The sedimentation and paleohydrology of the Mesoproterozoic stream deposits in a strike–slip basin (Svinsaga Formation), Telemark, southern Norway

Sedimentary Geology 236 (2011) 239–255

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The sedimentation and paleohydrology of the Mesoproterozoic stream deposits in a strike–slip basin (Svinsaga Formation), Telemark, southern Norway Juha Köykkä ⁎ Department of Geosciences, University of Oulu, P.O. Box 3000, 90014 Oulu, Finland

a r t i c l e

i n f o

Article history: Received 20 September 2010 Received in revised form 4 January 2011 Accepted 25 January 2011 Available online 1 February 2011 Editor: M.R. Bennett Keywords: Braided stream Lithofacies Mesoproterozoic Paleohydrology Telemark Southern Norway

a b s t r a c t Ancient braided fluvial systems were affected by a lack of binding vegetation, under variable paleoclimatic conditions, and possibly characterized by unique fluvial styles during the Precambrian. Thus, the paleohydrological investigations are important in interpreting ancient fluvial systems and a comparison of different fluvial styles/systems. This study reports new paleohydrological and sedimentological data from the ca. 1.17 Ga-old fluvial deposits (Svinsaga Formation) in the Fennoscandian Shield, which is a part of the East European Craton. The Mesoproterozoic Svinsaga Formation overlies a major unconformity (sub-Svinsaga unconformity) in the Telemark Supracrustal rocks (southern Norway). The unconformity corresponding to a ca. 170-Ma hiatus (ca. 1347–1170 Ma) and carry evidence for ancient periglacial climatic conditions. Eleven different sedimentary lithofacies and five lithofacies associations were recognized from the Svinsaga Formation. The lithofacies assemblages indicate that the formation represents a relatively high-energy braided fluvial environment. Massive breccias, massive and inversely to normally graded conglomerates, stratified sandstones and laminated and rippled mudstones, were transported by traction currents, sediment gravity flows and by suspension. The proximal part of the formation contains blockfield/blockslope accumulations and waning flood-flow cycles. These are overlain by hyperconcentrated and debris flow deposits, and ephemeral high-energy bedforms. The distal part of the formation contains dune bedforms and channel lag deposits, which indicate waning discharge fluctuations. The paleohydrological calculations from the channel-fill conglomerates and sandstone bedforms correlate with lithofacies assemblages. The estimation of paleoslope and discharge values supports the presence of a fluvial stream rather than an alluvial fan environment. The stream power and water velocity systematically decreases towards the distal parts of the formation. The abundance of the waning flood-flow cycles, ephemeral high-energy sedimentary deposits, and hyperconcentrated and debris flow deposits is a possible indicator of periglacial melt waters in basin margin strike–slip fault evolution. This melt water combined with abundantly available debris from the nonvegetated high-topographic landscape. © 2011 Elsevier B.V. All rights reserved.

1. Introduction Natural fluvial systems have been operated throughout the geological history, ever since continental masses were able to sustain external surface runoff. Ancient fluvial deposits occur in a wide range of tectonic settings, forming an important part of Precambrian rock strata. The Precambrian fluvial systems were characterized by a lack of binding vegetation, which affected the channel bank stability, channel migration, and bedload-suspension ratios (Mueller and Corcoran, 1998; Tirsgaard and Øxnevad, 1998; Long, 2004). In many settings, source rock material was subjected to intensive weathering, which caused an

⁎ Tel.: + 358 8 5531439; fax: + 358 8 5531484. E-mail address: juha.koykka@oulu.fi. 0037-0738/$ – see front matter © 2011 Elsevier B.V. All rights reserved. doi:10.1016/j.sedgeo.2011.01.010

abundance of bedload material and hyperconcentrated debris flows due to channel bank instability. Thus, a braided fluvial style is generally seen as the most dominant for pre-vegetational paleoenvironments (e.g. Tirsgaard and Øxnevad, 1998; Long, 2004, 2006; Eriksson et al., 2006). Fluvial systems display channel patterns that can be systematically related to sediment types that describe channel banks, sediment loads, and streamflow characteristics. Braided stream systems tend to usually have laterally confined channels with well-defined geomorphological elements (channels, banks, and floodplains), which result in complex facies assemblages, and they are thus part of a continuous series of fluvial forms that develop in quasi-equilibrium with external controls on the river system. The analysis of fluvial deposits usually comprises variations in flow patterns, sediment transport rate, stream flow velocity and depth control of bedform textures and structures, stratification patterns, lithofacies characteristics, and possible largescale evolutionary architectural elements. By utilizing sedimentology

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and the paleohydraulic reconstruction of events, the processes operating in ancient fluvial systems can be better understood. The sedimentology of the ancient braided stream sedimentation systems and channel-fills, in different tectonic settings, has been described in many studies (e.g., Eriksson et al., 2006, 2008; Long, 2006), but only very few paleohydrological studies have been made on ancient fluvial deposits (e.g., Van Der Neut and Eriksson, 1999; Eriksson et al., 2006, 2008; Davidson and Hartley, 2010; Köykkä, 2011). Eriksson et al. (2006, 2008) reported and discussed a possibly unique fluvial style from the Proterozoic (ca. 2.0–1.8 Ga) deposits, in the Kaapvaal Craton. This study contributes to this discussion of ancient fluvial systems and brings new paleohydrological and sedimentological data from the ca. 1.17 Ga-old fluvial deposits in the Fennoscandian Shield, which is a part of the East European Craton. The focus of this study is the approximately 200-m-thick Mesoproterozoic Svinsaga Formation, in southern Norway. The Svinsaga Formation was deposited in a strike–slip basin setting, and affected by periglacial climate before and possibly during sedimentation (Köykkä and Laajoki, 2009). It is composed mainly of sedimentary breccias, conglomerates, sandstones, and mudstones. This paper describes the detailed sedimentology of the Mesoproterozoic conglomerate, sandstone and mudstone sequences, to build a paleogeographic reconstruction for the sedimentary environment, and to calculate paleohydrological estimations of the ancient stream channels.

