Precambrian Research 145 (2006) 243–271

Depositional environments at a Palaeoproterozoic continentalmargin, Vastervik Basin, SE Sweden

Lena Sultan a,b,∗, Piret Plink-Bjorklund a

a Department of Geology, Earth Sciences Centre, Goteborg University, Box 460, SE-40530 Gothenburg, Swedenb Laboratory for Isotope Geology, Swedish Museum of Natural History, Box 50007, SE-10405 Stockholm, Sweden

Received 7 March 2005; received in revised form 13 December 2005; accepted 20 December 2005

Abstract

This study documents: (1) the sedimentary environments and depositional processes and (2) the temporal and spatial relation-ships of sedimentary environments within well-preserved areas in the Palaeoproterozoic Vastervik metamorphic siliciclastics insoutheastern Sweden.

Based on lateral and vertical facies transitions, seven facies associations were documented: fluvial deposits (Facies Association1), tidal flat deposits (FA 2), tidal channel deposits (FA 3), upper-flow-regime tidal flat deposits (FA 4), tidal bar deposits (FA 5),shoreface deposits (FA 6) and turbidite deposits (FA 7). The documented facies associations reflect deposition on a Palaeoproterozoiccontinental margin, where rivers drained the continent to the present north/northwest and carried sediment southeastwards. Therivers became tidally influenced further seawards. Sediments accumulated in tide-dominated estuaries, river-dominated deltas andin wave-dominated shallow-water environments. The documented tide-dominated estuary suggests that the coastline at least duringsome periods was macrotidal (>4 m).

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The metasediments of the Vastervik succession are locally strongly altered by tectonic deformation and metamorphism, andetailed vertical and lateral correlation was possible only in restricted separate areas. Vertical facies transitions show transgressiveo regressive deposition in two of the in detail studied areas as indicated by the vertical change from riverine to a shallow marinend back to river deposition in Area 1. In Area 3 the transition occurs from river-dominated deltas into tide-dominated estuaries,ucceeded above by river deposition. In the two remaining areas the deposition was overall transgressive as river deposits areucceeded above by deltaic turbidites.

Despite the somewhat “fragmentized” character of the Vastervik succession this work enabled to recognize general character ofhe Vastervik Palaeoproterozoic continental margin.

2006 Elsevier B.V. All rights reserved.

eywords: Palaeoproterozoic; Tidal deposits; Tide-dominated estuary; Transgressive; Regressive

. Introduction

This paper reports the first detailed and systematicedimentologic outcrop study of the Svecofennian

∗ Corresponding author. Tel.: +46 31 773 28 00;ax: +46 31 773 28 49.

E-mail address: lena@gvc.gu.se (L. Sultan).

Vastervik metamorphic siliciclastics in the BalticShield (Fig. 1) demonstrating that it is possible to dosedimentological analyses in tectonic and metamorphicareas. The Palaeoproterozoic Vastervik Basin haslong been known for locally well-preserved primarysedimentary structures and geometries (Hogbom etal., 1910; Gavelin and Russell, 1967; Russell, 1967;Gavelin, 1984). Vertical and lateral facies transitionscan be documented within the restricted well-preserved

301-9268/$ – see front matter © 2006 Elsevier B.V. All rights reserved.doi:10.1016/j.precamres.2005.12.005

244 L. Sultan, P. Plink-Bjorklund / Precambrian Research 145 (2006) 243–271

Fig. 1. (A) Simplified geological map of the Baltic Shield; (B) simplified geological map of Vastervik area, modified after Gavelin (1984). Separatestudy areas are marked on the map; (C) general stratigraphy of Areas 1–4.

L. Sultan, P. Plink-Bjorklund / Precambrian Research 145 (2006) 243–271 245

areas only. The primary sedimentary structures andgeometries are well preserved even in places where theprimary textures have been significantly altered and thedetrital grains have been obviously recrystallized.

The objectives of this study are to document: (1) thesedimentary environments and depositional processesand (2) the temporal and spatial relationships of sedi-mentary environments within the well-preserved areas,as well as establish the depositional regime of thesePalaeoproterozoic metasediments.

Earlier studies have mainly focused on recognitionof sedimentary structures, tectonic deformation andmetamorphism in the metasedimentary rocks aroundVastervik (Hogbom et al., 1910; Beunk and Page, 2001).The Vastervik succession is suggested to have accu-mulated in a large delta with tidal flats adjacent to thedelta (Gavelin and Russell, 1967; Gavelin, 1984). Thesuggested net sediment transport was from the presentnorth. In this paper we aim to demonstrate that a muchmore complicated mosaic of depositional environmentsexisted in the Vastervik Basin, including episodes ofdeltaic progradation and a net seaward transport of sed-iments. However, episodes of transgressive estuarinedeposition occurred, with a net sediment transport intothe estuaries, and a powerful storage of sediments withinthe estuaries.

2. Geological setting

The Vastervik succession occurs on the southern-mbnBprhPPfacip1

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et al., 1996), and the BABEL-B seismic profile (e.g.Midgley and Blundell, 1997), Beunk and Page (2001)suggest that the Vastervik succession was deposited ina back-arc basin formed in response to a northwardsubduction beneath the Oskarshamn-Jonkoping Belt.The Oskarshamn-Jonkoping Belt was first recognizedas a terrane separate from the surrounding Svecofennianrocks by Mansfeld (1996).

The newly presented 1859 ± 9 Ma age for the Lof-tahammar (TIB) granite (Bergstrom et al., 2002) thatintrudes into the northern part of the Vastervik succes-sion suggests a constraint for the minimum age of thebasin. The youngest documented detrital zircon Pb/Pbages of 1837 ± 22, 1872 ± 24 and 1870 ± 12 Ma (Sultanet al., 2005) suggest that deposition ended not longbefore intrusion of the granite.

The preservation degree of the metasedimentary suc-cession varies between different areas in the Vastervikarea, commonly the sedimentary structures are best pre-served in the hinges of large-scale folds (Gavelin, 1984).Between the well-preserved areas sedimentary structuresare totally obliterated, either by deformation or by thevicinity to intruding granitoids.

3. Facies associations

Sedimentary facies were studied measuring detailedvertical sections, lateral mapping and ‘walking out’stratigraphic levels laterally or downdip. The docu-mented sedimentary facies are grouped into seven facies

ost edge of the Svecofennian domain (Fig. 1), formedy a large crust-forming event during the Svecofen-ian orogeny, ∼1.9–1.86 Ga ago (Gorbatschev andogdanova, 1993). The Svecofennian orogen is inter-reted to represent several events of accretion of ter-anes (Beunk and Page, 2001). Svecofennian rocksave been documented in Sweden, Finland and belowaleozoic cover in the Baltic countries (Beunk andage, 2001). In Sweden these rocks are dominantlyelsic volcanics and metasedimentary rocks. West ofnd within the Svecofennian domain runs the Transs-andinavian Igneous Belt (Fig. 1), a suite of bimodalntrusions and volcanics that were emplaced in severalulses between 1.85 and 1.65 Ga (Larson and Berglund,992).

The Vastervik metasedimentary succession is sur-ounded by intruding granitoids mainly belonging to theransscandinavian Igneous Belt (TIB). The surroundingIB rocks are mainly of 1850–1800 Ma age (Bergstromt al., 2002). Based on structural data, geochronologi-al data and geothermobarometry, in combination withreviously published geochemical evidence (e.g. Beunk

associations based on textures, sedimentary structures,geometry, palaeocurrent direction indicators, and lateraland vertical facies transitions. Detailed descriptions ofindividual sedimentary facies are given in Table 1 andFigs. 2–8. Facies groups 1.1–1.3, 2.1–2.9, 3.1–3.2 and4.1–4.3 refer to dominating deposition agent, i.e. uni-directional currents, bi-directional currents, waves andgravitational processes, respectively.

The detrital grains are obviously altered, and in placestotally recrystallized. In places, e.g. large (1–2 cm)cordierite, andalusite or sillimanite porphyroblasts, char-acteristic for metamorphic rocks (Kresten, 1971) occur.Structures discordant to bedding (e.g. ripple cross-lamination, cross-stratification, troughs and channels)are easy to identify as primary sedimentary struc-tures. Such primary sedimentary structures can beused to provide direct evidence for the primary grainsize in recrystallized rocks, e.g. cross-stratificationis hydrodynamically possible only in sand-size andcoarser material; swaley cross-stratification has onlybeen documented from very fine-grained sandstones,etc. (e.g. Harms, 1975). Larger clasts, like pebbles,

246 L. Sultan, P. Plink-Bjorklund / Precambrian Research 145 (2006) 243–271

Table 1Characteristics of sedimentary facies

Facies Description Interpretation Occurrence

1.1. Trough cross-stratifiedsandstone

Typically 15–30 cm thick beds, ranges from 5 to100 cm of well-sorted trough cross-stratifiedsandstone (Fig. 2). Sets extend laterally 4–5 m

Deposited by migration of 3Ddunes

FA: 1, occurs interbeddedwith Facies 1.2 and 1.3

1.2. Quartzite conglomerate 25–30 cm thick cross-stratified matrix-supportedconglomerate, with rounded quartzite pebbles andangular clay clasts (Fig. 2)

Deposited by traction currentsin high-energy environment

FA: 1, occurs interbeddedwith Facies 1.1

1.3. Plane-parallel laminatedmudstone

5–7 cm thick plane-parallel laminated beds,flat-based, sheet-like (Fig. 2)

Deposited by suspensionsettling

FA: 1, occurs interbeddedwith Facies 1.1

2.1. Plane-parallel laminatedmudstone with polygonaland/or spindle-shapedcracks

5–10 cm thick flat-based and sheet-like beds with3–10 cm deep sandstone-filled spindle-shaped (Figs.3C and 5) and polygonal cracks (Fig. 3B). Bedsextend laterally for tens or hundreds of meters in FA2 but only 2–5 m in FA 5

