5.a. Lacustrine facies

At Clevedon, these consist of three interbedded sub-facies: 1) dominant yellow dolomites with common ooids and some siliciclastic grains, and various sedimentary structures, 2) beds, bands and nodules of Red silicified rock (chert) and 3) limited outcrops of Blue-grey marl. In addition, 4) in many parts of NW Somerset, there are Red marls (‘Keuper’ facies) but these are not exposed in the Clevedon area and are best seen at the classic Aust Cliff exposure (25 km NE) and along the North Somerset coast (e.g. St Audrie’s Bay to Blue Anchor). The yellow dolomites with the red cherts represent the sublittoral to littoral facies marginal to the extensive lake-inland sea to the south-southwest (wherein the red and blue-grey marls were deposited), and close to a hilly-rocky shoreline immediately to the east and northeast of Upper Palaeozoic strata, supplying siliciclastic material by fluvial-colluvial processes.

5.b. Fluvial-colluvial siliciclastic facies

The Dolomitic Conglomerate in the greater Bristol-Mendips area includes 1) Breccias, massive to thick-bedded, varying from clast to matrix supported with a red muddy-sandy dolomitic matrix, clasts composed of Devonian and Carboniferous sandstone and limestone-dolomite, along with clasts of Triassic soil/regolith, referred to as pedorelicts (Howson et al. Reference Howson, Tucker and Whitaker2022). These are colluvial-scree and local debris-flow deposits. 2) Bedded breccio-conglomerates and coarse sandstones: medium to thick-bedded sandy conglomerates, pebbly coarse sandstones to finer sandstones, commonly channelled with erosive bases and local cross-bedding. These are mostly stream flood deposits. 3) Well-bedded sandstones-siltstones-mudrocks, with cross-bedding, lamination and ripples, being sandy to muddy channel to floodplain deposits and 4) locally interbedded pedogenic facies (nodular calcretes) and hiatal surfaces (horizons of exfoliated-cracked pebbles on upper surfaces of breccia/conglomerate), representing extended periods of soil development and subaerial exposure. These various facies are well displayed further NE along the coast towards Charlcombe Bay, 3.5 km NE of Clevedon. See Milroy (Reference Milroy1998) and Howson et al. (Reference Howson, Tucker and Whitaker2022) for further information.

5.c. Field observations

Yellow dolomites are mostly poorly bedded with vague lamination and bedding on the mm to cm scale. In places, they display medium-scale cross-bedding (sets 20 cm high) (Figure 6 a) and flat bedding with relics of wave-formed ripples. Most beds consist of clay-grade (lime mud-dolomicrite) to sand (oolitic-peloidal) grade carbonate (Figure 6 b), with minor quartz and lithic sand grains. There are local concentrations of granules to small pebbles of basement clasts (ORS and Carboniferous Limestone), together with cm-size clasts of blue-grey and red marl (Figure 7 a). There are also thin micritic beds of mm-scale, irregular-crinkly lamination of likely microbial-mat-biofilm origin and some cm-thick micritic layers broken up into intraclasts (Figure 7 b). In places, the bedding is disturbed, manifested by flat marl clasts defining a foliation, which can be vertical. Up to 2 m of section are exposed on the Clevedon foreshore, but in the cliff below the Toll-House, on the north side of the fault, up to 8 m of yellow dolomite crop out above 2–3 m of breccia facies, becoming more oolitic upwards and seawards, with less siliciclastic material, similar to the succession at Ladye Bay, 1 km NE.

Figure 6. (a) Cross-bedded oolite, set height 20 cm, on south-facing cliff beneath Clevedon Pier. Field of view 1 m across. (b) Photomicrograph of oolitic dolomite. Field of view 2 mm across. Thin-section courtesy of Eric Squires.

Figure 7. (a) Yellow oolitic dolomite with scattered small pebbles and clasts of red and green marl-siltstone. Field of view 20 cm across. (b) Yellow oolitic dolomite with thin beds of dolomicrite disrupted into clasts. Field of view 40 cm across.

Within the yellow dolomites exposed in the mid to high part of the foreshore at Clevedon, there are conspicuous bands and nodules of extremely hard red silicified rock, which range from near-continuous flat to undulating ‘pavement’ surfaces through to discontinuous layers in the form of discrete lenses and nodules (Figures 5 a–e, 8 a, b). The cherty bands typically range in thickness from 5 to 10 cm and form a resistant cap to foreshore exposures of the much softer, more porous yellow dolomite (Figures 5 a–e). In some places, the red silicified bands are beds with coarser clastic sediment than the adjacent yellow dolomites (Figure 8 a), along with scattered small basement pebbles up to 3 cm in diameter. Elsewhere, the red bands are similar to adjacent carbonates in terms of clastic content or grain size, simply appearing as diagenetic nodules of red chert within the dolomite (Figure 8 b). There are cm-scale patches of yellow dolomite within the silicified red bands (Figure 8 a, b).

Figure 8. (a) and (b) Cut surfaces of chert bands within yellow dolomite. In (a) coarser sediment within the silicified band, an irregular base to the chert layer from erosion or loading into less dense dolomitic sediment. Cm-scale areas of yellow sediment within the chert. In (b) two cut surfaces of a large block with 2 chert bands: the lower is continuous to nodular, and the upper is seen (in the field) to be a more continuous band. The dark spots and streaks in the yellow dolomite are late diagenetic Mn-oxides and dendrites. All slabs are ∼ 10 cm across.

The bases of the red bands vary from fairly regular, near-horizontal and parallel to bedding in adjacent yellow dolomite to undulating with an apparent erosive base, to be more irregular, disrupted and deformed with downward penetrating lobes and protrusions (Figures 5 b–d, 8 a, b). In some places, there is a thin (mm), finer-grained basal layer to the silicified bands. The upper surfaces of the chert layers are smooth to undulating with hollows and ridges; locally, they appear to be distorted, with fold-like structures and local ‘gaps’ in the bands (Figures 5 b–d). These features suggest that at some earlier time, the bands were soft and ductile, unlike their present hard, brittle nature. Displacements of one band by up to 15 cm indicate local soft-sediment faulting.