youngest part of the Shield and includes several sectors with distinct geneses and/or tectonometamorphic histories (Bingen et al., 2005) that were later influenced by Sveconorwegian deformation and metamorphism between ca 1.14 and 0.90 Ga (Bingen et al., 2008). The Southwest Scandinavian Domain is divided into a number of continental domains along Proterozoic shear zones and Phanerozoic faults, which are classified as terranes, sectors, segments or blocks (Bingen et al. 2001, 2005, 2008; Andersen 2005; Åhäll and Connelly 2008) (Fig. 1). The main domains are, from east to west, the Eastern Segment, the Idefjorden terrane, the Bamble–Kongsberg sectors/blocks/ terranes, the Telemark sector/block, and the Hardangervidda–Rogaland block. The Eastern Segment consists of reworked Palaeoproterozoic felsic orthogneisses related to the Transcandinavian Igneous Belt. The Idefjorden terrane exposes the typical “Gothian” rocks, reflecting a volcanic arc setting between 1.64 and 1.52 Ga (Åhäll and Connelly, 2008). The Bamble–Kongsberg sectors include metasedimentary rocks and orthogneisses formed after 1.57 Ga. The Telemark block is characterized by a thick supracrustal sequence, known as the Telemark Supracrustal rocks, which were deposited at and following 1.51 Ga. The Hardangervidda–Rogaland block is exposed to the west of the Mandal– Ustaoset fault and shear zone. It contains gneisses with an increasing metamorphic grade towards the southwest and hosts the Rogaland anorthosite–mangerite–charnockite complex. 2.1. Regional lithostratigraphy

2. Geological setting The Fennoscandian Shield is the northern part of the (proto)Baltica or the East European Craton (review in Bogdanova et al., 2008). It is composed of an Archean core in the northeast and progressively younger Proterozoic crustal domains toward the southwest. The Southwest Scandinavian Domain (Gaál and Gorbatschev, 1987) is the

The Mesoproterozoic volcano-sedimentary Telemark Supracrustal rocks are an important component of the Fennoscandian Shield and the Telemark Block (Dons, 1960a,b; Sigmond et al., 1997; Bingen et al., 2001, 2005; Laajoki et al., 2002). The rocks form an approximately 10-kmthick succession that has been metamorphosed under greenschist– amphibolite facies (Brewer and Atkin, 1987; Atkin and Brewer, 1990).

Fig. 1. Simplified geological map of the Sveconorwegian belt and domains affected by ductile deformation after 0.97 Ga (modified from Bingen et al., 2006). Study area marked. HR = Hardangervidda-Rogaland block, T = Telemark block, B = Bamble block, K = Kongsberg block, I = Idefjorden block, ES = Eastern segment, MU = Mandal-Ustaoset fault and shear zone. The Oslo fjord is not part of this classification.

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The Telemark Supracrustal rocks can be divided into two main successions separated by a first-order unconformity (Fig. 2). The lower succession was deposited between ca. 1.512 and 1.347 Ga and is referred to as the “Vestfjorddalen Supergroup,” which was deposited in the Rjukan Rift Basin (Lamminen and Köykkä, 2010; Fig. 2). The overlying succession was deposited between ca. 1.170 and 1.019 Ga and is separated by the sub-Svinsaga unconformity, which records a ca. 170Ma hiatus in the stratigraphic record (Fig. 2). The Rjukan Group forms the basal part of the Vestfjorddalen Supergroup (Fig. 1B). No depositional basement to the Rjukan Group has been identified because the base of the group progressively grades northwards into a gneiss complex. Bingen et al. (2001, 2005) argued that the detrital zircon age distributions indicate that the Hallingdal complex, located at the northern boundary of the Telemark block near the Åmot–Vardefjell shear zone, is a possible depositional basement to the Rjukan Group. Nevertheless, these rocks are not in contact with one another. The Rjukan Group consists of the Tuddal Formation, which is made up of ca. 1.51 Ga continental felsic volcanic rocks (Dahlgren et al., 1990a,b; Bingen et al., 2005), overlain by an approximately 2-km-thick succession of the volcano-sedimentary Vemork Formation (ca. ≤ 1.495 ± 2 Ga, Laajoki and Corfu, 2007), which is characterized mainly by mafic volcanism (Fig. 2). The stratigraphic position of the Vemork Formation has been a matter of debate (Dahlgren et al., 1990a,b; Laajoki and Corfu, 2007; Köykkä, 2010). According to Köykkä (2010), the volcano-sedimentary Vemork Formation represents the axial part of the Rjukan Rift Basin, where the

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extrusion of abundant mafic lavas reflects a change in the magmatic regime in the rift basin, which was possibly coeval with a syn-rift tectonic phase. The Rjukan Group is overlain by the approximately 5-km-thick sedimentary-dominated Vindeggen Group (Fig. 2), which contains several sandstone- and mudstone-bearing formations deposited in diverse environments and complexly faulted into several internally folded blocks. Sedimentation of the Vindeggen Group is bracketed between ca 1.495 and 1.347 Ga (Laajoki and Corfu, 2007; Corfu and Laajoki, 2008). The tidal-shallow marine Brattefjell Formation represents the topmost part of the Vindeggen Group and the depositional basement for the approximately 200-m-thick Svinsaga Formation, which is the lowermost unit of the Oftefjell Group and focus of this study (Fig. 2). The Vindeggen and Oftefjell Groups are separated by the sub-Svinsaga angular unconformity, which is associated with an at least 170-Ma hiatus (Lamminen et al., 2009). Köykkä and Laajoki (2009) concluded that cryogenic weathering played an important role in the genesis of the pre-Oftefjell Group fracture system associated with the sub-Svinsaga unconformity. The Svinsaga Formation consists mainly of massive breccias, mono- and polymictic conglomerates, stratified sandstones, and mudstones. It is overlain by the Robekk and Ljosdalsvatnet formations, which consists mainly of acid volcanics, volcanic conglomerates and sandstone interbeds (Laajoki et al., 2002). The depositional age of the Oftefjell Group and younger supracrustal units coincides with the early to middle evolution of the Sveconorwegian orogeny. The younger overlying succession consists of coarse-grained sedimentary rocks interbedded with bimodal volcanic rocks. These units are more complex due to deformation and faulting, and the lithostratigraphical correlation of different units is problematic. On a regional scale, the formations are laterally restricted and variable, which suggests that they were laid down in separate basins (Laajoki et al., 2002; Andersen et al., 2004) and possibly in a transtensional tectonic regime (Lamminen et al., 2009). 3. Lithofacies characteristics

Fig. 2. A stratigraphic summary of the Telemark supracrustal rocks. The package is subdivided into the Vestfjorddalen Supergroup (ca. 1.5–1.35 Ga) and overlying early Sveconorwegian units (ca. 1.17–1.019 Ga) separated by the sub-Svinsaga unconformity. Age limits are taken from Laajoki et al. (2002), Bingen et al. (2003, 2006), and Lamminen et al. (2009). SHU = sub-Heddersvatnet unconformity, HF = Heddersvatnet Formation, and GF = Gausdalen fault zone. Dash lines indicate unconformities.