Deposited by suspensionsettling. Spindle-shapedcracks interpreted assynaeresis cracks. Polygonalcracks formed duringsubaerial exposure

FA: 2, 5, occursassociated with Facies2.2, 2.3, 2.4, 2.6 and 2.9

2.2. Plane-parallel laminatedmudstone and sandstone

5–25 cm thick flat-based and sheet-like beds. Acontinuum of deposits from dark gray plane-parallellaminated mudstone with thin sandstone interlayersto light sandstone with thin mudstone interlayers(Figs. 3 and 5)

Alternating deposition fromsuspension and from tractioncurrents, deposited by tidalcurrents

FA: 2, 5, occursassociated with Facies2.1, 2.3, 2.4, 2.5, 2.6, 2.7and 2.9

2.3. Lenticular bedding ∼10 cm thick flat-based sheet-like beds of pinksandstone with unidirectional sandstone ripplesseparated by dark gray mudstone flasers (Fig. 3).Beds extend laterally for tens of meters

Intertidal deposits whereripples form by tractioncurrents and mudstone drapesform during slack-waterperiods by suspension settling

FA: 2, associated withFacies 2.1, 2.2 and 2.6

2.4. Bi-directionalripple-laminated sandstonewith mudstone drapes

5–15 cm thick, bi-directional ripple-laminated setswith mudstone drapes between individual laminae(Figs. 3 and 5). Beds appear very disturbed and bedcontacts are undulating and erosional. Reactivationsurfaces occur within sets

Deposited as migratingripples from reversing tidalcurrents

FA: 2, 5, associated withFacies 2.1, 2.2, 2.5, 2.6and 2.9

2.5. Compoundcross-stratified sandstone

2–25 cm thick cross-stratified sets with low-anglemaster bedding surfaces with superimposed ripplesclimbing up or down

Deposited from tidal currents,where the master beddingplanes form by the dominantcurrent and the superimposedripples are deposited from thesubordinate current

FA: 5, associated withFacies 2.2, 2.4, 2.6 and2.9

2.6. Bi-directionalcross-stratified sandstonewith mudstone drapes

5–50 cm thick sets of cross-stratified bi-directionalsandstone with erosional contacts (Fig. 5).Mudstone drapes, reactivation surfaces within sets,clay chips and rip-up clasts occur in places

Deposited from tidal currents FA: 2, 3, 5, associatedwith Facies 2.1, 2.2, 2.3,2.4, 2.5, 2.7 and 2.8

2.7. Large-scale bi-directionalcross-stratified sandstone

60–500 cm thick sets of cross-stratifiedbi-directional sandstone (Fig. 4), in places withmudstone drapes along laminae. Sets typically fineupwards, and occur above significant erosionsurfaces

Deposited from tidal currents FA: 3, associated withFacies 2.6 and 2.8

2.8. Plane-parallel laminatedsandstone

25–75 cm (in FA 4) or 100–125 cm (in FA 3) thickbeds of plane-parallel laminated sandstone

Deposited fromupper-flow-regime currents

FA: 3, 4, associated withFacies 2.6, 2.7 and inplaces 2.1

2.9. Thick mudstone beds 50–90 cm thick beds of mudstone. Large-scalewater-escape structures occur at bases of beds

Rapid deposition from fluidmuds

FA: 5, occurs on top of analternation of Facies 2.1,2.2, 2.4, 2.5 and 2.6

3.1. Alternatingsymmetrical-ripple andplane-parallel laminatedsandstones

5–25 cm thick sets of alternating ripples andplane-parallel laminated sandstones. Mudstonedrapes occur within ripples in places. Beds areflat-based and lenticular (Fig. 6)

Deposition by traction inoscillatory flow conditions

FA: 6, interbedded withFacies 3.2

3.2. Swaley cross-stratifiedsandstone

5–40 cm thick flat-based beds of low-angle swaleystratification (Fig. 6). Bed contacts are nearly flatand beds are lenticular

Deposited by high-energywave action

FA: 6, interbedded withFacies 3.1

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Table 1 (Continued )

Facies Description Interpretation Occurrence

4.1. Upward-fining sandstoneto mudstone beds

Typically 2–5 cm thick sets, in places up to 20 cmthick of plane-parallel laminated sandstone that finesupwards into dark gray plane-parallel laminated orripple cross-laminated sandy mudstone to mudstone.Bed contacts diffuse and flat, beds are sheet-like(Fig. 7). Load casts are common at bases of beds.Bed thickness is typically 2–5 cm; in places up to15–20 cm. The thicker beds have occasionallyungraded or planar tabular cross-stratified sandstoneintervals at the bases of beds. These thicker bedscommonly display abundant water-escape structures(Fig. 7)

Waning, surge-type turbiditycurrents

FA: 7

4.2. Thick upward-finingsandstone to mudstonebeds

Intervals of upward-fining beds of sandstone tomudstone (Fig. 8). Beds typically consist ofintervals of ungraded sandstone, commonlydeformed by ball and pillow structures, load castsand water-escape structures, followed by planartabular cross-stratified sandstone, plane-parallellaminated sandstone, asymmetrical-ripplecross-laminated sandstone and finally plane-parallellaminated sandy mudstone or mudstone. Beds arecommonly incomplete, lacking one or several of theabove described structures. Beds have flat bases anda sheet-like geometry, and are 10–110 cm thick

Waning, surge-type turbiditycurrents

FA: 7, interbedded withFacies 4.3

4.3. Soft-sediment deformedsandstone beds

15–25 cm thick soft-sediment deformed beds(Facies 4.3) occur. Lamination within beds is oftenpreserved, but is commonly deformed by folding. Inplaces beds are totally deformed and show nopreserved lamination (Fig. 8)

Soft-sediment slump bedsthat form on slopes

FA: 7, interbedded withFacies 4.2

are clearly recognizable. The prefix “meta” has beenomitted from “sandstone”, “mudstone”, etc. for easierreadability.

3.1. Facies Association 1: fluvial deposits

Facies Association 1 consists of erosionally basedsandstone bodies, 3–12 m thick. The facies associationis dominated by light gray to reddish, well-sorted troughcross-stratified sandstone of Facies 1.1, in places basedby a cross-stratified, matrix-supported rounded quartzpebble and angular clay-clast conglomerate of Facies1.2, and capped by plane-parallel laminated mudstoneof Facies 1.3 (see Table 1 and Fig. 2).

The cross-sets of sandstones in Facies 1.1 are typ-ically 15–30 cm thick. Set boundaries are commonlyparallel, with the troughs cutting into the underlying sets.Sets commonly extend 4–5 m laterally, and are erosion-ally based. Lee-faces of dunes dip systematically towards154–174◦ SE, with a few cross-sets in opposed directiontowards 335–350◦ NW, in Area 1, towards 218–226◦ SWin Area 3 and towards 135–139◦ SE in Area 4. The cross-

stratified sandstones in many places contain ripped-upclasts of mudstones, 2–10 cm in diameter. The clasts arealigned on erosion surfaces. In many places the lee-facesof the cross-strata have systematically been erosion-ally modified, and have a characteristic gently dippingconvex-up shape (Fig. 2). In a few places mm-thick mud-stone layers drape these surfaces. The conglomerate bedsof Facies 1.2 are commonly 15–20 cm thick. The mud-stone of Facies 1.3 occurs in flat-based and sheet-likebeds.

Soft-sediment deformation, especially overturnedcross-stratification, is common in Facies Association 1.The degree of overturning of cross-strata varies, in placesit is only the topmost centimeters and in places almostthe whole set. The degree of deformation characteristi-cally increases vertically towards the top of the sets. Insome places, soft-sediment deformation occurs as foldsin the lower parts of the sets. Load casts occur in placesat bases of cross-sets.

Facies Association 1 occurs in northwestern reachesof Area 1, and is interbedded with gravitational depositsof Facies Association 7 in Areas 3 and 4.

248 L. Sultan, P. Plink-Bjorklund / Precambrian Research 145 (2006) 243–271

Fig. 2. Tide-influenced fluvial deposits (Facies Association 1). Rose diagram shows dominant palaeocurrent direction towards present S–SE, this isan example from Area 1. In Area 3 palaeocurrent directions are towards SW and in Area 4 towards SE. (A) Rip-up clast within trough cross-stratifiedsandstone (Facies 1.1) and (B) typical appearance of overturned trough cross-stratified sandstone (Facies 1.1). Pen for scale.

3.1.1. InterpretationThe extensive unidirectional (SE) cross-stratification

in conglomerates and sandstones, together with basalerosion surfaces, and lack of wave influence suggestdeposition in fluvial channels. The trough cross-stratifiedconglomerates and sandstones (Facies 1.1 and 1.2) weredeposited in 3D dunes and in longitudinal bars from trac-tion currents (Collinson, 1996). Plane-parallel laminatedmudstones of Facies 1.3 are a characteristic productof suspension fallout, and suggest deposition in fluvial

overbank environment or indicate channel abandonment(Collinson, 1996; Miall, 1996). Lack of lateral accretionbeds, occurrence of conglomerate in the channel fills,low abundance of overbank deposits and relatively lowpalaeocurrent variability suggest that the channels hadrelatively low sinuosity (Bridge, 2000).

The reactivation surfaces on steeply dipping cross-strata indicate that a bedform lee-side is changed intostoss-side by reversals in flow direction (e.g. Boersma,1969; Visser, 1980; Boersma and Terwindt, 1981; Allen

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Fig. 3. Tidal deposits (Facies Association 2). (A) Lenticular bedding (Facies 2.3); (B) sand-filled desiccation cracks suggesting subaerial exposure attimes; (C) typical appearance of muddy units (of Facies 2.1 and 2.3) with spindle-shaped synaeresis cracks and ripples; (D) double mud drapes withinbi-directional ripple-laminated sandstone (Facies 2.4), lower picture with mud drapes drawn on photograph. Rose diagram shows palaeocurrentdirections towards SW and NE. These tidal facies have been interpreted as tidal flat deposits (see text for interpretation).