In one place in a silicified red bed, there is an intriguing feature of two divergent elongate lobes about 8 cm long (Figure 9). The feature resembles structures with three lobes in similar-aged Marginal Triassic sediments in South Wales, interpreted as dinosaur footprints (Tucker & Burchette, Reference Tucker and Burchette1977). Overall, the observations of deformation, faulting and loading indicate that at a time either during or soon after its formation these layers were soft, ductile and malleable.

Figure 9. Red silicified sediment layer with a possible dinosaur footprint (20 cm long) on the upper surface showing 2 toes.

Discrete silicified nodules are present in the yellow dolomites, notably near the low tide mark of the beach at Clevedon, several metres below the horizons of the red silicified bands. The nodules, white/cream to pink-red in colour, are a few cm up to 10 cm across, round to oval to elongate in shape (Figures 10 a, b). In cut section, however, the outer surface of the nodules is cauliform (Figure 10 b). The nodules mostly consist of fine-grained quartz, but coarser, clear, colourless, glassy crystals are present in their centre, or they are hollow, like geodes. Similar nodules are present in Ladye Bay (Milroy & Wright, Reference Milroy and Wright2000).

Figure 10. (a) Quartz nodules, white to red in colour, some with central voids containing quartz crystals, within micritic-oolitic dolomite. Field of view 15 cm across. (b) Cut surface showing red quartz nodule with botryoidal outer margin. Field of view 10 cm across.

Blue-grey marl crops out over a square metre or more below the beach gravel about mid-tide, 5 m south of Clevedon Pier, 50 m west of the seawall. This outcrop appears to be within the yellow oolitic dolomites. Marl also occurs between the major fault and a minor splay fault 30 m from the seawall (Figure 4 b). In a temporary excavation in 2000, at least 80 cm thickness of blue-grey marl was observed behind the seawall above yellow dolomites and adjacent to the major fault itself (Mathieson in Ashley et al. Reference Ashley, Wright, Marriott and Lavis2024). These blue-grey marls at Clevedon Pier are dolomitic mudstones varying in appearance from fissile to blocky, in some places breaking with a conchoidal fracture. Although in the field they appear homogeneous, massive and structureless, like typical mudstones, on cut surfaces, a remarkable range of features is revealed: mm-scale lamination with silty graded layers, wispy and vaguely folded lamination, and brecciated units with rounded to angular clasts (Figure 11 a, b).

Figure 11. (a) and (b) Cut surfaces of grey-blue marl showing a range of features: Laminated, brecciated, angular to elongate clasts of different shades of grey, some red sediment. From exposure on Clevedon foreshore. Scale mm.

6.a. Yellow micritic-oolitic dolomites

This microfacies consists of both fine- and coarse-grained/crystalline varieties. The finely crystalline carbonate micrite matrix contains local to abundant carbonate grains and scattered siliciclastic grains. Numerous small dolomite rhombs, 10–20 µm diameter, commonly vague to poorly defined in thin-section, occur throughout the fine matrix, but there are also scattered larger, better-formed dolomite rhombs up to 100 µm across (Figures 12 a–c). In many places, the dolomicrite has a vague peloidal appearance (grains 250–500 µm diameter), composed of tiny (5 µm) carbonate particles. Sand-sized carbonate grains are mostly 0.5–1 mm in diameter and many consist of micritic carbonate, such that they could be peloids. These are interpreted as micritic grains of faecal or microbial origin (synsedimentary); some could be reworked and altered fragments of Carboniferous Limestone which formed the upland areas adjacent to the lake margin. The ooids themselves are round through to ovoid-subrounded in shape, mostly in the range of 300–700 µm diameter (Figures 6 b, 12 a, b). They mostly consist of dolomite crystals within an obvious outer lamella, 5–10 µm thick composed of very fine micritic dolomite. With some ooids, the outer lamella is a red coating, a few µm thick, not always continuous (Figure 6 b). The dolomite crystals within the ooids are fine to coarse, subhedral (planar-s) to non-planar in shape, varying from limpid to cloudy and poorly defined (Figures 6 b, 12 a, b). In some cases, empty spheroidal voids are inferred to be oomolds. Nuclei in the ooids are rare. These ooids show the same features as the dolomitized ooids described from Ladye Bay (Milroy & Wright, Reference Milroy and Wright2002).

Figure 12. Yellow oolitic–micritic dolomite. (a) Three dolomitized ooids (oo), 2 grains of Carboniferous Limestone (grey, cl), quartz silt and scattered dolomite rhombs, in a fine dolomitic micrite. (b) Ooid with well-defined outer lamella in fine dolomite sediment with quartz silt and patches of oxidized pyrite (dark brown). (c) Fine-grained dolomite with vague sand-sized grains, scattered larger dolomite rhombs and fragments of carbonate rafts (thin sheets). Scale bars each 100 microns.

The silicate grains are mostly of quartz with minor feldspar and are present in two size grades: silt (20–50 µm), disseminated throughout, and sand-size, 500 µm to 1.5 mm. Some beds would classify as a silty dolomite. Within the yellow carbonates are conspicuous 100–500 µm-sized areas of rusty brown iron oxide-hydroxide, likely goethite-limonite (Figures 13 a–c). They vary from spheroidal to crystal shapes (cubic-triangular-hexagonal) and appear to consist of sub-particles. Under reflected light, areas of bright yellow-silver colour observed within the brown patches are relics of pyrite (Figure 13 b). Elsewhere there are dendrite-looking features of the same rusty brown colour.

Figure 13. (a) Oxidized pyrite and dolomitized ooid in yellow micritic dolomite with quartz silt. Scale bar 250 microns. (b) Reflected light view of (a) pyrite relics in oxidized material. Scale bar 100 microns. (c) Oxidized pyrite spheres and crystal forms, 200 microns diameter, in dolomitic micrite with some larger scattered dolomite rhombs.