This study focuses on the mountainous Brattefjell, Svafjell, Meien, and Sjånuten areas (Fig. 3A, B), where the Svinsaga Formation is the best exposed and primary sedimentary features are well preserved. The approximately 200-m-thick Svinsaga Formation (Fig. 4) is subdivided into Svafjell, Meisetjørn, Meien, Robekkheim, and Robekktjørn members based on their distinctive lithofacies assemblages and sedimentary characteristics (Table 1). In Brattefjell, Svafjell and Meien, the Svinsaga Formation defines the western flank and the NE hinge of the steeply plunging Meien syncline and is cut by the Marigrønutan fault (Fig. 3A, B). In Sjånuten, the SF defines a northwest trending syncline cut by a fault near Lake Meisetjørn (Fig. 3B). The sedimentary characteristics of the approximately 40-m-thick Svafjell member include mainly massive breccias, matrix- and clastsupported conglomerates, and interbedded cross-stratified sandstones (Fig. 5A). The massive breccias are deposited directly above the fracture zone of the upper member of the Brattefjell Formation, which is defined as a sub-Svinsaga angular unconformity. The Svafjell member is also exposed in the Sjånuten area, which shows features similar to those in the Svafjell area (Figs. 3A, B, and 4). The approximately 40-m-thick Meisetjørn member is exposed in the Sjånuten area, and contains laminated mudstone, clast-supported conglomerates, and cross-stratified sandstones (Fig. 5B). The lowest part of the approximately 30–40-m-thick Meien member, which is best exposed in the Meien area, contains mainly matrix- and clastsupported conglomerates, whereas the upper part of the member contains horizontally laminated sandstones (Fig. 5C). The approximately 90-m-thick Robekheim member is exposed in the Brattefjell, Svafjell and Meien areas (Fig. 3A, B), and contains horizontallaminated sandstones, clast-supported conglomerates, and laminated

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Fig. 3. A) A geological map and member occurrences in the Svafjell and Brattefjell areas (modified from Köykkä and Laajoki, 2009). B) A geological map and member occurrences in the Meien and Sjånuten areas (modified from Köykkä and Laajoki, 2009). Maps illustrate the distribution of the Svinsaga Formation. Locations of the photographs are given. A–F = the locations of the paleohydrological measurements. MF = the Marigrønutan fault zone.

mudstones (Fig. 5D). The topmost member of the Svinsaga Formation, known as the Robekktjørnn member, contains clast-supported conglomerates and trough cross-bedded sandstone (Fig. 5E). The member is approximately 60–70-m-thick and is exposed in the Brattefjell, Svafjell, and Meien areas (Fig. 3A, B). Based on the characteristic lithology, grain size, and sedimentary structures, eleven different sedimentary lithofacies were recognized in the Svinsaga Formation (Table 2). The estimation of the total lithofacies types and the lateral distribution of the different members are shown in Figs. 3A, B, and 6. Petrographically, the Svinsaga Formation conglomerates are quartz-pebble, whereas the sandstones are medium- to coarse-grained subarkoses. The detrital framework of the conglomerate clasts contains N95% quartz in a sericite-rich matrix. The texture is mature to supermature and well- to very well-sorted with angular to subrounded clasts. The detrital framework of the sandstone samples contains 75–95% quartz in a sericite-rich matrix, and the feldspar vs. rock fragment ratio is 3:1. Rock fragments consist of sandstones and chert. The texture is submature to mature and moderately- to well-sorted with subrounded to rounded clasts. 3.1. Lithofacies Bcm (clast-supported massive breccias) The lithofacies Bcm is visible in the Svafjell member, specifically in the Brattefjell, Svafjell and Sjånuten areas (Fig. 3A, and B), where it forms a several-meter-thick succession of clast-supported massive breccias (Figs. 4 and 7A). The nature and the interpretation of the breccias are discussed in Köykkä and Laajoki (2009). The breccias are derived from the Brattfjell Formation quartz arenite with a clast size of b256 mm up to 4.0 m and a mudstone matrix, which may contain small

pebbles. The lower part of the breccia succession is tightly packed, but the clast size and sorting decrease up-section. The breccia clasts are moved and rotated only slightly, and the long-axis orientation is more variable in the upper parts of the succession. The breccias do not show evidence of any preferred orientation of tectonic origin. The margins of the breccia clasts are sharp, rounded, or partly corroded. Köykkä and Laajoki (2009) concluded that the breccias overlying the Brattefjell Formation shallow marine sedimentary deposits indicate an abrupt change in the depositional system. The fact that the breccias are not rotated, or are rotated only slightly, indicates that the material was not transported any long distance. This excludes solifluction, avalanching, rockfall on a talus slope, and debris flows as significant processes in its origin (Köykkä and Laajoki). The authors concluded that the packing of fragments in the breccias favors block accumulation, such as blockfield, blockslope, or a debris-mantled slope, where breccia fragments at the surface are firmly embedded in the matrix grains, associated with possible periglacial processes (e.g., frost heaving). The origin of the breccias was explained by in situ frost shattering of the bedrock, which is supported by the nature of the underlying fracture zone, clast angularity, and the absence of grussification (see Köykkä and Laajoki, 2009). Based on previous studies, this lithofacies is interpreted to represent blockfield or blockslope accumulation. 3.2. Lithofacies Gcm (clast-supported massive conglomerate) Lithofacies Gcm is common in the Svafjell and Meien members, of the Svafjell, Sjånuten and Meien areas (Fig. 3A, B), where it forms a ca. 20m-thick unit of clast-supported conglomerate beds (Figs. 5A, C, and 7B). The conglomerate beds are generally 20 cm to 6.0 m thick and a few

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Fig. 4. A) A stratigraphic summary column of the Svinsaga Formation in the Svafjell–Brattefjell area. Paleoflow measurements from the trough cross-bedded sets and imbricated conglomerates indicate a southwest transport direction (modified from Köykkä and Laajoki, 2009). B) A stratigraphic summary column of the Svinsaga Formation in the Sjånuten area (modified from Köykkä and Laajoki, 2009).