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Fig. 4. Tidal channels (Facies Association 3). Rose diagram shows palaeocurrent directions towards SW and NE. (A) Large-scale cross-stratifiedsandstone (Facies 2.7) and (B) tidal channel fill cuts into tidal flat deposits. Channel bottom can be traced downcurrent for hundreds of meters.

and Homewood, 1984; Shanley et al., 1992), or migrationof superimposed bedforms (Dalrymple, 1984; Shanleyet al., 1992). The systematic occurrence of reactivationsurfaces together with the occasional mud drapes mightsuggest that the fluvial channels were influenced by tidalcurrents (see Allen, 1980).

The load casts are a result of temporary loss ofstrength shortly after deposition, commonly due to rapiddeposition. Overturned cross-stratification is caused bythe drag of the overriding current (Allen and Banks,

1972; Owen, 1996). The overturned cross-stratificationis suggested to be characteristic for fluvial environments(Allen and Banks, 1972). Deformation in the lower partsof beds suggests liquefaction-induced gravitational col-lapse (Owen, 1996).

3.2. Facies Association 2: tidal flat deposits

Facies Association 2 consists of packages of bi-directional ripple-laminated sandstones with ubiquitous

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Fig. 5. Tidal bars (Facies Association 5). (A) Typical appearance of tidal bars with bi-directional cross-stratified sandstone (Facies 2.6) withwater-escape structures and erosional contacts; (B) bi-directional cross-stratified sandstone (Facies 2.6) and plane-parallel laminated mudstonewith synaeresis cracks (Facies 2.1); (C) bi-directional ripple-laminated sandstone (Facies 2.4) eroded by the overlying set. Rose diagram showspalaeocurrent directions towards NE and SW.

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Fig. 6. Shoreface deposits (Facies Association 6) with alternating symmetrical-ripple-laminated and plane-parallel laminated sandstone (Facies 3.1)and swaley cross-stratified sandstone (Facies 3.2). Rose diagram shows palaeocurrent directions towards NW. (A) Symmetrical ripples.

mudstone drapes (Facies 2.4 and 2.3), plane-parallellaminated sandstone and mudstone beds (Facies 2.2),plane-parallel laminated mudstone beds with sandstone-filled cracks (Facies 2.1) and in places bi-directionalcross-stratified sandstones with mudstone drapes (Facies2.6) (see Table 1 and Fig. 3). The plane-parallel lam-inated sandstones and mudstones (Facies 2.1 and 2.2)volumetrically dominate Facies Association 2.

Facies 2.1 consists of dark gray plane-parallel lami-nated flat-based and sheet-like mudstone with abundantpink sandstone filled cracks, 3–10 cm deep (Table 1and Fig. 3). Spindle-shaped cracks as well as polygo-nal cracks are common. Beds are 15–80 cm thick andextend laterally for tens or hundreds of meters. Thespindle-shaped cracks occur as narrow, slit-like features,typically 0.7–1 cm wide and 8–10 cm long. They typ-ically display a slight bending or branching near thebase (Fig. 3C). The polygonal cracks are v-shaped andunfilled from above. They are typically 1.5–2 cm wideand 7–10 cm deep. Load casts occur in places. Charac-teristically the overlying sand layer is thin and does notdisplay loading structures (Fig. 3B). Load casts occur inplaces.

Facies 2.2 varies from dark gray plane-parallel lam-inated mudstone with thin sandstone interlayers to lightplane-parallel laminated sandstone with thin mudstone

interlayers (Fig. 3). Beds are sheet-like and laminaethickness varies typically between 2 and 50 mm, com-monly the sandier parts are thicker and the muddier partsthinner. Individual beds are typically 5–25 cm thick.Beds typically extend laterally for tens of meters. Loadcasts occur in places.

The plane-parallel laminated beds of Facies 2.1and 2.2 rhythmically alternate in the vertical section(Fig. 3C), and occur as 50–100 cm thick packages, inter-bedded by 20–30 cm thick erosionally based packagesof light gray cross-stratified sandstones of Facies 2.6.

The lenticular and bi-directional ripple-laminatedsandstones and mudstones (Facies 2.3 and 2.4) and lessabundant, and occur as 5–10 cm thick erosionally basedbeds. Facies 2.3 consists of pink sandstone with unidi-rectional sandstone ripples separated by dark gray muddrapes (Fig. 3). Facies 2.4 is a light gray to pink bi-directional ripple-laminated sandstone, with mudstonedrapes along laminae, within the ripples. Beds extendlaterally 0.2–2 m. Lee-faces of ripples commonly facetowards SW and NE.

Facies 2.6 occurs in a few places and consists ofpink cross-stratified sandstone with opposite dip of lee-faces in adjacent sets. The cross-sets of Facies 2.6 aretypically 5–15 cm thick in Facies Association 2. Lee-faces of cross-strata show convex-up erosion surfaces

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Fig. 7. Surge-type turbidites (Facies 4.1, Facies Association 7). (A) Diffuse laminated sandstone and mudstone and (B) typical appearance of Facies4.1. Although the muddier deposits have been tectonically deformed at a higher degree than the sandy deposits, the syn-depositional soft-sedimentdeformation can be recognized as the loading and water escape occur locally, at specific bed boundaries or within individual beds, whereas thetectonic deformation folds thick stratigraphic packages or larger areas.

and are draped by mudstone laminae, 1–3 mm thick.Sets extend laterally for 0.5–5 m. Palaeocurrent measure-ments derived from lee-faces of dunes show a dominantcurrent towards 250–265◦ SW and a subordinate currenttowards 70–78◦ NE.

Facies Association 2 occurs in Area 3 interbeddedwith tidal channel deposits of Facies Association 3,upper-flow-regime deposits of Facies Association 4 andtidal bar deposits of Facies Association 5.

3.2.1. InterpretationThe rhythmic alternation of sheet-like mudstone and

sandstone beds (see Fig. 3C), together with bipolarrippled sandstones with ubiquitous mudstone drapesand bimodal cross-strata with double mud drapes, sug-gests deposition by tidal currents (e.g. Reineck and

Wunderlich, 1968; Reineck and Sing, 1980; Visser,1980; Nio and Yang, 1991).

The bipolar cross-strata (Facies 2.6) indicate migra-tion of dunes in reversing currents of approximatelyequal strength (Allen, 1980), even though several tidalcycles may pass between deposition of individual sets.The double mudstone drapes in cross-strata indicate muddeposition on lee-faces of dunes during both slack-waterperiods (Visser, 1980).

The bipolar rippled sandstone beds with ubiquitousmudstone drapes between individual laminae (Facies2.4) is interpreted as deposited by reversing tidal cur-rents. Lenticular bedding (Facies 2.3) typically indi-cates deposition from reversing tidal currents (Reineckand Sing, 1980) and represents deposition in an inter-tidal environment (Dalrymple, 1992). Rippled sands are

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Fig. 8. Thick-bedded surge-type turbidites (Facies 4.2, Facies Association 7). (A) Rhythmical alternation of sandy to muddy turbidite beds andsoft-sediment deformed beds. Water-escape structures indicate rapid deposition. (B) Close-up shows cordierite-rich mud-prone beds and a cross-stratified set with few and small cordierites along laminae. Cordierites have formed during metamorphism. (C) Rhythmical alternation of sandstoneand mudstone. Arrows point to layer where strong water-escape has removed much sand. (D) Close-up on sandstone from picture C, showingpreserved internal lamination.

deposited during maximum tidal flow, whereas muddrapes form during ensuing slack-water periods. Whenthe subordinate current is strong enough, a second rip-pled unit with an opposite palaeocurrent direction willbe deposited on top of the first mud drape, followedby another mud drape formed by another slack-waterperiod.

When the tidal currents are too slow to produce rip-ples, they produce sand layers from suspension thatalternate with mud layers (Dalrymple, 1992), as seenin plane-parallel laminated sandstones and mudstones ofFacies 2.1 and 2.2. The spindle-shaped synaeresis cracksof Facies 2.1 are commonly interpreted to form sub-aqueously at the water-sediment interface due to salinitychanges (Plummer and Gostin, 1981). The polygonalcracks are interpreted to represent desiccation cracksformed during subaerial exposure (Plummer and Gostin,1981).

Alternatively, Plummer and Gostin (1981) suggestthat the spindle-shaped cracks are often formed sub-stratally due to sedimentational compaction inducedby deposition of overlying sediments. Some workerseven suggest that synaeresis cracks do not form at thesediment–water interface in natural systems (Tanner,1998), but form from interstratal cracking caused bylayer-parallel contraction. The bending near the baseof the cracks is then probably a result of compaction(Kidder, 1990). This alternative explanation is unlikely,as loading or other compaction and soft-sediment defor-mation features are not associated with the spindle-shaped cracks in the Vastervik succession.

The transition from bipolar cross-stratified sand-stones of Facies 2.6, to bipolar ripple-laminated sand-stones of Facies 2.4, to lenticular bedded sandstonesand mudstones of Facies 2.3, and into laminated mud-stones with sandstone layers and synaeresis and desic-

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cation cracks of Facies 2.2 and 2.1, indicates decreas-ing ebb and flood current speeds, and deposition intidal flat environments (e.g. Dalrymple, 1992; Readingand Collinson, 1996). The bimodal cross-stratified sand-stones and lenticular bedded sandstones with mud drapescommonly occur on subtidal to intertidal sand flats(Reineck and Sing, 1980; Dalrymple, 1992). Lenticu-lar beds and laminated mudstones with sandstone layersaccumulate on intertidal mixed to mud flats, where thecurrent energy is lower, and subaerial exposure frequent.

