside.
The Carrizo basal Glossifungites surface, and the upper part of the Riverside can be examined in several deep gullies. From sand-dominated tidal flats there is a rapid transition to about 0.75m of
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highly bioturbated (BI6) muddy sandstone (Fig. 2.6). The Cruziana Ichnofacies is indicated by abundant small diameter Ophiomorpha, with Planolites, Skolithus and possible Asterosoma in a churned background.
Figure 2.6: Highly bioturbated uppermost part of the Riverside; sand infilling some burrows is orange colored. A combination of dark (?organic–rich) burrow walls, sand filling some burrows, with grey/brown muddy sandstone fill of other burrows, suggests this was originally a heterolithic laminated deposit before being churned. This highly bioturbated muddy sandstone transitions rapidly upwards into 1.0m of featureless, probably completely bioturbated grey/brown siltstone. Where the Glossifungites surface at the base of the Carrizo is well exposed in these deep gullies, vertical to sub-vertical sand-filled Thalassinoides (not roots) penetrate as much as 1.0m into the Riverside. From such limited exposures, the origin of this distinctly open marine succession is highly speculative. It may represent a marine transgression across the Riverside tidal flats. The Anchor Mine Tongue in the Sego Sandstone of Utah may be an analog. Good exposures of Carrizo sedimentary structures are hard to find, even when surfaces are scraped clean. Liesegang are often more prominent than bedding. However, there are a few examples of silt rip-up drapes on toes of sigmoidal beds (Fig. 2.7).
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Figure 2.7: Silt rip-up drapes (oxidized, orange) on bedding surfaces. Interpretation: with a basal Glossifungites surface, tidal bedding signatures and marine trace fossils all pointing to a marine depositional environment, the Carrizo is not a fluvial channel deposit, but is more likely a prograding tidal delta. Some significant basal tidal incision may be present regionally, but it is essentially disconformable at this location, with no clear evidence of down-cutting. In a sequence stratigraphic context, the Carrizo is probably part of a Highstand Systems Tract (HST). Dickey & Yancey (2010) and Yancey et al. (2013) mis-interpreted the elongate Thalassinoides as roots, and the featureless siltstone as a paleosol, at an unconformable sequence boundary, and the overlying Carrizo as a back-stepping marine Transgressive Systems Tract (TST). Golf Course exposure Riverside Formation incised by Pine Forest Formation, both truncated by erosional base of overlying Carrizo Formation (Figure 2.8). Proceeding up the path, on the right note the tidal flat heterolithics of the Riverside, the Glossifungites surface at the base of the Carrizo, and the more uniform sandstones of the Carrizo above. Towards the top of the hill there is a channelized incision into the Riverside. Note that the basal Carrizo Glossifungites surface is continuous across the top of this channel-fill. Pine Forest Formation. Inclined, discontinuous sandstones, siltstones and carbonaceous shales. Small trace fossils are rare to very rare, possibly indicating marine influence.
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Figure 2.8: Pine Forest channel fill incising into Riverside. Carrizo above, with basal Glossifungites surface.
Figure 2.9: Basal Glossifungites surface with sand-filled burrows penetrating into underlying carbonaceous material.
The uppermost central part of the channel fill becomes very rich in carbonaceous material, but this is not rooted and appears to be fragmentary, suggesting mostly or entirely drifted material. A piece of fossil wood is deeply indented, suggesting prolonged transport and partial decomposition. Where the Glossifungites surface at the base of the Carrizo is in contact with this carbonaceous material, burrows are filled with pale sand, in marked contrast to the black carbonaceous material (Figure 14). Note that this does not conform to the definition of a ‘woodground’. Although previous accounts have interpreted this channelized cut-and-fill unit as fluvial sands in a channel cut down from the basal Carrizo (Dickey & Yancey, 2010; Yancey et al., 2013), the continuous Glossifungites surface demonstrates that both the Riverside and Pine Forest formations were scalped by a marine erosion process – the scour at the base of a tidal delta (Willis 2005; Willis & Gabel 2001) - before deposition of the Carrizo. The temporal context of the Pine Forest is uncertain. It could be argued that it merely represents a tidal channel incision into tidal flats and is coeval with the Riverside, i.e. a Member. However, the carbonaceous material suggests proximity to highly vegetated overbank areas, which is quite different from anything we have seen today. Although highly speculative, the Pine Forest is possibly an up-dip expression of the Yoakum Canyon fill, and entirely younger than the Riverside. Subsurface maps of the Yoakum Canyon (Hoyt, 1959; Dingus, 1987; Conwell et al., 2015) show an extension or sidearm of the canyon extending into southern Bastrop County, about 25 kilometers from Stop 2.