Table 1 Svinsaga Formation members classification, thickness and sedimentary characteristics. Member

Thickness

Sedimentary characteristics

Svafjell member

40 m

Meisetjørn member Meien member

40 m

Massive breccias, matrix- and clast-supported conglomerates, and interbedded cross-stratified sandstones. Laminated mudstone, clast-supported conglomerates, and cross-stratified sandstones. Matrix- and clast-supported conglomerates (in the lower part), and horizontal-laminated sandstones (in the upper part). Horizontal-laminated sandstones, clast-supported conglomerates, and laminated mudstones. Clast-supported conglomerates and trough cross-bedded sandstone.

Robekkheim member Robekktjørnn member

30–40 m

90 m 60–70 m

meters wide, forming a sheet-like geometry. Basal bed contacts, in the Svafjell member, are transitional and partly channelized, eroding the underlying sandstone beds. However, in the Meien member, some of the basal bed contacts are more erosional with the underlying matrixsupported conglomerates. The texture and fabric of the conglomerate beds vary from moderately to poorly sorted and from ungraded to a weak normal coarse tail grading. The conglomerate clasts do not have any significant stratification or imbrication. Clast size varies from 5 to 20 cm, and beds may contain breccia boulders and fragments. The size of the breccia boulders and fragments varies from a few centimeters up to 3 m. The average maximum clast size of the clast-supported conglomerates correlates with the bed thickness. The conglomerate clasts are generally well-rounded and are composed of the Brattefjell Formation (basement) quartz arenite. The matrix is medium to coarse-grained sandstone and

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Fig. 5. A detailed stratigraphical log and lithofacies associations (I–V) in different members of the Svinsaga Formation. A) The Svafjell member is dominated by clast-supported massive conglomerate beds and low-angle cross-bedded sandstones. B) The Meisetjørn member is dominated by a thick mudstone unit and overlying inversely graded conglomerates and sandstones. C) The Meien member is dominated by the clast-supported massive conglomerates overlain by matrix-supported massive conglomerates with a mudstone matrix. D) The log and sketch drawn from the lower and upper parts of the Robekkheim member, which is dominated by clast-supported imbricated conglomerates and horizontally-laminated sandstones with minor interbedded mudstones. E) The Robekktjørn member is composed of clast-supported imbricated conglomerates and trough crossbedded sandstone beds. C = correlation.

does not contain any stratified features. The matrix content increases in the upper part of the beds before passing to the next lithofacies. The clast-supported and poorly sorted nature of the conglomerate beds indicates a cohesive, highly viscous, sluggish, and low-mobility flow, which is likely of subaerial origin. The transitional contacts, in the Svafjell member, indicate depositional processes closely spaced in time, whereas more partly channelized basal contacts in some of the conglomerates are the results of fluctuating and pulsating depositional processes (Larsen and Steel, 1978; Nemec and Postma, 1993). The erosive base, in the Meien member, may indicate a more turbulent flow in an earlier phase or water-rich sediment flows, the movement of which involved intense shearing and turbulent churning (Larsen and Steel, 1978; Nemec and Postma, 1993; Went, 2005). The lack of significant imbrication caused by clast-to-clast collisions or clast shape, the disorganized nature, and the rare vertical clast orientations indicate a rapid movement and/or high viscosity of the flows (Feyter and Molenaar, 1984; Nemec and Steel, 1984). A lack of stratification likely resulted from a highly concentrated dispersion of silt, sand, and gravel that were developed beneath a turbulent heavily sediment-laden flow (Sohn et al., 1999).

conglomerate beds are approximately 3–5 m thick and a few meters wide with a sheet-like geometry. Basal bed contacts are not clearly exposed, but appear to have erosional features with the underlying mudstone and sandstone beds. The texture and fabric of the conglomerate beds vary from moderately- to well-sorted with clear inversely grading due to clast size. The conglomerate clasts do not contain significant stratification, but show a weak imbrication. Smaller clasts tend to occur in the basal parts of the beds, passing upwards into larger clasts. The clasts range from 4 to 20 cm in size, are generally wellrounded, and are composed of the Brattefjell Formation (basement) quartz arenite. The matrix is medium- to coarse-grained sandstone and does not contain stratification. The inversely graded conglomerate beds represent a cohesive, highly viscous, and low-mobility flow deposits, which has a possible subaerial origin. Inverse grading was a result of dispersive pressure where viscosity was low and shear stress was high, which allowed upward migration of the larger clasts to the upper parts of the bed (e.g., Nemec and Steel, 1984). 3.4. Lithofacies Gcim (clast-supported imbricated conglomerate)

3.3. Lithofacies Gci (clast-supported inversely graded conglomerate) The lithofacies Gci is only visible in the Meisetjørn member, in Sjånuten (Fig. 3B), where it forms two individual units of clastsupported inversely graded conglomerates (Figs. 5B and 7C). These

The lithofacies Gcim is a common feature in the upper part of the Robekkheim member, and in the Robekktjørn member, in the Brattefjell, Svafjell, Brattefjell, and Meien areas (Fig. 3A, B), where it forms units of clast-supported normally graded conglomerate beds

J. Köykkä / Sedimentary Geology 236 (2011) 239–255 Table 2 Svinsaga Formation lithofacies classification, associated structures and depositional process interpretation. Sedimentary Lithofacies lithofacies

Sedimentary structures

Depositional process/ interpretation

Bcm

Mostly massive, but clast-size sorting through succession Ungraded, disorganized to weak normal grading Clast-size sorting; inversely graded Normally graded; crudely stratified and imbricated Ungraded and disorganized

Blockfield or blockslope accumulation and bedrock frost weathering Subaerial plastic debris flow (high strength and viscous)

Gcm

Gci

Clast-supported massive breccia; matrix may contain pebbles Clast-supported massive conglomerate; may contain breccia blocks and fragments Clast-supported inversely graded conglomerate

Gcim

Clast-supported imbricated conglomerate

Gmm

Matrix-supported massive conglomerate

Sl

Medium- to coarse grained sandstone; may contain pebbles Medium- to coarse grained sandstone; contains pebbles

Sli

Fl

Medium- to fine grained sandstone; may contain breccias fragments Medium- to coarse grained sandstone; may contain pebbles Mudstone (clay, silt)

Flr

Mudstone (clay, silt)

Sh

St

Low-angle cross-beddings (b15°) Low-angle cross-beddings (b15°); inversely graded Horizontallylaminated

Subaerial plastic debris flow (high strength and viscous) Channel fill and lag deposits; braid bars

Subaqueous resedimented debris flow or mass flow deposits Waning current strength streamflow deposits Hyperconcentrated flow deposits

High-energy bar tops or front dunes and sheets

Trough crossbedded

3-D channel fill dune bedforms

Laminated

Waning stream flow deposits 2-D migrating ripples in waning flows of seasonal or climatic controlled ponded-lake

Laminated and rippled

Fig. 6. Estimation of the total lithofacies distribution in the study area. See lithofacies codes in Table 2 and Fig. 5.