3.3. Facies Association 3: tidal channel deposits

Facies Association 3 occurs as erosionally basedsandstone bodies, 5–12 m thick. Large-scale bi-directional cross-stratified sandstone with mudstonedrapes (Facies 2.7) dominates and occurs as 5–7 m thickunits interbedded with 1–2 m thick units of bi-directionalcross-stratified sandstone (Facies 2.6) or plane-parallellaminated sandstone (Facies 2.8) (see Table 1 and Fig. 4).In some places Facies 2.6 and Facies 2.8 occur interbed-ded as 2–4 m thick units.

Facies 2.7 consists of white or pink well-sorted sand-stone with large-scale bi-directional cross-stratification(Fig. 4). The large cross-sets fine upwards and are cov-ered with mudstone in places. The sets, 0.6–5 m thick,extend more than 10 m laterally.

Facies 2.6 consists of pink cross-stratified sandstonewith opposite dips of lee-faces in superimposed sets. Setsof Facies 2.6 are thicker than in Facies Association 2,td

p1

wflb

3

tsg

witf1t

inated sandstone (Facies 2.8) contains no mud alonglaminae and is interpreted to be deposited from upper-flow-regime currents (Harms, 1975). The depositionalunits are based by significant erosion surfaces with rip-up clasts, interpreted to represent channel bases. Theseerosion surfaces are traceable for approximately 100 malong channel axis. The upward-fining large-scale cross-stratified sets (Facies 2.7) are interpreted to representbars in the tidal channel since the upwards fining reflectsupwards decrease in channel energy, as opposed to dunesthat reflect highest current energy at the top (Miall,1996). This is confirmed by the great height of thesecross-strata.

3.4. Facies Association 4: upper-flow-regime tidalflat deposits

Facies Association 4 consists of white well-sortedplane-parallel laminated sandstones of Facies 2.8, andoccasional cross-stratified sandstones of Facies 2.6. Theplane-parallel laminated sandstones occur as erosionallybased 25–75 cm thick packages. This facies associationis completely sand prone, with no mud along the laminae(see Table 1).

Facies Association 4 occurs in Area 3 interbeddedwith the tidal bars Facies Association (FA 5).

3.4.1. InterpretationThese plane-parallel laminated sandstone beds that

contain no mud at all along laminae are interpreted to

ypically 25–40 cm. Palaeocurrent measurements showirections towards 250–254◦ SW and 64–80◦ NE.

Facies 2.8 consists of white or pink well-sorted plane-arallel laminated sandstone (Fig. 4). Beds are typically00–125 cm thick.

Facies Association 3 occurs in Area 3 interbeddedith tidal flat deposits of Facies Association 2, upper-ow-regime deposits of Facies Association 4 and tidalar deposits of Facies Association 5.

.3.1. InterpretationThe large upwards-fining bi-directional cross-strata,

ogether with the extensive bi-directional cross-tratification, mudstone drapes and erosional bases, sug-est deposition in subtidal channels.

The bi-directional cross-stratified sets of Facies 2.6ere formed as migrating dunes by traction from revers-

ng tidal currents. Reactivation surfaces are formed ashe subordinate current erodes the lee-face of the duneormed by the preceding dominating current (Dalrymple,992). Mud drapes are deposited from suspension set-ling during slack-water periods. The plane-parallel lam-

be deposited from upper-flow-regime (UFR) currents.The UFR plane-parallel laminated intervals form duringmaximum tidal flow velocities (see Kreisa and Moiola,1986). UFR tidal sand flats have been reported fromthe axial portions of modern tide-dominated macrotidalestuaries (Hamilton, 1979; Lambiase, 1980; Dalrympleet al., 1990, 1992). Current speeds that exceed 2 m/s,and water depths less than 2–3 m have been measured inmodern environments, where upper-flow-regime planebeds form in fine sand (Dalrymple et al., 1990).

3.5. Facies Association 5: tidal bar deposits

Facies Association 5 consists of erosionally based,sand prone, 2–15 m thick packages of gray sandstone(see Table 1 and Fig. 5). Facies Association 5 is volumet-rically dominated by compound cross-stratified sand-stone (Facies 2.5) and bi-polar cross-stratified sandstonewith mudstone drapes (Facies 2.6) that occur super-imposed on larger master surfaces. The cross-stratifiedfacies are interbedded or capped by ripple-laminatedsandstones with mudstone drapes (Facies 2.4), plane-

256 L. Sultan, P. Plink-Bjorklund / Precambrian Research 145 (2006) 243–271

parallel laminated mudstone and sandstone (Facies 2.2),mudstones with polygonal and spindle-shaped sand-filled cracks (Facies 2.1) and in places 50–90 cm thickmudstone beds of Facies 2.9.

The bipolar cross-stratified sandstones of Facies 2.6overlie the basal erosion surface in most places and con-sist of gray sandstone with 5–15 cm thick cross-sets. InFA 5 Facies 2.6 has convex-up erosion surfaces on lee-faces of almost every cross-strata. These erosion surfacesare draped by 1–3 mm of mudstone. Facies 2.6 showsmore erosional features in FA 5, and in places wholeunderlying sets have been eroded and only small ero-sional remnants have been preserved. In some placesload casts occur at bases. In other places clay chips andrip-up clasts with preserved internal structures are doc-umented. Palaeocurrent measurements show directionstowards 222–240◦ SW and 40–52◦ NE.

The beds of Facies 2.1 are much more discontinuouscompared to FA 2, as they pinch out in 2–5 m.

Facies 2.5 consists of light gray sandstone that occursas cross-strata with superimposed ripples (Fig. 5). Themaster bedding surfaces are of low-angle and the super-imposed ripples climb up or down these inclined sur-faces. Sets, typically 5–25 cm thick, are bounded bysharp contacts.

Facies 2.4 is a light gray to pink bi-directional ripple-laminated sandstone, commonly with mudstone drapesalong laminae, within the ripples (Fig. 5). Beds aretypically 5–15 cm thick and commonly extend laterally0.2–2 m. Basal erosion is very abundant and beds appear

3.5.1. InterpretationThe erosional base, extensive trough cross-

stratification, lateral accretion, coarse grain-size,bipolar palaeocurrent directions in adjacent deposi-tional units, together with the significant height ofthe sandbodies suggest deposition in tidal sand barsin a subtidal environment. Such large tidal bars arecharacteristic for seaward portions of most macrotidalenvironments (Dalrymple et al., 1990). Tidal sand barshave been described from estuaries, tide-dominateddelta fronts and tide-dominated shallow-marine settings(sand ridges; Swift, 1975; Dalrymple et al., 1990).

Plane-parallel laminated mudstone with desiccationand/or synaeresis cracks (Facies 2.1) is interpreted tohave formed on top of bar surfaces, which is why theycommonly do not extend laterally for more than a cou-ple of meters. Spindle-shaped cracks are interpreted torepresent synaeresis cracks formed by shrinking due tosalinity changes (see, e.g. Plummer and Gostin, 1981).Polygonal cracks are interpreted to represent desiccationcracks formed during subaerial exposure (Plummer andGostin, 1981; Kidder, 1990; Tanner, 1998).

The low-angle inclined bedding with superimposedripples (Facies 2.5) is interpreted as compound cross-stratification deposited by reversing currents (see Allen,1980). The ripple cross-lamination separated by low-angle inclined set boundaries shows that the dominantand subordinate currents did differ in strength (Allen,1980; Kreisa and Moiola, 1986; Nio and Yang, 1991;Johnson and Baldwin, 1996). The master bedding forms

very disturbed. The overlying set may erode almost thewhole set of ripples. Load casts at bases and reactivationsurfaces within sets occur in places. Bed contacts areundulating and sharp or erosional. Lee-faces of ripplescommonly face towards SW and NE.

Facies 2.2 consists of a continuum of deposits fromdark gray plane-parallel laminated mudstone with thinsandstone interlayers to light plane-parallel laminatedsandstone with thin mudstone interlayers (Figs. 3–5).Load casts occur in places. Bed contacts are flat andbeds are sheet-like, and extend laterally for tens ofmeters. Laminae thickness varies typically between 2and 50 mm, commonly the sandier parts are thicker andthe muddier parts thinner. Individual beds are typically5–25 cm thick.

Facies 2.9 consists of 50–90 cm thick beds ofstructureless mudstone with water-escape structures atbases.

Facies Association 5 occurs in Area 3 interbeddedwith tidal flat deposits of Facies Association 2, tidalchannel deposits of Facies Association 3 and upper-flow-regime deposits of Facies Association 4.

by migration of dunes in the dominant current. Thesubordinate current deposits these superimposed ripplesthat migrate up the master surfaces. The ripple sets thatmigrate down the master surfaces are deposited by thedominant current.

Bi-directional cross-stratification (Facies 2.6) repre-sents deposition as migrating 3D dunes deposited fromreversing tidal currents.

Bi-directional ripple-lamination (Facies 2.4) isdeposited from traction by reversing (i.e. tidal) currents,by migration of ripples. Mud drapes are deposited fromsuspension fallout during tidal slacks. Reactivationsurfaces are formed as the subordinate current erodesthe lee-face of the ripple formed by the precedingdominating current (Dalrymple, 1992). The disturbedappearance of beds also reflects the reversing currents,which commonly erode the underlying set beforedepositing the next. Load casts suggest rather rapiddeposition.

Interlayered mudstone and sandstones (Facies 2.2)indicate alternating deposition from suspension and fromtraction currents. Suspension settling of mud occurs dur-

L. Sultan, P. Plink-Bjorklund / Precambrian Research 145 (2006) 243–271 257

ing tidal slacks while traction deposition of sand occursduring ebb and/or flood currents.