In areas of the foreshore close to the major fault, the yellow dolomites have numerous fractures with a spacing of 5–10 cm containing galena in patches up to a cm across. Greenish copper carbonate minerals (malachite) are rare and white to pink baryte is a common fracture fill. Pyrolusite (MnO₂) occurs as scattered mm-sized black spots in the dolomite (as in Figure 8 b) and as dendrites, branching out from the fractures. Liesegang rings are present on dolomite surfaces (Ashley, Reference Ashley2014).

Silt to fine sand-sized feldspar crystals of K-feldspar and albite were observed in the SEM (see Figures 17 a, 18 a) and could be authigenic, although rarely reported in limestone (e.g. Singewald & Milton, Reference Singewald and Milton1929; Hearn & Sutter, Reference Hearn and Sutter1985). Tiny (<5 µm) baryte crystals are evenly distributed through the sediment and minor rutile is noted (Figure 17 a).

6.b. Red silicified bands and nodules (cherts)

This microfacies mostly consists of a red-brown fine-grained matrix of microquartz with vague micritic sand-sized grains (peloids) (Figures 14 a, b), along with other grains. There are scattered well-developed dolomite rhombs, 20–50 µm in size. Some of these larger dolomite rhombs have irregular margins when in contact with the silica (Figure 17 b), indicating reaction with the fluid precipitating the silica. Around 10% of the rock consists of detrital quartz grains, minor feldspar and lithic grains, fragments of siltstone, mudrock and Carboniferous Limestone, including individual fossils, especially crinoid debris. Many grains are medium sand but larger grains are present, up to 10 mm across. Some grains are partially to fully coated with a thin dark red-brown rim (Figure 14 c, d).

Figure 14. (a) In upper part, red sediment with silicified ooids (oo, one with fractured outer lamella, red arrow) and other grains, fine silica matrix and scattered silt-sized quartz grains, and below, dark fine-grained (yellow) micritic dolomite, vaguely peloidal. (b) Red silicified sediment with vague relics of sand-sized grains and scattered quartz silt (q). (c) and (d) Red silicified sediment with grains of quartz, some of which silicified ooids (oo); one grain with partial clay coat (red arrow), plane polarised light (PPL) and crossed polars (XP). All scale bars are 500 microns.

Round sand-size grains composed of quartz crystals commonly show a drusy fabric (i.e. increasing crystal size towards the grain centre), with a thin, dark outer rim (5 µm thick) of very fine dolomite (Figures 15 a–d, 17 c, 18 b). In some cases, this outer rim has fractured and been displaced, like a broken eggshell (Figures 15 a–d), and the quartz crystals include the fractured lamella and extend outside the former grain into the original inter-grain pore space. The drusy texture of this quartz indicates it is a cement filling a void, attributed to dissolution of an aragonitic ooid. The minor fracture and breakage of the outer lamellae before the quartz precipitation indicate very early quartz precipitation just after aragonite dissolution, as compaction was beginning but before any significant burial.

Figure 15. Red silicified sediment with ooids (oo) showing clear outer lamella, in some cases fractured (red arrows), and fill of drusy quartz, which extends into adjacent original pore space, in a matrix of silicified lime/ dolomitic mud. (a) and (b): scale bars 500 microns, PPL and XP. (c) and (d): scale bars 100 microns, PPL and XP.

The red colour of this silicified sediment is attributed to finely disseminated iron oxide (hematite). There are no relics of pyrite present. One case of a gypsum pseudomorph is inferred from a swallow-tail crystal shape (500 µm across) with a fine-grained red-brown clay-micrite rim (Figure 16) and a quartz interior. Some detrital grains contain dolomite rhombs, clear to cloudy, which are floating in a medium crystalline quartz matrix. These dolomite rhombs were clearly present before the silicification; they could be reworked fragments of Carboniferous dolomite, with some early dissolution then silica cement.

Figure 16. Former gypsum crystal with a red micritic coating, now consisting of quartz in a lacey-alveolar texture (microbial/pedogenic), within red silicified sediment containing silicified ooids. PPL, scale bar 1000 microns.

7.a. Sediment geochemistry

We present major and trace element analyses of two samples of each rock type (Table 2). The two analyses of each sample type are extremely close in element contents. The most obvious and unsurprising difference between the compositions is the much higher silica content of the chert.

Table 2. Chemical data for Yellow dolomite and Red chert; oxides in % and group 1 and 2 trace elements in ppm

The main mineral assemblage of the yellow carbonate is dolomite, calcite and detrital quartz with minor feldspar and Fe oxide (see Figures 17, 18). This mineral assemblage is consistent with the major element chemistry. Dolomite and smaller amounts of calcite account for the inventory of MgO, CaO and LOI (assumed to be almost entirely CO₂). SiO₂, Al₂O₃, Na₂O and K₂O can be explained by detrital quartz and alkali feldspars. Fe₂O₃ is explained by the iron oxide-hydroxides. There is minimal S, indicating low contents of sulphide (relic pyrite) or sulphate (baryte). We can use mass balance on the average of the two compositions to calculate approximate wt% proportions of dolomite (72%) and calcite (10.5%). The Al₂O₃, Na₂O and K₂O contents are consistent with a few percent of detrital feldspar and small amounts of clay.

Figure 17. (a) Scanning electron microscope-back-scattered electrons (BSE) image of yellow dolomite sediment (point analyses 7, 8; Mg, Ca, O, C), with authigenic feldspar (point 1; Si, O, Al, K), quartz (points 2, 5; Si, O), albite (points 3, 4; Si, O, Al, Na) and rutile (point 6; Ti, Si, O). (b) SEM-BSE image of silica enclosing corroded dolomite rhombs, 10-15 microns in diameter. (c) SEM-EDS map of silicified-dolomitized ooid with calcite vein cutting through. Colours: green = Mg (dolomite), blue = Si (quartz), cyan = Al (K-spar), yellow = Ba (baryte), pink = Ca (calcite), red = Fe (iron oxide).