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(Figs. 5D, E, and 7D, E). The thickness of the conglomerate beds varies from around 15 cm to 0.5 m and they are generally a few meters wide with a sheet-like geometry. Basal bed contacts are erosional and penetrate the underlying sandstone beds (Fig. 7D). The texture and fabric of the conglomerate beds vary from moderately- to well-sorted with tight packing. Conglomerate is crudely stratified with significant imbrication (Fig. 7E). The size of the conglomerate clasts varies from 5.0 to 40 cm. The larger clasts and thicker beds tend to occur in the upper part of the Robekkheim member, whereas the Robekktjørn member is mostly dominated by thinner conglomerate strata. The conglomerate clasts are generally well-rounded and are composed of the same material as in previous lithofacies. The matrix is medium to coarse-grained sandstone and contains some stratification. The well-sorted fabric, tight packing, significant clast imbrication and stratified features in the matrix suggest the deposition in a fluvial environment as small transverse or side bar types (cf. Kleinspehn et al., 1984; Long, 2006). The thinner channel lag strata in the Robekktjørn member probably indicate rapidly shifting, shallow braided channels from shallow flows on the tops of braid bars or gravel sheets. Thus, they suggest grain-by-grain tractional bedload conglomerates. Clast-supported, well-sorted nature, imbrication and tight-packing indicate that high mechanical strengths were caused by grain friction (Sohn et al., 1999). Larger clasts probably moved by high-density dispersion along the bottom of stream channels during floods, where the driving mechanism is tangential shear stress exerted by the overflowing turbulent portion of the sediment laden flood flow. Cross-bedded features in the matrix suggest that the conglomerate was likely deposited as residual open-work gravels (Krainer, 1988; Long, 2006). The tightly packed and clast-supported framework indicates that the current was capable of winnowing finegrained sediment; therefore, the matrix of these conglomerates likely presents a late-stage infiltrate. Bed thickness was likely controlled by stream depth, slope and stream power. 3.5. Lithofacies Gmm (matrix-supported massive conglomerate) The lithofacies Gmm is visible in the lower part of the Svafjell member and in the Meien member, in Svafjell and Meien (Fig. 3A, B), where it forms a few-meters-thick unit of matrix-supported conglomerate beds (Figs. 5C and 7F). The conglomerate beds are from 1 to 4 m thick and generally a few meters wide, forming a sheet-like geometry. Basal bed contacts in the Svafjell member are erosional, but in the Meien member, they are transitional, where mudstone partly penetrates the underlying clast supported conglomerate beds. The texture and fabric of the conglomerate beds are generally massive without any stratification or imbrication. The size of the conglomerate clasts is from 5.0 to 20 cm. The Svafjell member beds may contain occasional breccia boulders and fragments with sizes varying from a few centimeters up to half a meter. The conglomerate clasts vary from well- to moderately rounded and are similar to those in the clast-supported conglomerate beds. However, most of the clasts in the Svafjell member tend to be more angular than in the Meien member. The matrix is medium- to coarse-grained sandstone in the Svafjell member, whereas the matrix of the Meien member is composed dominantly of mudstone (Fig. 7F). In the Meien member, the mudstone matrix content increases in the upper part of the beds before passing to the next lithofacies. The massive nature of the matrix-supported beds indicates rapid episodes of debris flow sedimentation, during which sediment concentration is high and probably deposited en masse (Went, 2005). Alternatively, Eriksson (1978) interpreted similar matrixsupported conglomerate beds with well-rounded clasts as a mass flow origin from more distal sources. A lack of grading and stratification as well as poor sorting supports a mass flow origin. The lack of any imbrication suggests deposits from non-sheared high strength or high-viscosity flows (Nemec and Steel, 1984). The matrix maintained clasts in dispersion during the laminar flow. The erosive base in the

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Fig. 7. Photographs of outcrops. A) Massive breccias blocks in the Svefjell member (lithofacies Bcm). The compass is 12.5 cm wide. B) Clast-supported and tightly-packed massive conglomerates in the Svafjell member (lithofacies Gcm). UB = angular breccia block. C) Clast-supported and massive conglomerates in the Meisetjørn member showing distinct imbrication of clasts (lithofacies Gci). The compass plate is 12.5 cm. D) The clast-supported imbricated conglomerate penetrating the underlying horizontal-laminated sandstone bed in the upper part of the Robekkheim member (lithofacies Gcim and Sh). The scale is 15 cm. E) The clast-supported and imbricated channel-fill conglomerates in the Robekkheim member (lithofacies Gcim). F) The matrix-supported massive conglomerate from the Meisetjørn member (lithofacies Gmm). The compass plate is 12.5 cm wide. White arrows indicate stratigraphic top.

Svafjell member may indicate a more turbulent flow in earlier phase or water-rich sediment flows, whereas gradational contacts in the Meien member suggest depositional processes closely spaced in time (Larsen and Steel, 1978; Nemec and Postma, 1993; Went, 2005). The matrix-supported beds, in the Svafjell member, are interpreted to represent either a subaerial plastic debris flow with a high-strength or viscous nature or possible colluvial (hill-slope) deposits. The angularity of the clasts suggests the latter interpretation. However, in the Meien member, the lack of any stratified sandstones, mud content, and upward increase in mudstone matrix-content near the top of the beds may indicate that the beds are of subaqueous origin (e.g., Kelling and Holroyd, 1978; Nemec and steel, 1984). The mudstone matrix suggests the increased mobility of the flow through lubricating clast interaction and maintaining high pore fluid pressures.