Based on the large thickness of beds of Facies 2.9and the occurrence of water-escape structures (indicat-ing rapid deposition), Facies 2.9 is interpreted to repre-sent fluid mud deposits. Fluid muds occur in estuarieswith high suspended sediment concentrations (Wells,1995).

3.6. Facies Association 6: shoreface deposits

Facies Association 6 consists of a ∼5 m thick unit ofalternating symmetrical-ripple and plane-parallel lami-nated sandstones (Facies 3.1) and swaley cross-stratifiedsandstone (Facies 3.2) (see Table 1 and Fig. 6). In a verti-cal section, Facies 3.2 typically occurs as 0.25–2 m thickpackages, interbedded by 0.20–0.25 m thick packages ofFacies 3.1.

Facies 3.1 consists of alternating dark graysymmetrical-ripple cross-laminated mudstone andplane-parallel stratified sandstone (Fig. 6). The rippleshave straight or slightly bifurcating crests. Beds aretypically flat-based and lenticular, extending laterally2–4 m in 5–25 cm thick beds. Facies 3.1 always occursinterbedded with Facies 3.2 (Fig. 6).

Facies 3.2 consists of gray to red swaley cross-stratified sandstone. Beds extend 2–4 m laterally. Inplaces mudstone drapes occur. Bed thickness is 5–40 cm.Facies 3.2 always occurs associated with Facies 3.1(Fig. 6).

A

3

cald(

dopbt

3

a(

stone beds (Facies 4.2) and soft-sediment deformed bed(Facies 4.3) (see Table 1 and Figs. 7 and 8).

Facies 4.1 consists of plane-parallel laminated lightgray sandstone that fines upwards into dark gray plane-parallel laminated or ripple cross-laminated sandy mud-stone or mudstone (Fig. 7). Bed thickness is 2–20 cm.

Facies 4.2 consists of 10–110 cm thick flat-basedsheet-like beds of with ungraded sandstone, followedby cross-stratified sandstone, plane-parallel laminatedsandstone and plane-parallel laminated sandy mudstoneor mudstone.

Interbedded with Facies 4.2 15–25 cm thick soft-sediment deformed beds (Fig. 8) occur. Facies Associ-ation 7 occurs in Area 2 and interbedded with fluvialdeposits of Facies Association 1 in Areas 3 and 4.

3.7.1. InterpretationUpward-fining, laminated or normally graded sand-

stone to mudstone beds (Facies 4.1 and 4.2) with flator scoured bases and sheet-like geometry are a char-acteristic product of waning, surge-type turbidity cur-rents (Walker, 1967; Middleton, 1993; Kneller, 1995;Kneller and Branney, 1995), similar to Tabc of “Boumasequences” (Bouma, 1962). The surge-type currentshave a general decrease in both mean velocity andsediment concentration backwards from the head (e.g.Kneller and Buckee, 2000; Felix, 2002). Such turbiditycurrents gradually lose their capacity to carry sediment(Hiscott, 1994). Load casts and water-escape structuressuggest rapid deposition.

Facies Association 6 occurs in southward reaches ofrea 1.

.6.1. InterpretationSymmetrical ripples (Facies 3.1) are characteristi-

ally deposited by waves. Mud drapes between ripplesre deposited from suspension settling. Plane-parallelamination is interpreted to be deposited by waves andisplays higher wave orbital velocity than wave ripplesAllen, 1982).

The swaley cross-stratification (Facies 3.2) indicateseposition by storm waves. The sand-prone character,ccurrence of swaley cross-stratification and wave rip-les indicate deposition above the fair weather wavease, in a shoreface, where waves commonly reach downo the bottom (Johnson and Baldwin, 1996).

.7. Facies Association 7: turbidite deposits

Facies Association 7 consists of 2–15 m thick pack-ges of upward-fining sandstone to mudstone bedsFacies 4.1), thick upward-fining sandstone to mud-

Alternation of laminated, traction-deposited intervalswith ungraded intervals (as in Facies 4.2) indicatesdeposition triggered by loss of flow competence, andthrough alternation of the fallout rates. The ungradedintervals reflect rapid fallout rates. The laminatedintervals occurred when the vertical grain flux inter-mittently decreased and enabled a distinct interfacebetween the current and the substrate (Arnott and Hand,1989; Kneller and Branney, 1995) allowing tractionand the development of plane beds and ripples. Thecommonly occurring ball and pillow structures, loadcasts and water-escape structures indicate rapid depo-sition. Plane-parallel mud or plane-parallel laminatedsandy mud is deposited from suspension as the flowwanes.

Soft-sediment deformed beds (Facies 4.3) form onslopes. Displacement is concentrated on a basal slipsurface and the material is unconsolidated, becominginternally deformed as the sediment moves downslope(Stow et al., 1996).

Facies 4.1 and 4.2 differ by: (1) Facies 4.1 hasmonotonous thin beds of sand and mud, while Facies 4.2

258 L. Sultan, P. Plink-Bjorklund / Precambrian Research 145 (2006) 243–271

has substantially thicker beds that indicate longer livedevents; (2) Facies 4.2 is interbedded with slumped beds(Facies 4.3), which are absent in Facies 4.1 deposits. It ispossible that Facies 4.1 represent deposits situated fur-ther away from the main turbidite channel, while Facies4.2 and 4.3 represent deposits in or close to the channel.It is not likely that Facies 4.1 represent distal depositson the basin floor since successions of this facies areinterbedded with fluvial deposits.

4. Lateral and vertical facies transitions

The lateral and vertical facies transitions were docu-mented within the well-preserved areas (Fig. 1). Theseareas are separated from each other by poorly exposedareas or by areas with high degree of tectonic deforma-tion or metamorphic overprint. Faults are not detectedwithin the studied areas, as seen on aeromagnetic mapsand in the field. Block rotation is restricted to the vicinity

Fig. 9. Stratigraphy in Area 1: (1) fluvial deposits with increasing tidal inflAssociation 6) and (3) tide-influenced fluvial deposits (Facies Association 1).

uence towards SE (Facies Association 1), (2) wave deposits (FaciesInset map shows location of logs.

L. Sultan, P. Plink-Bjorklund / Precambrian Research 145 (2006) 243–271 259

of large faults, outside the studied areas. The sedimentarystrata are folded, but sedimentary structures reveal thatthe folding is simple and none of the sedimentary rocksin the studied areas are inverted. This is supported by thefact that the sedimentary structures looks primary, theyare not flattened or sheared. We therefore suggest thatpalaeocurrent measurements together with lateral faciestransitions can be used to infer palaeoshoreline loca-tion. Areas 1 and 2 in the northern portion of the basin,and Areas 3 and 4 in the southern portion of the basinare separated by granitoid intrusions. The correlationbetween these areas is unclear and is not objective for thispaper.

4.1. Area 1: rivers, tide-influenced river andshallow-marine deposition

Stratigraphic way-up in Area 1 is towards NE.The general stratigraphy is following: (1) fluvialdeposits with increasing tidal influence towards SE, (2)tide-influenced fluvial deposits and related shorefacedeposits and (3) tide-influenced fluvial deposits(Fig. 9).

The lower portion in Area 1 consists of ca. 30 m thickfluvial deposits of FA 1. These fluvial deposits showan increasing tidal influence towards southeast acrossa distance of ca. 5 km. In Area 1 fluvial deposits ofFacies Association 1 dominate. Pure fluvial deposits withunidirectional trough cross-stratification (Facies 1.1)with palaeocurrent directions towards SE can be traceddowncurrent into sandstones with regularly spaced reac-tivation surfaces and bi-directional cross-stratification(Facies 2.6).

Approximately 40 m stratigraphically above thetide-influenced fluvial deposits shoreface deposits ofFacies Association 6 occur, exposed in 2–7 m thicksections.

The stratigraphically upper portion of Area 1 consistsof a 10 m thick unit of tide-influenced fluvial deposits(Facies Association 1), dominated by overturned cross-stratified sandstone (Facies 1.1) and regularly spacedreactivation surfaces.

4.1.1. InterpretationThe lateral facies transitions from fluvial to tide-

influenced fluvial towards SE, together with palaeocur-

F ck-beddo rbiditet f logs.

ig. 10. Stratigraphy in Area 2: white glassy quartzite overlain by thiverlain by tide-influenced fluvial deposits (Facies Association 1). Tuhat stratigraphy resembles Areas 3 and 4. Inset map shows location o

ed surge-type turbidites (Facies 4.2 and 4.3 of Facies Association 7)deposits represent deposition in or close to the channel. It is possible

260 L. Sultan, P. Plink-Bjorklund / Precambrian Research 145 (2006) 243–271

Fig. 11. Stratigraphy in lower portion of Area 3: alternation of tide-influenced fluvial deposits (Facies Association 1) with thin-bedded surge-typeturbidites (Facies 4.1 of Facies Association 7). This alternation suggests that deposition took place in a delta front setting. Inset map shows locationof logs in Area 3.

L. Sultan, P. Plink-Bjorklund / Precambrian Research 145 (2006) 243–271 261

rent measurements indicate that main fluvial transportdirections were towards SE. The vertical facies transi-tions from fluvial to shallow marine and back to fluvialindicate a transgressive to regressive succession in Area1 (Fig. 9).

4.2. Area 2: delta distributary and mouth bars ordelta front deposition

Stratigraphic way up in Area 2 is towards SE. In thelower portion in Area 2 well-sorted glassy quartziteswithout clear sedimentary structures occur.

In upper portion and further towards W of Area 2thick sandy turbidites of Facies 4.2 and slumped bedsof Facies 4.3 occur documented in 6–10 m thick succes-sions (Fig. 10). The turbidites vary from mud-prone bedsto sand-prone beds in upwards coarsening successions∼10 m thick.