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STOP 3 WHISKEY BRIDGE (STONE CITY FORMATION TYPE LOCALITY) N 30º 37.641´, W 096º 32.689´ In south-central Mississippi, Hilgard (1860) first named the Claiborne Group, with reservations that is seemed to be a subordinate member of the Northern Lignitic Group (=?Wilcox). In Texas, Deussen (1924) mapped the Claiborne, divided into three formations, as far as the Colorado River. In the Brazos Valley, the Claiborne is about 790m thick. Penrose (1889) was the first to describe the outcrop on the banks of the Brazos River as ‘Burleson Shell Bluff’. Penrose and Stenzel (1931) named the Moseley limestone in a detailed stratigraphic section along the Brazos River. Stenzel (1936) formally named the ‘Stone City beds’, about 85 feet thick. The outcrop has changed considerably since these early accounts. Controlled release of water from upstream dams means that the lower 4m of the succession is always submerged. Until the 1980’s the Crockett was exposed in a small gully; after construction of the Hwy 21 bridge, this uppermost part is now covered. Macrofossils have been documented by Stenzel et al. (1957), and Stanton and Nelson (1980). Jiang 1991 (in Yancey and Davidoff, 1991), found moderately well preserved calcareous nannofossils in only seven of 48 samples, these being from the Main Glauconite bed and the 2m of sand and shale immediately below. These indicate NP16 (Bartonian). Jones and Gennett (1991) noted a lack of substantial differences in pollen and spore assemblages through the formation; dinocysts are rare, except in one sample near the base (currently inaccessible). In contrast, McMahon (1997) recorded 52 dinocyst species as well as 157 pollen species in assemblages from core through the Stone City Formation. Despite many references to the prominent green beds as glauconitic, including the ‘Main Glauconite Bed’, mineralogical analysis by Harding et al. (2014) shows that various verdine minerals are present, predominantly odinite, a dark green serpentine-rich mixed layer clay. Few glauconitic minerals are present. The verdine facies is found in modern tropical environments of nearly normal salinities, in a depth range of 15 to 60m, commonly associated with fecal pellets,. Siderite and apatite are the cements in large burrow systems. Yancey (1995a) interpreted the Stone City Bluff section as shelf sediments deposited in shallow water depths initially, with progressive marine deepening, and a maximum flooding horizon at the top. An upward change in sedimentary structures shows a shift from tidal current-dominated deposition in the lower layers to storm-dominated deposition. He interpreted the Moseley Limestone as an amalgamated storm deposit. Trace fossils are abundant in the green beds, but much sparser in intervening siltstones. Very large Thalassinoides form a boxwork in the upper part of some green beds (Fig. 3.2). From trace fossils, Stanton and Warme (1971) suggested a gradual change from restricted and perhaps brackish delta-margin environments to a normal marine environment. They divided the
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succession into 11 Units. A Glossifungites surface may be present at the base of their beds 7, 9 and 11. Subsequently, Stanton (1979) defined 10 Units. Using the Units of Stanton and Warme (1971), at river level grey siltstones with sparse macrofossils is the upper part of Unit 3 (Fig. 3.1), with a thin, green, muddy sandstone at its top. Unit 6 is a sandy siltstone, underlying Unit 7, the ‘Main Glauconite Bed’.
Figure 3.1: Unit 5 at river level. Unit 7, Main Glauconite Bed’ in shadow at upper left. There is evidence of numerous stratigraphic breaks in the succession. Stenzel (1936) noted numerous diastems – each glauconite bed has a break at the base, sometimes with clasts of the underlying shale. At the top, rolled fossils, including otoliths and shark teeth may indicate hardgrounds. Thornton and Stanton (1994) interpreted the Moseley Limestone as a hardground. Stenzel’s (1936) bed aa (0.2 feet thick), a ‘conglomerate of glauconitic limestone pebbles’, disconformable on the Moseley limestone, may be a transgressive lag, with Crockett Formation above. Bed ac (0.3 feet thick), a limonitic, nodular, ledge-forming limestone is also a candidate transgressive lag.
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Figure 3.2: Very large Thalassinoides at top of Unit 7 (Main Glauconite Bed) Siltstones and lenticular sandstones of Unit 10 are well exposed north of the bridge abutment (Fig. 3.3). Exposures of Unit 11, the Moseley Limestone are poor here, but there are many large blocks along the river bank.