3.6. Lithofacies Sl (low-angle cross-bedded sandstone) The lithofacies Sl is only visible in the Svafjell member, in the Svafjell and Sjånuten areas (Fig. 3A, B), where it forms several interbeds of lowangle (b15°) cross-bedded sandstone approximately 5 to 40 cm thick (Fig. 5A). The subarkose sandstone beds are medium- to coarse-grained with poor to moderate sorting. The grains are typically sub-angular to rounded. Basal bed contacts are gradational with the conglomerate and beds contain occasional clasts. The size of the conglomerate clasts varies from 0.5 to 8.0 cm. Clasts are sparsely present throughout beds and show no significant imbrication. The low-angle features of the sandstone foresets indicate a flow regime transitional from dune to upper plane bed stability conditions or vice versa (antidunes), likely represent the migration of low-relief

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bedforms (Røe, 1987; Fielding, 2006). The occurrence of the conglomerate clast throughout the beds indicates that short-lived traction carpets developed beneath a turbulent flow that carried abundant sediment load with high fall-out (Sohn, 1994). 3.7. Lithofacies Sli (inversely graded low-angle cross-bedded sandstone) The lithofacies Sli is found only in the Meisetjørn member, in Sjånuten (Fig. 3B), where it forms an approximately 5-m-thick unit of inversely-graded low-angle cross-bedded sandstone (Figs. 5B, 8A). The thickness of the beds varies from 1.0 to 1.5 m. These sandstone beds are composed of the same material as the matrix of the conglomerate beds. The subarkose is medium- to coarse-grained with poor to moderate

247

sorting. The grains are typically sub-angular to rounded. Basal bed contacts are erosional with the underlying conglomerate beds. The sandstone beds contain small, well-rounded conglomerate pebbles and angular mudstone fragments. The size of the pebbles and fragments varies from 0.5 to 6.0 cm. These pebbles are sparse, and their long-axes are generally aligned parallel to the bedding planes. They form thin strata in the upper parts of the beds (Fig. 8A). This aspect results in a notable inversely graded feature for the beds. The foresets are low-angle (b15°) with wavy features cutting each other. Similar inversely graded and stratified sandstone beds containing pebbles and occurring in conjunction with the debris-flow conglomerates have also been described in other fluvial deposits and were interpreted as hyperconcentrated flow deposits (e.g., Smith and Lowe,

Fig. 8. Photographs of outcrops. A) The inversely graded low-angle cross-bedded sandstone bed from the Meisetjørn member (lithofacies Sli). The mud content usually increases in the upper part of the beds, which contain a layer of the rounded sandstone and angular mudstone clasts. The compass plate is 12.5 cm wide. B) The trough cross-bedded sandstone unit in the Robekkjørn member. The arrow indicates paleoflow direction. The scale is 15 cm. C) The thick unit of the laminated and rippled mudstone in the lower part of the Meisetjørn member (lithofacies Flr). The stick is 1.5 m. D) Lithofacies association I in the Svafjell member. The clast-supported conglomerates are overlain by the low-angle crossbedded sandstone beds with gradational contacts. E) Lithofacies association IV from the Robekkheim member (compared to Fig. 8A). The clast-supported imbricated conglomerates are associated with the horizontal-laminated sandstone beds. The contact is erosional, where conglomerate clasts penetrate the underlying sandstone beds although not clearly seen on the photograph. The pencil is 12 cm. F) Lithofacies association V in the Robekktjørn member. The thin strata of channel lag conglomerates are overlain by trough cross-bedded sandstone beds. The scale is 15 cm. White arrows indicate stratigraphic top.

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1991; Sohn et al., 1999; Marren, 2005; Rodriguez-Lopez et al., 2010). Thus, this interpretation is also suitable for these beds. Hyperconcentrated flows are characterized by a suspended sediment concentration of 40% to 80% by weight that behaves as a moderately turbulent to laminar non-Newtonian fluid and is a mixture between stream flow and debris flow (e.g., Costa, 1988; Marren, 2005). The inversely grainsize segregation is a result of intergranular collisions in a basal traction-carpet layer, which was driven by the shear of a powerful overriding current (Collinson et al., 2006). The development of a dispersive pressure within the layer forces pebbles upwards, which may be augmented by kinematic sieving in which smaller particles tend to fall through the mass of colliding grains. This pebble layer freezes when shear stress fails or if more pebbles are added from suspension. The angular mudstone fragments in the upper parts of the beds probably indicate short transport distance and possible rip-ups. 3.8. Lithofacies Sh (horizontal-laminated sandstone) The lithofacies Sh is the dominant feature in the Robekkheim member and is exposed in several areas. The horizontal-laminated sandstone unit is tens of meters thick, and the bed thickness varies from 0.5 m to 2.0 m (Figs. 5D and 7D). The basal bed contact is sharp and erosional with the underlying sandstone or mudstone beds. The sorting of the subarkose is similar to that observed in the other sandstone units in the study area but is generally finer-grained. These beds locally contain angular fragments in the lower part of the unit. In the upper part of the unit, fragments disappear, and rounded-clasts become more common. The size of the fragments varies from approximately 8.0 to 15.0 cm, and they do not have any preferred imbrication or stratification. Horizontal-laminated sandstone beds are interpreted to represent upper flow regime high-energy flow deposits, which are similar to those described in many other fluvial deposits (e.g., Miall, 1977a,b; Ashley, 1990; Bridge and Gabel, 1992; Skelly et al., 2003; Jensen and Pedersen, 2010). Thus, they likely represent bar tops or front dunes and sheets. The solitary breccia fragments may indicate minor plowing, avalanching, or solifluction processes (see Köykkä and Laajoki, 2009).

This lithofacies is interpreted to represent the deposits of waning current strength, which was weak enough to deposit mud and silt size detritus from suspension (e.g., Ashley, 1990; Bridge and Gabel, 1992; Skelly et al., 2003). The sharp-based layers in the mudstone suggest relatively sudden events, superimposed upon a background of constant sedimentation of fine-grained detritus (Collinson et al., 2006). The graded bases suggest the waning of the suspended load followed by more active episodes (Köykkä and Laajoki, 2009).

3.11. Lithofacies Flr (laminated and rippled mudstone) This lithofacies is only visible in the Meisetjørn member in the Sjånuten area (Fig. 3B), where it forms a laminated and rippled mudstone unit that is at least 5 m thick and generally 100 m wide (Figs. 5B and 8C). It is characterized by distinct, alternating layers of silt and clay-sized detritus. The basal contact of the facies is not exposed. The entire unit forms a wide, tabular-shaped unit. A dark clay-sized detritus dominates the mudstone layers, whereas the lighter and coarser silt and fine-grained silty layers tend to be thinner than clay-sized layers. These layers are generally a few millimeters thick and have gradational boundaries. Some layers contain small scale asymmetric ripple structures with a height b1.2 cm and a length b2.2 cm. The thick and laterally extensive nature of this lithofacies suggests that it may represent a seasonal or other climate-controlled fluctuation in the suspended sediment load within a ponded lake in the basin. The lack of any climbing ripples or convoluted bedding in the pale, silty, and fine-grained layers suggests that the mudstone deposits do not represent a flood basin, crevasse splay, or natural levee. According to Collinson et al. (2006), gradational contact suggests gradually increasing and waning high-discharge episodes, which are common in lake environments. The small-scale ripple marks represent migrating 2-D ripples deposited from a single current event during a waning stage of the flow in the lower flow regime.