Figure 18. X-ray maps showing distribution of major elements for an area of (a) yellow dolomite sediment and B) red silicified sediment. Colours: green = Mg (dolomite), red = Ca (calcite), blue = Si (silica), cyan = Al (K-spar), yellow = Fe (pyrite-goethite). In (b), note the very thin dolomite outer lamella around the silicified ooids. Porosity measurements under SEM: dolomite sediment 25-35%; silicified sediment 5-6%, over an area roughly 10 mm × 10 mm.

The silicified red rocks have substantial silica (chert) (∼65%), as to be expected, and 32% dolomite. In Table 2 we list the ratio of each major element oxide in the yellow dolomite relative to that in the silicified red rocks. A key observation is that the ratio for most oxides is approximately 2 and this suggests that the system is largely closed except for the ingress of silica. Thus, the main geochemical differences between the rock types can simply be explained by the addition of silica.

Data on trace elements can be divided into two groups in terms of the ratios of each element in the two rock types. In one group (Ga, La, Nd, Rb, Zr, V, Y), the ratio is close to a value of 2. These are a diverse set of elements that in many geochemical systems might behave differently through interaction with fluids or physical separation of different mineral components due to sedimentary processes. However, whatever fluids were involved in the silicification process, they have not perturbed the geochemistry significantly for these elements beyond dilution as a consequence of the ingress of silica. There is also no significant difference in the ratio of dolomite to detrital silicate grains in the two rock types in the 2 samples analyzed. The second group of elements (As, Mo, Ba, Co, Cu, Pb, Zn, Sr) shows ratios much greater than 2 and these are likely to be those associated with the Pb–Cu–Zn mineralization. In the yellow dolomite, these elements can also be linked to the higher content of S and As as these elements are found in sulphide and sulphate mineral species (Ixer et al. Reference Ixer, Pattrick and Starkey1993). The mineralization took place well after the silicification, so the simplest explanation of group 2 elements is that they are associated with infiltration of mineralizing fluids into the porous and permeable yellow dolomite, but these same fluids were prevented from ingress into the low porosity and permeability red silicified bands. There is one notable exception in that U is enriched in the silicified red bands compared to the yellow dolomite. U solubility in aqueous fluids is very sensitive to pH (Smedley & Kinniburgh, Reference Smedley and Kinniburgh2023), reaching a minimum at about pH = 8.5, and is also somewhat sensitive to Eh and activities of carbonate-bearing and other ionic species (Goodwin, Reference Goodwin1982; Chudasama et al. Reference Chudasama, Porwal, González-Álvarez, Thakura, Wilde and Kreuzer2018; Stanley & Wilkin, Reference Stanley and Wilkin2019). Thus, the U behaviour could be explained by changing fluid composition during silicification. The observation is consistent with a decrease in pH from around 9 to 7.5 through meteoric water input and mixing with lake water (discussed below).

A final issue is the extent to which the red cherts are formed by silica filling pore space or replacement of a fine carbonate matrix by silica. Of particular significance are the ratios of CaO and MgO since these relate to one of the major mineral constituents, dolomite. We observe that the ratio is 2.21, which is slightly greater than the ratio observed in other major elements and group 1 trace elements. If we assume the silica only fills pores, then using mass balance, we find that adding silica to the yellow dolomite composition requires about 55% porosity. This interpretation can explain the typical ratio in major elements, including MgO, and Group 1 trace elements as a closed system dilution effect. If we then consider that the yellow dolomite contained 10% calcite and assume it was replaced by silica, the initial porosity would then reduce to 45%. The current porosity of the yellow dolomite is 25–35% and it has undergone burial compaction, so it seems plausible that the initial surface sediment in the lake had a porosity of 45%. Modern carbonate sediments do have porosities of these high values (oolitic sands are 45%, lime muds up to 70%; Enos & Sawatsky, Reference Enos and Sawatsky1981). If this was the case, then there is no requirement to replace the dolomite, assuming that dolomitization preceded silicification.

7.b. Stable isotopes

Although we present no new isotope data, it is instructive to discuss published analyses to inform interpretation of the depositional environment and diagenesis of the sediments. Leslie et al. (Reference Leslie, Spiro and Tucker1993) presented whole-rock data from the Tea Green Marl (and the red Mercia mudstones below) from Somerset and South Devon and Milroy (Reference Milroy1998) analysed samples from the Ladye Bay Oolite Member, all dolomites, equivalent to the yellow dolomites of Clevedon (Table 3, Figure 19). The values from the Ladye Bay oolites are all very similar, with an average δ¹³C _PDB of 0.3 ‰ _PDB (range −0.13 to +0.46 ‰ _PDB) and δ¹⁸O _PDB of 1.5 ‰ _PDB (range +1.3 to +2.2 ‰ _PDB). For the Tea Green Marl, the δ¹³C _PDB values have a small heavier average of 1.5 ‰ _PDB although there is a little more scatter (range +0.9 to +2.2 ‰ _PDB); the δ¹⁸O _PDB values average 1.2 ‰ _PDB (range 0.5 to 1.6 ‰ _PDB), very similar to the oolites.

Table 3. Carbon-oxygen isotope data (‰ _PDB) for Ladye Bay oolites (Milroy Reference Milroy1998), Tea Green Marl (Blue Anchor Fm.) from Somerset (Leslie et al. Reference Leslie, Tucker and Spiro1992) and calcrete-tufa-travertine from the Marginal Triassic, Glamorgan (Leslie et al. Reference Leslie, Spiro and Tucker1993)

Figure 19. Carbon-oxygen isotope cross-plot for Ladye Bay oolites (Milroy Reference Milroy1998), Tea Green Marls (Blue Anchor Fm.) (Leslie et al. Reference Leslie, Tucker and Spiro1992) and calcrete-tufa-travertine (Leslie et al. Reference Leslie, Spiro and Tucker1993).