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Figure 3.3: Unit 10 siltstones and lenticular sandstones. Unit 11, Moseley Limestone, poorly exposed above.
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STOP 4 LAKE SOMERVILLE SPILLWAY N 30º 19.078´, W 096º 31.024´ Introduction Russell (1955) interpreted Manning strata to be almost entirely terrestrial, with the exception of the ‘Tuttle’ and ‘Yuma’ sandstones in the uppermost part, which include abundant, but low diversity, brackish water or marine body fossils (Renick 1936), and fairly common examples of the ‘fucoid’ Halymenites major (?Ophiomorpha sp., ?Thalassinoides sp.). However, according to Fisher et al. (1970; their Figure 2) the spillway outcrop is in a proximal part of one of the principal delta lobes feeding sediment to the Manning delta plain. From cores at the Gibbons Creek lignite mine in Grimes County, Mathewson and Bishop (1979) interpreted two intervals of non-marine delta plain deposits, separated by an interval of bioturbated marine deposits. In the Mississippi delta analog, this marine interval is the result of sea-level rise and transgression across the delta plain, involving landward translation of a barrier bar complex, followed by shallow marine deposition. Their core descriptions are too cryptic to be re-interpreted. With this increasing evidence of widespread marine conditions in the Manning, Yancey (1995b, 1997) rejected the deltaic model as having little utility, and proposed a strandplain-barrier bar model. Wave- and storm-dominated deposition is implicit in this model, with hummocky crossbedding (HCS) as the dominant sedimentary structure during strandplain progradation, followed by ravinement and landward migration of back-barrier lagoons and the barrier bar during transgression. Sancay (2000; O’Keefe et al., 2005) examined a new core, taken just before the bend in the emergency access road to the base of the spillway, and found similar depositional sequences but interpreted them to represent perideltaic shorezone facies. Detailed sedimentological logging, in preparation for this visit, has revealed 4 marine-dominated parasequences (P1 through P4) at the Main Central (MC) outcrop. Smaller adjacent exposures, to the southwest (SW) and northeast (NE), are separated by intervals of vegetation. Bioturbation is commonplace, except in tidal flats. Very fine grained sands in the upper part of parasequences are highly bioturbated, with minor current ripples and climbing ripples. There is no HCS. Above these parasequences is a marked change in depositional style. An interval of sandstone and siltstone heterolithics is succeeded by trough-cross bedded sandstone. Both are potentially non-marine. As the spillway is oriented NE-SW, approximating the local strike direction, the three outcrops there is the opportunity to examine lateral changes along the 200m of outcrop. The most prominent of these are a white siltstone only present in the NE outcrop, capped by thin lignite, a cross-bedded channel in P1 in the main outcrop, a very thick lignite in P4 in the SW outcrop, and the siltstone/sandstone alternations towards the top of the main outcrop. There are less
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pronounced changes in thickness of the tidal flats in P1, and in the distribution of trace fossils. Overall, these changes defy summary in a single stratigraphic column. NE outcrop Two main lithological units are exposed (Fig. 4.1). A lower siltstone is light grey to white, possibly due to kaolinite clays from alteration of volcanic ash. This lower siltstone is not present at MC, suggesting that it is a possible channel fill, but the stratigraphic relationships are obscured in a vegetated area. If this the upper ash bed of Yancey (1995), then the light brown-stained top few centimeters should contain abundant sponge spicules. Fossil roots descend up to 0.25m from the overlying lignite. Note that modern roots are also common.
Figure 4.1: Light grey to white siltstone, capped by thin lignite. Tan siltstone of P1 above. A 25cm thick lignite caps the siltstone. It contains a thin discontinuous ash bed, probably a continuous ash bed originally, but subsequently reworked by currents. Lenses of ash have a white weathered exterior. Note that at MC, this lignite has thickened to 0.6m. Above the lignite is a tan siltstone, approximately 1.5, thick, becoming more grey/white upwards and coarsening to very fine grained sandstone. The siltstone may be extensively bioturbated, but no definitive traces can be discerned. The sandstone upper part is poorly exposed.
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MC outcrop Exposure is poor at the floor of the spillway, but the uppermost few centimeters of a siltstone with roots can be seen. At MC, a 0.6m lignite below P1 has a thin, discontinuous ash bed near its base. Its relationship to the thin lignite at NE is obscured by a vegetated interval, but the ash bed suggests that it is the upper part of this lignite that is not represented at Northeast. Across MC the tan siltstone in the lower part of P1 has a sharp contact with the underlying lignite (Fig. 4.2). Trace fossils are difficult to discern in the siltstone, but a few Thalassinoides have weathered out.