3.9. Lithofacies St (trough cross-bedded sandstone)

4. Lithofacies associations

The lithofacies St is only visible in the Robekktjørn member, where it is exposed in several areas. The trough cross-bedded sandstone sets are generally 15 to 110 cm thick, and troughs are generally 2 to 5 m wide, depending on the locality (Figs. 5E and 8B). The sets are commonly grouped as cosets. The subarkose sandstone beds are medium- to coarse-grained and are moderately sorted. The grains range from moderately to well-rounded. These beds do not contain any pebbles or rip-ups, unlike other sandstone lithofacies in the study area. This lithofacies is interpreted to represent low sinuosity channel-fill dunes that are commonly described in fluvial deposits (e.g., Miall, 1977a,b; Galloway, 1981; Røe, 1987; Ashley, 1990; Bridge and Gabel, 1992; Skelly et al., 2003; Scherer et al., 2007; Jensen and Pedersen, 2010). The channels were probably shallow because the large-scale trough cross-bedded structures do not appear to form parts of larger scale macroforms. The cosets are the results of the migration of ripples, which are a combination of a net accumulation of sediment on the bed.

Five different genetically related lithofacies associations were identified from the Svinsaga Formation (Table 3 and Fig. 5), which are described and interpreted here.

3.10. Lithofacies Fl (laminated mudstone) The lithofacies Fl forms only a small part of the total lithofacies in the study area and occurs only as interbeds in the lower part of the Robekkheim member. Laminated mudstone beds are only a few centimeters thick and are characterized by distinct, alternating layers of silt- and clay-sized detritus. The basal bed contacts with the underlying sandstone beds vary from sharp to transitional. The siltstone and mudstone layers are generally a few millimeters thick and have gradational boundaries.

4.1. Lithofacies association I (clast-supported massive conglomerate–lowangle cross-bedded sandstone) Lithofacies association I is common in the Svafjell member, where clast-supported massive conglomerates become sandier up-section and pass into the low-angle cross-bedded sandstone (Figs. 5A and 8D). These associations are approximately 0.50 to 2.0 m thick and form an overall fining-upward succession with transitional contacts. These successions are interpreted to represent waning-flood cycles. Similar fining- and thinning-upward sheet-like successions of clast-supported conglomerates in association with pebbly horizontal to low-angle cross-stratified sandstones have been noted in other alluvial deposits (e.g., Martins-Neto, 1984; Köykkä, 2011). Each sheet records a single depositional/aggradational flooding event, where a waning-flood cycle begins with a clast-supported conglomerates followed by low-angle cross-bedded and pebbly sandstone. Likely, the rising-flood phase scoured the fluvial surface and eroded some parts of the top of the preceding cycle. The gradational contacts indicate the same sedimentation process, and the cross-stratified sandstone cappings of the association suggest a waning current system that can only pick up smaller particles, forming a pebbly bed (Nemec and Steel, 1984).

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Table 3 Svinsaga Formation lithofacies associations, occurrence and depositional process interpretation. Lithofacies associations

Lithofacies

Occurrence in member

Depositional process/interpretation

Lithofacies association I

Clast-supported massive conglomerate–lowangle cross-bedded sandstone Clast-supported massive conglomerate–matrixsupported massive conglomerate Horizontally-laminated sandstone–laminated mudstone Clast-supported imbricated conglomerate– horizontally laminated sandstone Clast-supported imbricated conglomerate–trough cross-bedded sandstone

Svafjell member Meien member Lower part of the Robekkheim member

Waning-flood cycles Subaerial debris flows followed by resedimented subaqueous debris flows or mass flow deposits Upper flow regime flows followed by waning flows

Upper part of the Robekkheim member

Ephemeral flashy flood deposits

Robekktjörn member

Discharged stream fluctuations of waning stream

Lithofacies association II Lithofacies association III Lithofacies association IV Lithofacies association V

4.2. Lithofacies association II (clast-supported massive conglomerate– matrix-supported massive conglomerate) Lithofacies association II is found only in the Meien member (Fig. 5C). The clast-supported massive conglomerate beds pass upwards into the matrix-supported massive conglomerate with a mudstone matrix. Associations are generally more than 2.0 m thick and wide with gradational or erosional contacts. It is likely that the muddy matrix-supported conglomerates can be interpreted as being of subaqueous origin, and they thus represent resedimented subaerial clast-supported conglomerates or mass flow deposits. Possibly, the resedimentation began with partial liquefaction and continued as a series of surging densities and modified grain flows. It is possible that clast-supported conglomerates became draped by mud, were thus rendered unstable and caused resedimentation and mixing. Similar associations are usually described from fandelta settings (e.g., Massari, 1984; Postma; 1984); nevertheless, these reworking associations can also be found in the fluvial settings (cf. Eriksson, 1978; Singh, 1983). 4.3. Lithofacies association III (horizontal-laminated sandstone–laminated mudstone) Lithofacies association III is common in the lower part of the Robekkheim member (Fig. 5D). Associations that are a few meters thick and wide contain horizontal-laminated sandstone beds, which are capped by laminated mudstone. The contacts are predominantly sharp. These associations are common in high-energy braided fluvial settings, where upper-flow regime horizontal-laminated sandstone deposits represent high flow velocities followed by waning flows and deposition of the muddy bed (cf. Miall, 1977a,b; Ashley, 1990; Bridge and Gabel, 1992; Skelly et al., 2003). Thick horizontal-laminated sandstone beds with an association of thin mudstone cappings suggest that high-energy flows have been the more dominant feature forming these associations. 4.4. Lithofacies association IV (clast-supported imbricated conglomerate– horizontal laminated sandstone) Lithofacies association IV is found in the upper part of the Robekkheim member before passing to the next member (Figs. 5D and 8E). This lithofacies contains clast-supported imbricated conglomerates associated with horizontal-laminated sandstone beds with erosional contacts (Fig. 8E). The features of the horizontal laminated sandstone beds are similar to those of lithofacies association III, except for a lack of mudstone interbeds. These associations are typically a few meters thick and wide. The texturally immature conglomerates, clast size, imbrication, crude stratification rudites and association with the upper flow regime sandstone indicate ephemeral flashy flood deposits (cf. Hein, 1984; Nemec and Steel, 1984). The close association between the conglomerate lithofacies and upper flow regime horizontal stratification