Leslie et al. (Reference Leslie, Tucker and Spiro1992) presented data on calcretes and travertine-tufa from the Marginal Triassic of Glamorgan, and all δ¹³C _PDB and δ¹⁸O _PDB data therein are quite different from the oolites and marls at Clevedon, with low to very negative δ¹³C _PDB and moderately low to quite negative δ¹⁸O _PDB (average calcrete δ¹³C _PDB −4.4 ‰ _PDB, δ¹⁸O _PDB −4.8 ‰ _PDB; average travertine-tufa δ¹³C _PDB −2.7 ‰ _PDB, δ¹⁸O _PDB −5.1 ‰ _PDB).

8.a. Depositional environments and processes

The yellow dolomites possess many features consistent with deposition in a low to moderate energy location in the vicinity of a lacustrine – inland-sea shoreline. Cross-bedding and flat bedding indicate moderate wave-current activity, water depths in the region of 0.5–3 m, as in the equivalent facies exposed in Ladye Bay, 1 km to the north of Clevedon (Milroy & Wright, Reference Milroy and Wright2000). Ooids are typical of such locations and are forming today in the Great Salt Lake and Pyramid Lake, USA (Eardley, Reference Eardley1938; Popp & Wilkinson, Reference Popp and Wilkinson1983). The development of oolites in the Marginal Triassic is only documented from this eastern side of the Bristol Channel Basin, probably related to the shoreline orientation relative to prevailing winds. Periods of lower energy are indicated by finer-grained carbonates; these are likely to have been aragonitic-high magnesium calcite (HMC) lime muds. Crudely developed irregular lamination in this facies could be relics of microbial mats, as to be expected in lower energy areas of a shallow saline lake; local mud intraclasts indicate erosion of lithified crusts or subaerial exposure and desiccation. Occasional influx of terrigenous sediment through flash floods is indicated by the presence of siliciclastic silt and sand, either disseminated or in lenses and thin beds. The presence of dolomite and the dolomitization of the ooids are discussed below in the diagenesis section, but it is likely that the fine-grained dolomite was forming on the lake floor or just below within the sediment. A schematic model for deposition in this Triassic lake is presented in Figure 20.

Figure 20. Schematic model for deposition and early diagenesis of the Triassic marginal lacustrine sediment, Bristol Channel Basin.

The red silicified chert bands in some cases are associated with more clastic-rich sediment than the enclosing carbonates, and with their coarser grains up to small pebble size, the influx of sediment as events is indicated. This is consistent with the generally sharp, erosive bases of these red chert beds and the presence of clasts of yellow carbonate within these red beds as well. The influx of coarser sand and pebbles into the lake system attests to the influx of run-off waters and these could well have been in the form of a hyperpycnal flow, whereby a dense sediment-laden flow enters a water body and flows along the lake floor (e.g. Yang et al. Reference Yang, Jin, Van Loon, Han and Fan2017 for a Triassic example from the Ordos Basin) (see Figure 20). An input of terrigenous clastic material mixing with lake carbonate is envisaged in this case, resulting from a flash flood-debris flow from the hinterland following a major storm. The presence of ooids in these red chert bands would be the result of the mixing of incoming terrigenous material with oolitic sand forming in the lake. Tucker (Reference Tucker, Matter and Tucker1978) and Tucker & Burchette (Reference Tucker and Burchette1977) interpreted laterally extensive 10 cm-thick siliciclastic beds in the similar-aged Marginal Triassic of Glamorgan as sheet-flood deposits formed from shallow debris-stream flows into a lake-playa system.

However, as noted, there are also patches of red chert in the form of nodules and discontinuous bands where the sediment itself appears to be of similar grain size and composition (ooids + clastic grains) to the surrounding yellow carbonate. In these cases, the red chert appears to cut across the bedding, suggesting that the contact is a diagenetic front, indicating post-depositional silicification. With these silicified nodules, it is suggested there were still influxes of freshwater, but in the form of hypopycnal flows (flows less dense than that of the water body) transporting suspended clay-dust particles, which gave rise to the red colour (Figure 20).

The pervasive red colour of the silicified bands and nodules is attributed to detrital iron-containing clay/dust in suspension within the incoming flow. A dry-arid desert environment favours the accumulation of ferric oxide within terrigenous sediment, and the presence of ORS in the adjacent landscape would also have contributed fine-grained hematite. The Fe-rich coats around some grains could reflect pedogenesis in the hinterland.

In arid-zone lakes, sand and coarser sediment are invariably deposited relatively close to shore with finer clay in suspension being carried much farther offshore in a hypopycnal flow. This phenomenon is clearly seen in images from news outlets from the recent huge flash flood in Derna, Libya (September 2023) and has been observed by MET in Marsa Alam, Red Sea, Egypt (Tucker, Reference Tucker2003). Fine sediment in suspension can be carried great distances offshore, in low-density hypopycnal flows, travelling out over dense (saline) lake water for several km. When the clay is eventually deposited out of suspension, and this may be after several days, there will be a red clay layer on the lake floor, but without any associated coarser material, sand or pebbles. It is suggested that red chert nodules with no coarser clastic material are developed in these situations by water mixing (see below) within the yellow carbonates.

Soft-sediment deformation: Both red chert bands and dolomitic limestones show evidence of soft-sediment deformation. In a few places, the bedded dolomitic sediments are disturbed some distance away from the chert bands. The red bands themselves are conspicuously disturbed and disrupted. The shapes of the silicified bands, together with their presence as isolated nodules within the yellow dolomitic limestone, and with clasts of yellow sediment within the chert bands, suggest these sediments were ductile and easily deformable at the time. These observations are taken to imply that the lake was repeatedly an environment that was being disturbed both prior to and after formation of the red chert bands and nodules.

Syn-sedimentary deformation structures are common in lake sediments, especially those in active rift basins (e.g. Varejao et al. Reference Varejao, Warren, Simoes, Cerri, Alessandretti, Santos and Assine2022). The overall depositional environment was an extensive lake basin in a desert/semi-arid location with an upland/hilly hinterland likely with contemporaneous active faulting. Soft-sediment deformation has been reported in the Blue Anchor Fm. elsewhere (Lavernock, Glamorgan; North Somerset coast) and in the overlying Rhaetic sediments (Mayall, Reference Mayall1983), related to seismic activity, although Simms (Reference Simms2007) attributed a deformed bed in the Cotham Member to meteorite impact.