Figure 4.2: Lignite with ash bed near base. Sharp contact at base of tan siltstones of P1. In gullies the siltstone is dark brown to dark grey. About 1.0m above the base, trace fossils become common and diverse, including Cylindrichnus, Ophiomorpha, Asterosoma, Planolites
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and Skolithus in an otherwise churned background, representing the Cruziana Ichnofacies (Fig. 4.3).
Figure 4.3: Brown/grey siltstone in lower part of P1, below tidal flats. Highly bioturbated. A remarkable feature of P1 is an interval of tidal flat lithologies (Fig. 4.4). Lower and upper boundaries are transitional. Trace fossils become sparse. Wavy and lenticular bedding, with some tidal doublets form a unit 0.25m thick in the northeast, thickening to 0.75m to the southwest.
21
Figure 4.4: Tidal sedimentary structures in P1, Domichnus in upper part.
Figure 4.5: Overview of P1 at MC. Tidal flats are 0.5m thick, below the grey/white sandstone. Lignite cap recessive; P2 above.
Figure 4.6: Tidal flats 0.75m thick at southwestern part of MC.
22
Figure 4.8: Planar cross beds in channel in P1. Note sharp boundary with overlying white sandstone. Roots are modern.
Figure 4.7: Schaubcylindrichnus freyi in the upper part of P1; walls probably kaolinite.
The tidal flats transition upwards into white/grey very fine grained sandstone. The color is probably due to kaolinite, from degradation of volcanic ash dispersed into the sandstone. Current ripples and climbing ripples can be seen in part of central MC and at the southwestern part. Most of this sandstone is highly bioturbated, with no sedimentary structures preserved. Trace fossils indicate the Cruziana Ichnofacies. Schaubcylindrichnus freyi is a minor component (Fig. 4.7); burrow walls are probably a concentrate of kaolinite clays. Cylindrichnus is locally common.
23
Figure 4.9: Lenticular ash bed in lignite cap of P1. Within the sandstone of P1 is a channelized unit of limited lateral extent (Fig. 4.8); the southwestern pinchout of the channel margin is present, whereas the northeastern margin is obscured by vegetation. Channel fill is planar cross-bedded, very fine grained sandstone, with a grey/pale tab color, possibly from most kaolinite being winnowed by currents. P1 is capped by a lignite 0.35m thick. An ash bed in the upper part has been heavily modified by currents into a series of discontinuous lenticular bodies (Fig. 4.9). P2 is thinner than P1 and lacks tidal flats at the transition from siltstone to very fine grained sandstone. Tan weathering siltstones coarsen up to very fine grained, grey to white, extensively bioturbated sandstone. Trace fossils are abundant (BI6) in the sandstone. Cylindrichnus, Teichichnus, small Ophiomorpha and Asterosoma are components of the Cruziana Ichnofacies. There is no lignite cap to P2. P3 has a sharp lower contact with top P2. Tan siltstone, probably highly bioturbated, coarsens up to very fine grained sandstone. Abundant trace fossils (BI6) indicate the Cruziana Ichnofacies. The upper part is ?silica cemented, hard. The next 3.3m has been covered by gravel and rip-rap to control erosion. This is probably an interval of lignites and siltstones, but this cannot be confirmed.
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Figure 4.10: P4 cemented, bioturbated, very fine grained sandstone.
Figure 4.11 Glossifungites surface
Above the covered interval is 1.85m of ?silica cemented, hard, very fine grained sandstone (Fig. 4.10). This is interpreted as the upper part of P4, the lower part being in the covered interval. Darker tan sand in burrows at 14.15m indicates a Glossifungites surface (Fig. 4.11), related to temporary cessation of deposition. This unit is highly bioturbated (BI5), but some bedding is preserved. At the top of this unit are numerous, unidentified trace fossils. These are curvilinear, up to 0.25m long, sub-vertical to almost horizontal. Traces do not intersect. Walls are mostly smooth: along their length, there are striae and/or elongate irregularities, but there are some areas of possible scratch marks. In crosssection probably flattened ovoidal to irregular. Traces apparently descend from the upper surface of the sandstone. Possibly crab or shrimp burrows.
25
Figure 4.12: Elongate trace fossils, top P4.
Figure 4.13: Elongate trace fossils, top P4.