Some 500 m stratigraphically above the turbidites a∼10 m thick unit of tide-influenced fluvial deposits of FA1 occurs, with palaeocurrent directions towards ∼140◦SE and 320◦ NW.

4.2.1. InterpretationThe vertical facies transition turbidites into tide-

influenced fluvial deposits may suggest depositionin a deltaic environment, where distributary chan-nels alternate with mouth bar or delta front turbidites(Porebski and Steel, 2003). The systematic alterna-tion of thick sandy turbidite beds of Facies 4.2 andslumped beds of Facies 4.3 also suggests deposi-tion in a delta front or mouth bar environment notfar from distributary channels. Palaeocurrent direc-tions in the tide-influenced fluvial deposits suggesta general fluvial sediment transport direction towardsSE–SW.

Fdlflct

ig. 12. (a) Tidal flat deposits (Facies Association 1) typically occur as finiesiccation cracks. (A) Sandier parts of the fining-upwards tidal flat units areaminated (Facies 2.2). Facies Association 2 is interpreted as marginal tidalats, Facies Association 2, from the estuarine environment in Area 3. In fineracks occur, suggesting subaerial exposure. (A) In places synaeresis cracks, she fining-upwards tidal flat units are interpreted as upper supratidal parts of t

ng-upwards units with upwards increasing density of synaeresis andtypically bi-directional cross-stratified (Facies 2.6) or plane-parallelflats of tide-dominated estuaries. (b) Fining-upwards cycles of tidal

r grained parts of the tidal flat units ubiquitous polygonal desiccationuggesting salinity changes, occur and (B) the muddier, upper parts ofhe marginal tidal flats.

262 L. Sultan, P. Plink-Bjorklund / Precambrian Research 145 (2006) 243–271

Fig. 12. (Continued ).

L. Sultan, P. Plink-Bjorklund / Precambrian Research 145 (2006) 243–271 263

Fig. 13. Tidal bar deposits (FA 5) are characteristically sandstone-prone and dominated by bi-directional cross-stratified sandstones (Facies 2.6).Depositional units, 2–15 m thick, are erosionally based and in many places capped by thinner mudstone units (A). (B) In some muddier partssynaeresis cracks occur. Tidal bar deposits are in places interbedded with plane-parallel laminated (Facies 2.8) upper-flow-regime tidal flat deposits.Tidal bar deposits are interpreted as outer-estuarine bars of the tide-dominated estuary. UFR tidal flats occurred landwards from tidal bars, whereestuarine funnel narrowed into a channel.

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4.3. Area 3: delta, tide-dominated estuary and riverdeposition

In Area 3 the stratigraphic way up is towards SSE.The general stratigraphy is following: (1) interbed-

ded fluvial deposits and turbidites (Facies Associa-tions 1 and 7, respectively), (2) tide-dominated estu-ary (Facies Associations 2, 3, 4 and 5) and (3) fluvialdeposits (Facies Association 1) (Figs. 11, 12a, b, 13, 14and 15).

Fig. 14. (a) Tidal channel deposits (FA 3) typically occur as sandstone-pronFacies 2.7), (A) capped by thin mudstone layers, (B) interbedded with sandstoare interpreted as inner-estuarine tidal channel of the tide-dominated estuary.Area 3. This small tidal channel consists mainly of bi-directional cross-stratifi

e bar deposits (large-scale bi-directional cross-stratified sandstones,ne-prone channel fill deposits (Facies 2.6–2.8). Tidal channel deposits(b) Tidal channel (Facies Association 3) in the upper estuarine part ofed sandstone (Facies 2.6).

L. Sultan, P. Plink-Bjorklund / Precambrian Research 145 (2006) 243–271 265

Fig. 14. (Continued ).

The lowermost portion of Area 3 is exposed in thenorthern part of the area (Fig. 11). Facies Association 1occurs as erosionally based 4–19 m thick successionsof trough cross-stratified sandstone (Facies 1.1), bi-directional cross-stratified sandstone (Facies 2.6), com-pound cross-stratified sandstone (Facies 2.5) and plane-parallel laminated sandstone (Facies 2.8). Lee-facesof dunes dip systematically towards S with subordi-nate N directions. Within Facies 1.1 20–35 cm thicksoft-sediment deformed beds occur. The tidal influ-ence decreases upsection. Alternating with Facies Asso-ciation 1 are 0.60–6.50 m thick packages of FaciesAssociation 7, with upward-fining sandstone to mud-stone turbidite beds (Facies 4.1). Monotonous thin tur-bidite beds of sandstone to mudstone are interbed-ded with 40–100 cm thick massive turbidite beds thatcommonly display water-escape structures. Bound-aries between turbidites and fluvial deposits are sharpand erosive in most places. Five cycles of fluvial-turbidite alternations are documented within Area 3(Fig. 11).

A few hundred meters stratigraphically above thelowermost portion tidal deposits of Facies Associations2, 3, 4 and 5 occur in ∼70 m thick succession (Figs. 12a,b, 13, 14 and 15). Two to 10 m thick tidal flat deposits(Fig. 12a and b) alternate with 2–27 m thick packagesof vertically stacked tidal bars (Fig. 13), 0.3–0.8 mthick units of upper-flow-regime tidal flats and 3–12 mthick units of tidal channel deposits (Fig. 14). One ofthe tidal channel fills, 9–12 m thick, consists mainlyo(((rL

with subordinate ∼50◦ NE directions. This channel canbe traced downcurrent for hundreds of meters, and hasmultiple erosion surfaces indicating channel migration.The second channel fill, 3–5 m thick, consists mainlyof bi-directional cross-stratified sandstone (Facies 2.6)and minor plane-parallel laminated sandstone (Facies2.8) (Fig. 14). Lee-faces of dunes display a dominantdip towards ∼250◦ SW with subordinate ∼50◦ NEdirections.

In the upper portion of Area 3, a few hundreds metersstratigraphically above the described tidal deposits flu-vial deposits occur (Facies Association 1). The flu-vial deposits consist of white well-sorted trough cross-stratified sandstone, typically in erosionally based40–80 cm thick sets (Fig. 15).

4.3.1. InterpretationThe alternation of tide-influenced fluvial deposits

with turbidites in the lower portion suggests deposi-tion in delta distributary channels and mouth bars or ona delta front. The alternation could be caused by rel-ative sea level changes (see, e.g. Porebski and Steel,2003), but may as well represent delta lobe switch-ing. The upsection decrease in tidal influence withinthe distributary channels suggest delta progradation, asthe uppermost distributary channels are above the tidallimit.

Occurrence of tidal flats, tidal channels, upper-flow-regime tidal flats and tidal sand bars suggest deposition ina tide-dominated estuary (sensu Dalrymple et al., 1992)

f large-scale bi-directional cross-stratified sandstoneFacies 2.7), bi-directional cross-stratified sandstoneFacies 2.6) and plane-parallel laminated sandstoneFacies 2.8) (Fig. 14a and b). Minor bi-directionalipple-laminated sandstone (Facies 2.4) also occurs.ee-faces of dunes dominantly dip towards ∼250◦ SW

(Fig. 16). The occurrence of upper-flow-regime tidal flatstogether with the tidal sand bars suggests deposition inthe inner estuary. The tidal channels thus represent inner-estuarine tidal channels. The presence of net landwardmovement of sediment from outside the estuary mouth(averaged over a period of several years) is one of the

266 L. Sultan, P. Plink-Bjorklund / Precambrian Research 145 (2006) 243–271

Fig. 15. General stratigraphy of Area 3: (1) interbedded fluvial deposits and turbidites, (2) tide-dominated estuary and (3) fluvial deposits.

L. Sultan, P. Plink-Bjorklund / Precambrian Research 145 (2006) 243–271 267

Fig. 16. Outcrops were ‘walked out’, and the lateral transition from fluvial to more and more tide-influenced environments suggests deposition ina tide-dominated estuary (sensu Dalrymple, 1992). Estuary model redrawn from Dalrymple (1992).

primary features that distinguish estuaries from delta dis-tributaries where the net transport is seaward. Estuariescan form only in the presence of a relative sea level rise(Dalrymple et al., 1992), and the middle portion of Area3 thus represents a transgressive interval.

Vertical transition to fluvial deposits in the upperportion of Area 3 (Fig. 15) suggests a new episode ofregression.

Palaeocurrent measurements indicate that main flu-vial transport directions were towards S in the lowermost

part and towards SW in the middle and upper part of Area3 deposits.

4.4. Area 4: river, tide-influenced delta plain anddelta front

Stratigraphic way up in Area 4 is towards NE. Innortheast a ∼35 m thick unit of fluvial deposits (FaciesAssociation 1) with palaeocurrent directions towards135–139◦ SE occurs. Two kilometers further SE a 5.5 m

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Fig. 17. Stratigraphy in Area 4: (1) fluvial deposits (Facies Association 1) with increasing tidal influence in the dominant palaeocurrent directiontowards SE, (2) surge-type turbidites (Facies Association 7) and (3) tide-influenced fluvial deposits and conglomerate (Facies Association 1). Insetmap shows location of logs.

thick unit of Facies Association 1 with tide-influencedfluvial deposits is overlain by 1.25 m of surge-type tur-bidites of Facies Association 7 (Fig. 17). The turbiditesare overlain by a new succession of fluvial deposits. Thisis a similar setting as in Area 3 and (possibly) Area 2,with alternating fluvial and turbidite deposits.

4.4.1. InterpretationThe lateral facies transitions from fluvial succession

in NW to turbidites in SE suggest a general sedimenttransport towards SE. The alternation of fluvial depositsand turbidites suggests deposition in delta distributarychannels and mouth bars or on a delta front.