Several processes common in such active lake–basin settings could be responsible for soft-sediment disturbances. Density inversion is a common cause of sediment deformation from Rayleigh-Taylor type instabilities, as in the sudden deposition of a relatively heavy sediment load onto a soft, muddy-sandy lake floor. Also, sediment compaction resulting in layered sediments with variable permeability can lead to later dewatering and transient trapping of water lenses with outbursts when these lenses suddenly break through (Nichols et al. Reference Nichols, Sparks and Wilson1994). Additionally, a build-up of dissolved silica or silica gel (see later section) in the pore spaces of what would become the silicified rock could have resulted in a density contrast with underlying sediment, leading to instability. Seismic disturbance might also be related to nearby faults, and it is possible that the Clevedon fault itself was active at the time of lake deposition. Seismogenic shaking can trigger density instabilities, dewatering and liquefaction of unconsolidated sediments. These various mechanisms are not mutually exclusive (Nichols et al. Reference Nichols, Sparks and Wilson1994).

Possible dinosaur footprints at Clevedon would be no real surprise, since they are recorded in abundance in similar facies of similar age in Glamorgan (Tucker & Burchette, Reference Tucker and Burchette1977; Lockley et al. Reference Lockley, King, Howe and Sharp1996; Falkingham et al. Reference Falkingham, Maidment, Lallensack, Martin, Suan, Cherns, Howells and Barrett2022). Although there is only one almost convincing footprint, albeit with only 2 clear toes, but with a space where a third toe could have been, there are many features of sediment disturbance within/upon the particular footprint-bearing silicified red bed that could be attributed to ‘dinoturbation’. Dinosaurs trampling across a soft-wet sediment surface, producing ‘squelch’ marks, were invoked by Tucker & Burchette (Reference Tucker and Burchette1977) and Falkingham et al. (Reference Falkingham, Maidment, Lallensack, Martin, Suan, Cherns, Howells and Barrett2021) for the footprints at Barry and Penarth, South Wales. Two palaeobiologists familiar with footprints have visited the Clevedon site (Mike Benton, University of Bristol and Susie Maidment, BM(NH)) and both considered the evidence to be equivocal but not to be discarded. Deposition of the yellow dolomites would have been in shallow water, quite likely less than a metre, but periods of subaerial exposure would also be relatively frequent in this arid-zone location. There could well have been very shallow water, leading to subaerial exposure, after deposition of the red clastic material introduced by a flood, such that a dinosaur could then have walked across the newly deposited sediment layer.

The blue-grey marls are interpreted as dominantly lacustrine, for the most part low energy, at a time when only terrigenous clay was being provided to and being deposited within the basin, along with fine CaCO₃-dolomite being precipitated in the lake. Sedimentary features in the marl (Figures 11 a, b) do suggest occasional moderate to high-energy events (storms) causing erosion and disruption of the lake-floor sediments, reworking and some in-situ brecciation. Sediment deformation, disruption and brecciation can be explained as seismogenic, possibly linked to the adjacent major fault at Clevedon. However, it should be noted that Mao et al. (Reference Mao, Liu and Zhou2021) (and others, previously) proposed that the red MMG marls (and by extension the blue-grey marls) could be of aeolian dust origin, by comparison to the aeolian red clay on the Chinese Loess Plateau, in view of their grain size, micromorphology and geochemistry. Although there are many aspects of the MMG mudstones and marls which indicate the presence of a lake (e.g. Tucker, Reference Tucker, Matter and Tucker1978), at Clevedon and Ladye Bay the oolite and cross-bedding (etc.) clearly demonstrate there was a lake here. The change from the typical red marls of the MMG to the blue-grey marls in the upper part (Blue Anchor Fm.) relates to increased humidity heralding the onset of marine conditions in the Rhaetic-Liassic.

The small outcrops of blue-grey marl near Clevedon Pier show all the features described from the ‘classic’ exposures of the Blue Anchor Fm. on the Glamorgan coast at Lavernock Point, Penarth and in Somerset at St Audrie’s Bay (Mayall, Reference Mayall1981; Waters & Lawrence, Reference Waters and Lawrence1987). This would indicate a younger age for the Clevedon yellow dolomites and Ladye Bay oolites than that suggested by Milroy & Wright (Reference Milroy and Wright2000), although there is no biostratigraphic evidence available to substantiate this.

8.b. Early diagenesis

8.c. Interpretations of stable isotope data

Values of the oolites and blue-grey marls are relatively close, being low positive in both δ¹³C _PDB and δ¹⁸O _PDB (Figure 19), and these are typical of lacustrine carbonates, including dolomites, precipitated in alkaline saline lakes, such as those in East Africa (see Della Porta, Reference Della Porta2015 and Guo et al. Reference Guo, Wen, Li, He and Sanchez-Roman2023 for reviews, and Zhao et al. Reference Zhao, Wei, Han, Han, Gao, Meng, Wang, Tucker, Minghui and Sánchez-Román2023 for dolomite from Qaidam lake, Tibet). The isotope data are also similar to alkaline lake carbonates of the Pre-salt (Cretaceous), Santos Basin, Brazil (Wright & Barnett, Reference Wright and Barnett2017; Gomes et al. Reference Gomes, Bunevich, Sartorato, Tedeschi, Tonietto, Tucker and Whitaker2024). The low positive δ¹³C _PDB values of ooids and blue-grey marl do not indicate any significant extreme microbial or meteoric water influence on carbonate precipitation; rather, they reflect an overall equilibrium between dissolved inorganic carbon in the lake water and atmospheric CO₂ (Leng & Marshall, Reference Leng and Marshall2004). The positive δ¹⁸O _PDB signatures are consistent with significant evaporation, driving off the lighter ¹⁶O. In contrast, Triassic calcretes and travertine-tufa from Glamorgan have low to moderately negative signatures in both δ¹³C _PDB and δ¹⁸O _PDB. These are typical values for non-marine, pedogenic – spring deposits with fluids of meteoric origin and some microbial influence (Andrews et al. Reference Andrews, Riding and Dennis1993; Andrews, Reference Andrews2006).