Forming the top of P4, a 0.8m thick lignite caps the sandstone. Above the lignite that caps P4 there is a complete change in depositional style. Two units can be differentiated, a lower sand/silt heterolithic unit and an upper sandstone dominated unit (Fig. 4.14). Both units can be seen at MC and SW, but exposure is of limited lateral extent. Sand/silt heterolithic unit At MC, there is 2.9m of interbedded pale grey siltstone and very fine grained sandstone. The grey siltstone is featureless. The sandstone is either planar cross-bedded or is current rippled. Both are markedly lenticular, with pronounced changes in thickness across the limited outcrop. Trace fossils are absent. Upper sandstone unit Only about 0.75m of pale grey very fine to fine grained sandstone is exposed, in a rather broken and weathered outcrop. Fractures have been differentially cemented, resulting in a rhomboidal to polygonal surface. Sedimentary structures are obscured, except for clasts of the underlying pale grey siltstone in the basal part. These clasts suggest an erosional base, but this cannot be confirmed.
26
Figure 4.14: Lignite (weathered brown) capping P4, overlain by lenticular siltstone/sandstone heterolithics. Grey sandstone at top includes siltstone clasts from underlying heterolithics. SW outcrop Between MC and SW is an interval covered in vegetation, including trees that block the view. However, the hard, cemented sandstone that forms the upper part of P4 can be traced through this intervening interval, until it re-appears at SW outcrop. However, less than 1.0m is present, and this thins rapidly, into a swale or scoop (Fig. 4.15). Some beds in the lower part are continuous, but others are truncated at the swale. The upper bed thins laterally, but is continuous across the swale. The sandstone is highly bioturbated (BI5), including unidentified curvilinear trace fossils. As at MC, these appear to descend from the upper surface of the sandstone.
27
Figure 4.15: Hard, cemented sandstone at top P4. Continuous beds in lower part; truncated beds above, with upper bed that thins rapidly into swale. Lignite (upper right) fills the swale.
Figure 4.16: Tan siltstone with branches/small trunks. Possible roots.
28
At the base of the lignite is a thin, tan colored siltstone. On this siltstone, and impressed into the underlying sandstone are branches or pieces of small diameter tree trunks (Fig. 4.16). Some small circular features may be cross-sections of roots, but this has not been confirmed on vertical faces of the sandstone.
At SW the lignite is about 2.0m thick, more than 1.0m thicker than at MC. In there is a more resistant bed, with some siltstone and fossil wood at about 1.0 above the base (Fig. 4.17). Potentially, the lower 1.0m is filling the swale, and the 1.0m above the siltstone is correlative with the lignite capping P4 at MC. As at MC, above this thick lignite that caps, P4 there is a complete change in depositional style. Two units can be also be differentiated at SW, a lower sand/silt heterolithic unit and an upper sandstone dominated unit. Sand/silt heterolithic unit At SW, the outcrops are poor. There is approximately 2.3m of interbedded pale grey siltstone and very fine grained sandstone. The grey siltstone is featureless. The basal sandstone includes siltstone rip-up clasts; sandstones are either planar cross-bedded or current rippled. Both are markedly lenticular, with pronounced changes in thickness across the limited outcrop. In the upper part, inclined beds suggest a cut-and-fill relationship. Trace fossils are absent. Upper sandstone unit
29
About 1.5m of pale grey very fine to fine grained sandstone is exposed in a gully. Clasts of the underlying pale grey siltstone in the basal part suggest an erosional base. Additional clasts, including some sheet-like pieces, can be seen in the upper part of the gully and onto the flat areas above. Trough cross bedding is present in the gully exposure and on the flat areas above. Trace fossils have not been observed. Discussion The four coarsening upwards cycles exposed in the spillway are clearly prograding marine parasequences, with generally high BI, and trace fossil assemblages of the Cruziana Ichnofacies. In general, each cycle has a sharp base, an interval of bioturbated siltstone which gradually coarsens upwards to a fine grained sandstone that is generally highly bioturbated, but may have some ripple lamination, and a cap of lignite, but each cycle records the particular depositional conditions in each cycle. A notable feature is the thinness of the four parasequences. P1 through P4 total 14.75m in thickness, which includes a covered interval of unknown lithologies: P1 3.15m siliciclastics, capped by 0.4m lignite. P2 2.35m siliciclastics, no lignite. P3 3.0m siliciclastics, possibly capped by unknown thickness of lignite. P4 1.85m siliciclastics, possibly more in covered interval, capped by 85cm of lignite. These thin parasequences are related to low accommodation space on a passively subsiding shelf. Each could be an extensive delta lobe that prograded a considerable distance across a shelf only a few meters deep. Proximal areas of each lobe became vegetated, eventually building up lignites that could represent thousands of years of elapsed time. After a hiatus, involving relative sealevel rise to generate accommodation space, lobe switching returned to deliver sediment again. P2 may be an example of a distal part of a lobe that never became sufficiently emergent for vegetation to develop. There is some indication of punctuated regression. In P1 the small cross-bedded channel sand may be the fill localized scour, probably in very shallow water, during temporary cessation of sediment supply. The upper boundary of the channel can be traced laterally into a surface with similar lithologies above and below. Very fine grained sandstone in P2 has a slight decrease in grain size, seen in the outcrop as a pronounced concavity in the vertical face. This may another temporary cessation of sediment supply, but pervasive bioturbation has homogenized the sediment. The tidal flats in P1 are enigmatic. Sedimentary structures are clearly tidal, and the BI is markedly lower than in the siltstones below and fine grained sandstones above. The tidal flat interval increases significantly in thickness from the NE to SW.