5. Discussion

5.1. Tidal signatures in the Vastervik Basin

Descriptions of Precambrian tidal deposits are rare.The metasedimentary succession in the Vastervik Basinis an interesting example of Palaeoproterozoic tide-dominated environments in an area that has experi-enced rather high-grade metamorphism during episodesof intrusion of magmatic bodies, as well as tectonic

deformation. The tidal signatures are displayed in mostdepositional settings in the Vastervik Basin. In places it ispossible to ‘walk out’ transitions from fluvial deposits totide-influenced fluvial deposits (e.g. in Area 1). The com-mon occurrence of tidal signatures together with occur-rence of tide-dominated estuaries (sensu Dalrymple etal., 1992) suggest a macrotidal (i.e. tidal range >4 m)regime on Vastervik palaeocoast, at least in estuariesand, perhaps, also in the river mouth.

5.2. The tidal cycle in the Vastervik Basin

Bi-directional cross-stratified and cross-laminatedsandstone with mudstone drapes occur frequently in thedeposits of the Vastervik Basin, especially in the tidal barfacies association. Many of these bi-directional sets dis-play tidal bundles, where single mud and sand coupletsrhythmically change thickness during the tidal cycle dueto neap-spring variations.

5.3. Basin configuration

The documented depositional environments in theVastervik Basin range from fluvial to estuarine, shal-

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low marine and deltaic. This suggests deposition at abasin margin. Occurrence of turbidites themselves doesnot necessarily require great water depths, but turbiditedeposition is much more common in shelf-edge deltasthan in shallow-water deltas (Porebski and Steel, 2003).

5.4. Location of provenance

Based on lateral facies transitions, main fluvial trans-port directions were towards SE in Areas 1 and 4, andtowards SE to SW in Areas 2 and 3. Main tidal sedi-ment transport directions by tidal currents were from SEin Areas 1 and 4, and from SE–SW in Area 3. Thesedirections are consistent with palaeocurrent measure-ments within the areas. This suggests that the generalpalaeoshoreline was roughly W–E trending, and sedi-ment was derived from at least two different sources. Afluvial drainage area was situated to the present north,and a marine source occurred in the present south. Loca-tion of the marine source is unclear as material can betransported by longshore currents for long distances.

5.5. Process regime change

In the Vastervik succession, in the transgressive inter-vals the depositional regime is tide dominated (tide-dominated estuaries), whereas the regressive intervalsshow river domination (deltas). Such process regimechange from river domination to tide domination on theregressive to transgressive transition has been a com-mtoDP1T(u(wpaMtiuTeFct

shoreline/delta is wave or river dominated. The shallow-marine wave-dominated succession in the VastervikBasin has limited exposures that do not allow to docu-ment whether this succession is regressive or transgres-sive. There are many modern examples, e.g. Nova Scotia,with wave-dominated coastline and tide-dominated Bayof Fundy estuary (Dalrymple and Zaitlin, 1994) whereestuaries occupy coastal valleys, where tidal currents areaccentuated and wave energy is excluded, whereas theopen coastline is wave-dominated.

6. Conclusions

This paper reports a much more complex deposi-tional history of the Vastervik Basin compared to thepreviously suggested deltaic deposition (Gavelin andRussell, 1967; Russell, 1967; Gavelin, 1984). Seven doc-umented facies associations: fluvial deposits, tidal flatdeposits, tidal channel deposits, upper-flow-regime tidalflat deposits, tidal bar deposits, shoreface deposits andturbidite deposits reflect deposition on a continental mar-gin. Fluvial transport directions towards SE–SW suggestthat the palaeoshoreline was roughly E–W trending. Sed-iment was derived from at least two sources, a fluvialdrainage area in the present N and a marine source in thepresent S.

The sediments have accumulated in tide-dominatedestuaries, river-dominated deltas and in wave-dominatedshallow-water environments. Both transgressive andregressive episodes of deposition occurred, with appar-

on assumption, and the tidal effects have been assumedo ‘switch on’ during the transgressive phase of devel-pment (Shanley and McCabe, 1991; Dalrymple, 1992;alrymple et al., 1992; Shanley et al., 1992; Allen andosamentier, 1993; Shanley and McCabe, 1994; Willis,997; Shennan and Andrews, 2000; Yoshida, 2000).idal regime change or enhancement is attributed to1) increasing tidal prism as the rising sea level grad-ally accesses a greater surface area (Dalrymple, 1992),2) increased tidal range associated with an increasedidth of continental shelf (Pugh, 1987) or (3) dissi-ation of wave energy due to the frictional effect onshallow shelf (Swift and Thorne, 1991; Emery andyers, 1996). However, the distributary channels of

he regressive river-dominated deltas also show tidalnfluence. The tidal influence decreases stratigraphicallypwards in successively younger distributary channels.his is an effect of delta progradation and does not nec-ssarily apply that the actual tidal influence decreased.looding of any irregular topography, here distributaryhannels and previous river valleys (estuaries) causes theidal currents to become accentuated, even if the general

ent tidal dominance during the transgressive episodesand fluvial dominance during the regressive periods.

The detailed vertical and lateral correlation was pos-sible only in restricted areas, due to tectonic deformationand metamorphism. Areas 1 and 2 occur in the northernpart of the Vastervik Basin, and Areas 3 and 4 in thesouthern part. In Areas 1 and 3 transgressive to regres-sive deposition occurred. In Area 1 with river depositionto shallow marine and back to river deposition, and inArea 3 river-dominated deltas and turbidites are followedby tide-dominated estuaries and later by river deposition.Area 4 (and possibly Area 2) is transgressive with riverdeposition followed by deltaic turbidites.

The 85 m thick estuarine succession in Area 3 indi-cates that the coastline was macrotidal (>4 m), at leastduring some periods, and suggests a powerful sedimentstorage and active sediment input also from a marinesource.

The study of Vastervik metasediments confirms thatprimary sedimentary structures and geometries can bewell preserved in successions where textures are obvi-ously altered by metamorphism, documented in other

270 L. Sultan, P. Plink-Bjorklund / Precambrian Research 145 (2006) 243–271

metasedimentary successions of Baltica (see Plink-Bjorklund et al., 2005). In Vastervik Basin the meta-morphism has in many cases reached amphibolite facies(Kresten, 1971).

Acknowledgements

This project is funded by grant 5103-1515 from theSwedish Research Council. We thank reviewers E.L.Simpson and D.R. Lowe for their constructive andthoughtful comments.

References

Allen, G.P., Posamentier, H.W., 1993. Sequence stratigraphy and faciesmodel of an incised valley fill: the Gironde estuary, France. J. Sed-imentary Petrol. 63 (3), 378–391.

Allen, J.R.L., 1980. Sand waves: a model of origin and internal struc-ture. Sedimentary Geol. 26, 281–328.

Allen, J.R.L., 1982. Simple models for the shape and symmetry of tidalsand waves; (1) statically stable equilibrium forms. Mar. Geol. 48(1–2), 31–49.

Allen, J.R.L., Banks, N.L., 1972. An interpretation and analysisof recumbent-folded deformed cross-bedding. Sedimentology 19,257–283.

Allen, P.A., Homewood, P., 1984. Evolution and mechanics of aMiocene tidal sandwave. Sedimentology 31 (1), 63–81.

Arnott, R.W.C., Hand, B.M., 1989. Bedforms, primary structures andgrain fabric in the presence of suspended sediment rain. J. Sedi-mentary Petrol. 59, 1062–1069.

Bergstrom, U., et al., 2001. Projekt Smaland, regionalt berg. In: Delin,H. (Ed.), Regional berggrundsgeologisk undersokning. Samman-

Dalrymple, R.W., 1984. Morphology and internal structure of sand-waves in the Bay of Fundy. Sedimentology 31, 365–382.

Dalrymple, R.W., 1992. Tidal depositional systems. In: Walker,R.G., James, N.P. (Eds.), Facies Models—Response to SeaLevel Change. Geological Association of Canada, Stittsville, pp.195–218.

Dalrymple, R.W., Knight, R.J., Zaitlin, B.A., Middleton, G.V., 1990.Dynamics and facies model of a macrotidal sand-bar complex,Cobequid Bay-Salmon River Estuary (Bay of Fundy). Sedimen-tology 37, 577–612.

Dalrymple, R.W., Zaitlin, B.A., 1994. High-resolution sequencestratigraphy of a complex, incised valley succession, the CobequidBay–Salmon River estuary, Bay of Fundy, Canada. Sedimentology41, 1069–1091.

Dalrymple, R.W., Zaitlin, B.A., Boyd, R., 1992. Estuarine facies mod-els: conceptual basis and stratigraphic implications. J. SedimentaryPetrol. 62 (6), 1130–1146.

Emery, D., Myers, K. (Eds.), 1996. Sequence Stratigraphy. BlackwellScience, 297 pp.

Felix, M., 2002. Flow structure of turbidity currents. Sedimentology49, 397–419.

Gavelin, S., 1984. The Vastervik area in south-eastern Sweden. Studiesin Proterozoic sedimentation, high-grade metamorphism and gran-itization. SGU serie Ba, 32. Sveriges Geologiska Undersokning,Uppsala, pp. 1–171.

Gavelin, S., Russell, R.V., 1967. Primary sedimentary structures fromthe Precambrian of southeastern Sweden. Geologiska Foreningeni Stockholm Forhandlingar 89, 74–104.

Gorbatschev, R., Bogdanova, S., 1993. Frontiers in the Baltic Shield.Precambrian Res. 64, 3–21.

Hamilton, D., 1979. The high-energy, sand and mud regime of theSevern estuary, S.W. Britain. In: Severn, R.T., Dineley, D., Hawker,L.E. (Eds.), Tidal Power and Estuary Management. TransatlanticArts Incorporated, Albuquerque, pp. 162–172.

Harms, J.C., 1975. Stratification produced by migrating bed forms,

fattning av pagaende undersokningar. Geological Survey of Swe-den, Uppsala, pp. 65–83.