8.d. Silicification

One hypothesis for the formation of the red silicified bands and nodules is by ingress of meteoric water, plausibly introduced by flash floods into the lacustrine environment. The water transporting coarse terrigenous sediment (hyperpycnal flows) and that with suspended red clay (hypopycnal flows) would have had a markedly different composition from that of the lake and sediment pore waters. As discussed above, from the mineralogy of the lacustrine carbonates and clays and the local presence of evaporites in the succession, alkaline, saline waters with a relatively high Mg content and moderately high pH are inferred, as has been documented for the Mercia Mudstone Group sediments generally (Jeans, Reference Jeans1978; Talbot et al. Reference Talbot, Holm, Williams and MR1994; Wright & Sandler, Reference Wright and Sandler1994; Milroy & Wright, Reference Milroy and Wright2000). Such alkaline water is expected to contain significant amounts of silica in solution, in a hydrolyzed form, Si(OH)₄, in view of the increased solubility of silica at high pH (Bustillo, Reference Bustillo, Alonso-Zarza and Tanner2010; Tucker, Reference Tucker, Tucker and Jones2023a). The storm water entering the shallow lake would be fresh-meteoric, with low salinity and near-neutral pH. We infer that the lake was supplied with low-silica water and that the silica concentration increased in the lake through a combination of evaporation and increased pH governed by the biology and chemistry as described earlier, as well as dissolution of clays imported in suspension or in wind-blown dust. Silica contents could reach a high enough concentration such that a colloid or gel could form. The presence of very high amounts of silica in gel form could also account for sinking of the domains that would eventually become silicified rock into the underlying sediment and their ductile character. Silicification might then be triggered when the Si content reached a threshold governed by the meteoric water influx with lowering of pH and evaporation. Suitable conditions for triggering might be expected to develop following prolonged drought and then a sudden influx of low-pH freshwater.

That the silicified beds are red and the sediment was prone to soft-sediment deformation indicates that the silicification is a very early event. As noted above, the yellow carbonate sediment became anoxic in the shallow subsurface and pyrite was precipitated. In modern shallow-marine-lacustrine environments, this takes place from a few mm to 10s of cm below the sea/lake floor as a result of the decomposition of organic matter and consumption of oxygen. Fe-oxide particles (red dust, clay) deposited out of suspension on the lake floor would have been dissolved in anoxic porewaters on shallow burial and then re-precipitated as pyrite, if not preserved within the chert/silica gel before this. Thus, the preservation of the red colour of the silicified layers and nodules is significant, indicating very early precipitation of silica to form a tight, low porosity-permeability rock with the Fe-oxide particles protected from reduction.

The silicification is in the form of precipitation of quartz crystals, as a micro- to mega- quartz cement within pore spaces within the sediment, as well as within the voids once occupied by ooids; some replacement of the fine-grained lime mud matrix in the red sediment could also have occurred. With the original aragonite mineralogy of the ooids, the influx of meteoric water would have favoured their dissolution. The partial fracture of the outer lamellae of ooids (Figures 15 a–d) and then quartz cementation, also indicate extremely early silica precipitation. The growth of the micro- to mega- quartz crystals could have taken place within an initial precipitate of silica gel.

Silicification of lacustrine carbonates has been reported from several other formations, notably the Tertiary of the Madrid Basin, Spain (e.g., Bustillo et al. Reference Bustillo, Arribas and Bustillo2002), the Aptian Pre-salt of the South Atlantic Basins, offshore Brazil and Angola (Wright & Barnett Reference Wright and Barnett2015; Pereira et al. Reference Pereira, Terra, da Silva, Pires and Ribeiro2023; Gomes et al. Reference Gomes, Bunevich, Sartorato, Tedeschi, Tonietto, Tucker and Whitaker2024) and the Permian Fengcheng Formation, China (Yu et al. Reference Yu, Du, Cao and Zhang2024). With the Pre-salt, a huge alkaline lake developed during the initial opening of the South Atlantic, freshening of lake waters during phases of a more humid climate is thought to have triggered silicification of carbonates through decreasing alkalinity and pH; however, some silicification is related to hydrothermal fluids (Teboul et al. Reference Teboul, Durlet, Girard, Dubois, San Miguel, Virgone, Gaucher and Camoin2019; Lima et al. Reference Lima, Tedeschi, Pestilho, Santos, Vazquez, Guzzo and De Ros2020), as is that in the Fengcheng Fm. (Yu et al. Reference Yu, Du, Cao and Zhang2024).

The initial precipitation of the hydrolyzed silica in solution in the lake water through freshening and reduction of pH could have been in the form of a precursor amorphous silica or a silica gel. This process has been described from some alkaline-saline East African lakes, where gels of Na and Mg silicates are also forming, with the silicon and other ions leached from volcanic rocks in the hinterland (Eugster & Jones, Reference Eugster and Jones1968; Renaut et al. Reference Renaut, Jones, Tiercilin and Tarits2002; Renaut et al. Reference Renaut, Owen, Lowenstein, de Cort, McNulty, Scott and Mbuthia2021; Raudsepp et al. Reference Raudsepp, Wilson and Morgan2023). Siliceous precipitates here include magadiite, a metastable hydrated sodium silicate, which converts to chert in a relatively short time (Owen et al. Reference Owen, Renaut, Muiruri, Rabideaux, Lowenstein, McNulty, Leet, Deocampo, Luo, Deinoh, Cohen, Sierj, Campisanol, Shenm, Billingsleyi, Mbuthiao and Stockheckep2019). The formation of gels as a precursor to silica has been invoked for the Aptian Pre-salt of the South Atlantic basins (Wright, Reference Wright2022; Wright & Barnett, Reference Wright and Barnett2015), with precipitation, again, taking place when pH is lowered, as through freshwater input. Evaporation then leads to Mg-silicate formation, notably stevensite and kerolite. An analogous process producing gelatinous opaline chert is occurring in several lakes in South Australia (Peterson & von der Borch, Reference Peterson and von der Borch1965; Ambrose & Flinn, Reference Ambrose and Flinn1981), attributed to the mixing of ground water and lake water during falling lake level, and/ or seasonal pH changes.