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Plate 1. Select Palynomorphs from the Calvert Bluff Formation Dinoflagellates and other algae frequently found in the laminated siltstones and mudstones: 1. 2. 3. 4. 5. 6.
Cleistosphaeridium sp. Spiniferites sp. Peridinoid cyst Unknown Lingulodinium sp. Unknown
Fungal spores common in the uppermost portions of the coal and lower portions of the laminated siltstones and mudstones: 7. 8. 9. 10.
Fusiformisporites crabbii (a saprophyte on Juncus sp.) Lacrimasporites sp. (a known salt-marsh taxon) Paraphaeosphaeria raoii (not yet formally validated; a saprophyte on Juncus sp.) Meliola niessleana (a known saprophyte on Calluna vulgaris)
Common Spores 11. Cicatricosisporites dorogensis. This particular form of Cicatricosisporites is very similar to Ceratopteris sp. spores, all of which tolerate standing water, and some of which are salt tolerant. It also has similarities to Acrostichum aureum, which is a known mangrove associate. O’Keefe has suggested, based on cluster analysis of spore occurrences, that this particular Cicatricosisporites may be salt-tolerant, as it reaches greatest abundances in the Calvert Bluff where Fusiformisporites, Lacrimasporites, and Spiniferites are also present. 12. Deltoidospora sp. This particular Deltoidospora is similar to Acrostichum danaifolium, which is highly salt tolerant. 13. Microfoveolatus pseudodentatus Common Pollen 14. 15. 16. 17. 18. 19. 20. 21. 22. 23. 24. 25.
Momipites coryloides Caryapollenites veripites Plicapollis sp. Momipites wyomingensis, tetraporate form Rhoipites sp. Nypa echinata; Nypa (Nipa) is a known salt-marsh and mangrove taxon. Thompsonipollis magnificus Proteacidites sp. Cupuliferoipollenites sp. This taxon has been suggested to be somewhat salt-tolerant Basopollis sp. Bombacacidites paulus Lanangiopollis lihokus
31
1
2
6
5
7
3
4
8
9
14
11
10
15
12 16 17
13
25 19
20
21 22
23
24 32
18
Plate 2. Select Palynomorphs from the Manning Formation Dinoflagellate found in the laminated siltstones and mudstones: 1. Wetzeliella sp. Common Spores 2. 3. 4. 7.
Cicatricosisporites dorogensis. See note about this taxon associated with plate 1. Verrucatosporites sp. Deltoidospora sp. See note about this taxon associated with plate 1. Selaginella perinata
Common Pollen 5. 6. 8. 9. 10. 11. 12. 13. 14. 15. 16. 17.
Picea sp. Ephedra distachya-type Alangiopollis sp. Rhoipites angustus Cupuliferoipollenites sp. See note about this taxon associated with plate 1. Cupuliferoidaepollenites sp. Retitrescolpites sp. Caryapollenites vetripites Momipites wyomingensis Momipites waltmanensis Intratriporopollenites stavensis Ericipites sp.
Fungal spores common in the coal and lower portions of overlying rocks: 18. 19. 20. 21.
Brachysporisporites sp. Foveodiporites sp. Lacrimasporites sp. See note about this taxon associated with plate 1. Fusiformisporites crabbii See note about this taxon associated with plate 1.
33
1 2
3 4
8
7 5
6 9
10
11 12
14
13
15 16 17
18
19
20
34
21
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