Beunk, F.F., Page, L.M., 2001. Structural evolution of the accretionalcontinental margin of the Paleoproterozoic Svecofennian orogenin southern Sweden. Tectonophysics 339 (1), 67–92.

Beunk, F.F., Page, L.M., Wijbrans, J.R., Barling, J., 1996. Deforma-tional, metamorphic and geochronological constraints from theLoftahammar-Linkoping Deformation Zone (LLDZ) in SE Swe-den: implications for the development of the Svecofennian orogen.GFF, 118, A9.

Boersma, J.R., 1969. Internal structure of some tidal mega-ripples ona shoal in the Westerschelde estuary, the Netherlands; report of apreliminary investigation. Geologie en Mijnbouw 48 (4), 409–414.

Boersma, J.R., Terwindt, J.H.H., 1981. Neap-spring tide sequences ofintertidal shoal deposits in a mesotidal estuary. Sedimentology 28,151–170.

Bouma, A.H., 1962. Sedimentology of Some Flysch Deposits: AGraphic Approach To Facies Interpretation, Elsevier, Amsterdam,168 pp.

Bridge, J.S., 2000. The geometry, flow patterns and sedimentary pro-cesses of Devonian rivers and coasts, New York and Pennsylvania,USA. In: Friend, P.F., Williams, B.P.-J. (Eds.), New Perspectives ofthe Old Red Sandstone. Geological Society Special Publications,pp. 85–108.

Collinson, J.D., 1996. Alluvial sediments. In: Reading, H.G. (Ed.),Sedimentary Environments: Processes, Facies and Stratigraphy.Blackwell Science Ltd., Oxford, pp. 37–81.

Depositional Environments as Interpreted from Primary Sedimen-tary Structures and Stratification Sequences, Society of Economi-cal Paleontologists and Mineralogists, Dallas, pp. 45–61.

Hiscott, R.N., 1994. Loss of capacity, not competence, as the fun-damental process governing deposition from turbidity currents. J.Sedimentary Res. A64 (2), 209–214.

Hogbom, A.G., Gavelin, A., Hedstrom, H., 1910. Excursions in theArchaean of Southern Sweden. Geologiska Foreningens i Stock-holm Forhandlingar 32, 985–1029.

Johnson, H.D., Baldwin, C.T., 1996. Shallow clastic seas. In:Reading, H.G. (Ed.), Sedimentary Environments: Processes,Facies and Stratigraphy. Blackwell Science, Oxford, pp. 232–280.

Kidder, D.L., 1990. Facies-controlled shrinkage crack assemblages inMiddle Proterozoic mudstones from Montana, USA. Sedimentol-ogy 37 (5), 943–951.

Kneller, B., 1995. Beyond the turbidite paradigm: physical models fordeposition of turbidites and their implications for reservoir pre-diction. In: Hartley, A.J., Prosser, D.J. (Eds.), Characterization ofDeep Marine Clastic Systems. Geological Society of London Spe-cial Publication 94, pp. 31–49.

Kneller, B., Buckee, C., 2000. The structure and fluid dynamics of tur-bidity currents: a review of some recent studies and their geologicalimplications. Sedimentology 47, 62–94.

Kneller, B.C., Branney, M.J., 1995. Sustained high-density turbiditycurrents and the deposition of thick massive sands. Sedimentology42, 607–616.

L. Sultan, P. Plink-Bjorklund / Precambrian Research 145 (2006) 243–271 271

Kreisa, R.D., Moiola, R.J., 1986. Sigmoidal tidal bundles and othertide-generated sedimentary structures of the Curtis Formation,Utah. Geol. Soc. Am. Bull. 97, 381–387.

Kresten, P., 1971. Metamorphism and migmatitization in theVastervik area, SE Sweden. Geologiska Foreningen i StockholmForhandlingar 93 (4), 743–764.

Lambiase, J.J., 1980. Sediment dynamics in the macrotidal Avon Riverestuary, Bay of Fundy. Can. J. Earch Sci. 17, 1628–1641.

Larson, S.A., Berglund, J., 1992. A chronological subdivision ofthe Transscandinavian Igneous Belt—three magmatic episodes?Geologiska Foreningens i Stockholm Forhandlingar 114 (4), 459–461.

Mansfeld, J., 1996. Geological, geochemical and geochronologicalevidence for a new Palaeoproterozoic terrane in southeastern Swe-den. Precambrian Res. 77 (1–2), 91–103.

Miall, A.D., 1996. The Geology of Fluvial Deposits, SedimentaryFacies, Basin Analysis and Petroleum Geology. Springer-Verlag,Heidelberg, 582 pp.

Middleton, G.V., 1993. Sediment deposition from turbidity currents.Annu. Rev. Earth Planet. Sci. 21, 89–114.

Midgley, J.P., Blundell, D.J., 1997. Deep seismic structure and thermo-mechanical modelling of continental collision zones. Tectono-physics 273, 155–167.

Nio, S.-D., Yang, C.-S., 1991. Diagnostic attributes of clastic tidaldeposits: a review. In: Smith, D.G., Reinson, G.E., Zaitlin, B.A.,Rahmani, R.A. (Eds.), Clastic Tidal Sedimentology. CanadianSociety of Petroleum Geologists, pp. 3–28.

Owen, G., 1996. Experimental soft-sediment deformation: structuresformed by the liquefaction of unconsolidated sands and someancient examples. Sedimentology 43, 279–293.

Plink-Bjorklund, P., Bjorklund, L., Loorents, K.-J., 2005. Sedi-mentary documentation of the break-up of Rodinia, OfferdalNappe, Swedish Caledonides. Precambrian Res. 136 (1), 1–26.

Plummer, P.S., Gostin, V.A., 1981. Shrinkage cracks: desiccation or

P

P

R

R

R

Russell, R.V., 1967. Paleocurrent analysis in deltaic Precambrianmeta-sedimentary rocks from Vastervik, Sweden. GeologiskaForeningens i Stockholm Forhandlingar 89, 105–115.

Shanley, K.W., McCabe, P.J., 1991. Predicting facies architecturethrough sequence stratigraphy—an example from the KaiparowitsPlateau, Utah. Geology 19, 742–745.

Shanley, K.W., McCabe, P.J., 1994. Perspectives in the sequencestratigraphy of continental strata. AAPG Bull. 78, 544–568.

Shanley, K.W., McCabe, P.J., Hettinger, R.D., 1992. Tidal influencein Cretaceous fluvial strata from Utah, USA: a key to sequencestratigraphic interpretation. Sedimentology 39, 905–930.

Shennan, I., Andrews, J. (Eds.), 2000. Holocene Land–Ocean Interac-tion and Environmental Change around the North Sea. GeologicalSociety Special Publication, 166, London, 326 pp.

Stow, D.A.V., Reading, H.G., Collinson, J.D., 1996. Deep seas. In:Reading, H.G. (Ed.), Sedimentary Environments: Processes, Faciesand Stratigraphy. Blackwell Science, Oxford, pp. 395–453.

Sultan, L., Claesson, S., Plink-Bjorklund, P., 2005. Proterozoic andArchaean ages of detrital zircon from the PalaeoproterozoicVastervik Basin, SE Sweden: implications for provenance and tim-ing of deposition. GFF 127 (1), 17–24.

Swift, D.J.P., 1975. Tidal sand ridges and shoal-retreat massifs. Mar.Geol. 18 (2), 105–133.

Swift, D.J.P., Thorne, J.A., 1991. Sedimentation on continental mar-gins, I: a general model for shelf sedimentation. In: Swift, D.J.P.,Oertel, G.F., Tillman, R.W., Thorne, J.A. (Eds.), Shelf Sand andSandstone Bodies: Geometry, Facies and Sequence Stratigraphy.International Association of Sedimentologists Special Publication,pp. 3–31.

Tanner, P.W.G., 1998. Interstratal dewatering origin for polygonal pat-terns of sand-filled cracks: a case study from late Proterozoicmetasediments of Islay, Scotland. Sedimentology 45 (1), 71–89.

Visser, M.J., 1980. Neap-spring cycles reflected in Holocene subti-dal large-scale bedform deposits: a preliminary note. Geology 8,543–546.

synaeresis? J. Sedimentary Petrol. 51, 1147–1156.orebski, S.J., Steel, R.J., 2003. Shelf-margin deltas: their stratigraphic

significance and relation to deepwater sands. Earth Sci. Rev. 62,283–326.

ugh, D.T., 1987. Tides, Serges and Mean-Sea Level. John Wiley andSons, Chichester, 427 pp.

eading, H.G., Collinson, J.D., 1996. Clastic coasts. In: Reading, H.G.(Ed.), Sedimentary Environments: Processes, Facies and Stratig-raphy. Blackwell Science, Oxford, pp. 154–231.

eineck, H.E., Sing, I.B., 1980. Depositional Sedimentary Environ-ments. Springer-Verlag, New York, 549 pp.

eineck, H.E., Wunderlich, F., 1968. Classification and origin of flaserand lenticular bedding. Sedimentology 11, 99–104.

Walker, R.G., 1967. Turbidite sedimentary structures and their rela-tionship to proximal and distal depositional environments. J. Sed-imentary Petrol. 37 (1), 25–43.

Wells, J.T., 1995. Tide-dominated estuaries and tidal rivers. In: Perillo,G.M.E. (Ed.), Geomorphology and Sedimentology of Estuaries.Developments in Sedimentology, vol. 53. Elsevier, Amsterdam.

Willis, B.J., 1997. Architecture of fluvial-dominated valley-filldeposits in the Cretaceous Fall River Formation. Sedimentology44, 735–757.

Yoshida, S., 2000. Sequence and facies architecture of the upperBlackhawk Formation and the Lower Castlegate Sandstone (UpperCretaceous), Book Cliffs, Utah, USA. Sedimentary Geol. 136,239–276.