Silicified sediments and soils forming silcretes develop through pedogenesis in the shallow subsurface of soil profiles and somewhat deeper at the groundwater table (Ambrose & Flinn, Reference Ambrose and Flinn1981; Ullyott & Nash, Reference Ullyott and Nash2016; Taylor & Eggleton, Reference Taylor and Eggleton2017). These are mostly composed of microquartz, but with coarser fabrics in vugs and there are commonly relics of plant roots present in the form of canals and tubes. These deposits can be a metre or more thick and may be autobrecciated. In the case of the Triassic red silicified sediments here, generally thin layers and discrete nodules of chert within ‘normally’-deposited carbonates, there are no textures suggesting pedogenesis, and no evidence of the former presence of rootlets. Chalcedonic quartz, typical of hydrothermal silica (Lima et al. Reference Lima, Tedeschi, Pestilho, Santos, Vazquez, Guzzo and De Ros2020), is also absent. However, a key observation which could support a precursor high silica concentration gel as an intermediate step to complete silicification is the ductile behaviour of the domains that eventually became hard silicified rock. These bands and layers have plastic shapes and became involved in soft-sediment deformation. In fact, early syn-sedimentary deformation is common in silicified sediments generally (e.g. Gimenez-Monsant et al. Reference Gimenez-Montsant, Calvet and Tucker1999, Eocene of Spain), including the Pre-salt of Brazil (Pereira et al. Reference Pereira, Terra, da Silva, Pires and Ribeiro2023). Such features could reflect the effects of density contrasts between a gel of hydrolyzed silica, micro- mega- quartz, and the host dolomite mud.

Although we conclude that silicification is very early and related to syn-sedimentary pore-fluid mixing, there are records of extensive silicification in the Mendips area within the Dolomitic Conglomerate near East Harptree and also in the Lower to Middle Jurassic coarse bioclastic limestones (Brockley Down Limestone) on the north side of the Mendip Hills (Simms, Reference Simms1995; King, Strategic Stone Study, Reference King2011). This silicification is more pervasive stratigraphically and is likely related to the hydrothermal fluids involved in the Mendip Pb–Zn mineralization, thought to be of Jurassic age (Simms, Reference Simms1995), and contrasts with the more bed-bound localized silicification of the red chert bands and nodules at Clevedon and their evidence for a very early diagenetic origin.

The silicified nodules present within the yellow dolomites lower in the succession at Clevedon (Figure 10 a, b) are similar to nodules noted by Milroy (Reference Milroy1998) in his units 1 and 2 at Ladye Bay. The overall shape and ‘cauliform’ external surface suggest they are replaced anhydrite nodules, which may well have formed from replacement of cm-scale gypsum crystals precipitated within the sediment. Such sulphates are common in arid-zone littoral lacustrine and high inter-supra-tidal marine facies and are typical of sabkha facies generally. Similar facies have been described from the Marginal Triassic of Glamorgan by Tucker (Reference Tucker, Matter and Tucker1978) and Leslie et al. (Reference Leslie, Spiro and Tucker1993). These nodules, also referred to as ‘potato stones’, are similar in external shape to Bristol diamonds (Tucker, Reference Tucker1976), with some also having a central cavity containing large megaquartz crystals, the feature of the ‘diamond’ geodes. The ‘full-bodied’ shape/uncompacted nature of the nodules indicates an early timing for their silicification and this probably occurred through fresh-water ingress in a similar manner to the formation of the red silicified bands and nodules higher in the succession.

8.e. Later diagenesis: burial and uplift

The Triassic sediments were buried to a depth probably in excess of 800 m until regional uplift began in the early Cenozoic. During this time, fracturing and faulting of the Triassic sediments took place in response to regional stresses related to opening of the Central Atlantic and formation of Mesozoic basins, such as the Wessex Basin to the southeast, with movement on reactivated Variscan and younger faults (Cosgrove et al. Reference Cosgrove, Morgan and Ghail2021). Apart from some pressure dissolution planes, stylolites and calcite veins, the main post-Triassic event was the passage of hydrothermal fluids up the Clevedon Pier fault and associated fractures and consequent precipitation of Pb-Cu sulphides and baryte there in the fault breccia and adjacent porous yellow dolomite (Ixer et al. Reference Ixer, Pattrick and Starkey1993). Galena (mm-cm scale patches) is conspicuous in fractures within the yellow dolomite at Clevedon, as is baryte, notably within the Dolomitic Conglomerate and fault breccia, but it is also disseminated as sub-mm particles within the yellow dolomite, as seen in SEM-EDS maps (Figure 18 a).

The yellow colour of the Triassic dolomite beds is attributed to the oxidation of disseminated pyrite and replacement by goethite-limonite. This could have taken place during moderate to deep burial if the porewaters were oxic, as is the case with many deeper groundwaters as a result of the extensive microbiota occurring in the subsurface and the generation of dark oxygen there (Ruff et al. Reference Ruff2023). However, oxidation of pyrite would also have taken place much later during Cenozoic uplift on interaction with surface-derived oxic groundwaters. In fact, many carbonate rocks in the UK, from the Permian, Triassic and Jurassic, are a buff-yellow-honey colour at outcrop, but a dark grey colour, even blue, in the shallow subsurface as a result of the presence of original organic matter and pyrite (Tucker, Reference Tucker2023c). During this late phase of exhumation and weathering of the Clevedon yellow dolomites in the last few million to few 10s of 1000s years, it is likely that the black Mn-oxide particles and dendrites also formed. Mn would be released from organic matter, clays and pyrite in the highly oxic near-surface groundwaters. The oxidation (Liesegang) rings would also have formed at this time from episodic migration of oxic chemical fronts (Ashley, Reference Ashley2014).