indicates catastrophic flood-flow deposits in a turbulent flow (cf. Sohn et al., 1999). The horizontal-laminated sandstone therefore represents falling stages of floods, where the water is not flowing rapidly enough to transport the gravel load, which are common on bar tops and channels (e.g., Middleton and Trujillo, 1984). Thus, the whole association records changes in sorting and reflects a carrying discharge and discontinuous accretion. 4.5. Lithofacies association V (clast-supported imbricated conglomerate– trough cross-bedded sandstone) Lithofacies association V is a common feature in the Robekktjørn member (Fig. 5E). The association contains clast-supported imbricated channel-lag conglomerates followed by trough cross-bedded sandstone beds with gradational contacts (Fig. 8F). These associations are usually a few meters thick and wide. Clast-supported and texturally mature conglomerates, imbrication, and association with the trough cross-bedded sandstone suggest a strong channelized transport, continuous runoff, and effective contemporaneous reworking. The interbedded pebbly trough-cross bedded sandstone associated with the channel-fill conglomerates indicates discharge fluctuations of the stream (e.g., Steel and Thompson, 1983). The deposits settled in frictional freezing of the gravel traction carpets during waning of the stream flood where sand is waning in the tops of gravel sheets from suspension. Thus, the entire association suggests prolonged stream energy weakening or sudden lack of flood streams (e.g., Bridge and Gabel, 1992; Skelly et al., 2003; Jensen and Pedersen, 2010). 5. Paleoflow measurements and paleohydrology 5.1. Paleoflow measurements Paleoflow measurements were obtained from imbricated channelfill conglomerates (lithofacies Gcim) and from trough cross-bedded sets (lithofacies St) in the Robekktjørn member (Köykkä and Laajoki, 2009). The field measurements were corrected for tectonic folding and tilting, which assume that the basin block was not rotated around its vertical axis. The paleoflow measurements indicate unimodal paleopaleoflow directions from the present NE to the present SW (Fig. 4A). 5.2. Paleohydrology Although the paleohydrological calculations for ancient fluvial deposits are estimations, they still have sedimentological significance (e.g., Miall, 1977a; Ethridge and Schumm, 1978; Turner, 1980; Costa, 1983; Eriksson et al., 2006, 2008; Davidson and Hartley, 2010) at least in comparison to other ancient deposits, in distinguishing proximal and distal parts within a formation, or to distinguish ancient channel patterns between alluvial fans and braided streams. Due to erosion of the stream channel-fills and compaction of sediments, these results only give an estimation of the possible minimum paleohydrological

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values. Ethridge and Schumm (1978) suggested that a conservative estimate of a 10% reduction in thickness is appropriate for post-burial compaction of sandstone beds. The paleohydrological parameters for the Svinsaga Formation were calculated and estimated based on the clast sizes of the channelfill conglomerates (lithofacies Gcim) and from trough cross-bed sets (lithofacies St) in the Robekkheim and Robekktjørn members. The measurement sites A–F are illustrated in Fig. 3. The formulas used here are based on studies by Schumm (1968), Costa (1983), and summarized in Eriksson et al. (2006, 2008). 5.2.1. Paleohydrological parameters from channel-fill conglomerates (lithofacies Gcim) According to Costa (1983), the fluvial discharge stream power and average flow velocity can be calculated and estimated from the equations 0:455

v = 0:2 × ðdiÞ

ð1Þ

1:686

ð2Þ

W = 0:09ðdiÞ

where v = velocity of water discharge in ms− 1, W = water stream power in Watts/m2 and di = length of the intermediate axis of the measured largest visible clast. If the cross-sectional area of the channel that contains the measured clast can be estimated, then the discharge and paleoslope of the channel can be calculated from the equations (Köykkä and Laajoki, 2009) ð3Þ

Q =v×A
W=


 9800 × Q × S W×w →S= w 9800 × Q

ð4Þ

where Q = water discharge in cubic meters per second, A = channel cross-section (estimated from dm × w, where dm is mean depth of the channel and w is width of the channel in meters) and S = paleoslope of the channel in mm− 1. 5.2.2. Paleohydrological parameters from trough cross-bedded sets (lithofacies St) According to Allen (1968), the mean water depth (dm) can be calculated from 1:19

h = 0:086ðdm Þ


→ dm =

ð 1 Þ 1:19 h 0:086

ð5Þ

where h = mean thickness of cross-bedded sets. By assuming that the coarse fluvial channel-fill consists of about 5% of silt- and clay-size material, which is also the case in the Svinsaga Formation, the channel width and depth ratio can be estimated from (Schumm, 1968, van der Neut and Eriksson, 1999) F = 225M

−1:08

ð6Þ

where F = alluvial channel width and depth ratio and M = estimated sediment load variable (5%). The width of the alluvial channel and the average daily discharge can be calculated from (Leopold et al., 1964; Schumm, 1968) w = Fdm

ð7Þ

Q m = vA

ð8Þ

where w = channel width, Qm = average daily discharge in m3s− 1, A = channel cross-section estimated from dm × w, and v = velocity of

water discharge (can be assumed as an intermediate value for subaqueous trough cross-bedded bedforms, which is 0.75 ms− 1). The mean bankfull channel depth and width can be calculated and estimated from (Schumm, 1969; Leeder, 1973) db = 0:6M

0:34

0:29

Qm

1:40

ð9Þ ð10Þ

wb = 8:9db

where db = bankfull alluvial channel depth in meters and wb = bankfull alluvial channel width in meters. The average daily discharge can be re-calculated for better accuracy using the following equation (Osterkamp and Hedman, 1982): 1:71

Q m = 0:027wb

ð11Þ

The paleoslope and bankfull water discharge can be calculated from two different equations (Schumm, 1968; Schumm, 1972; Williams, 1978): S = 60M

−0:38

F 0:95 S = 30 w0:98

−0:32

Qm !

ð13Þ

1:21 0:28

Q b = 4:0Ab

ð12Þ

S

ð14Þ

where S = paleoslope of the river channel in mm− 1, Qb = bankfull water discharge, and Ab = db × wb. The drainage area and fluvial stream length can be estimated/ calculated from (Leopold et al., 1964): 0:75

Qb = Ad

0:6

L = 1:4Ad

ð15Þ ð16Þ

where Ad = drainage area in km2 and L = fluvial stream length in kilometers. In natural fluvial systems, although the increase in stream power may lead to bank erosion and the increase of channel width to sustain sediment supply equilibrium, the increase of stream power and velocity should be associated with the bedload grain size and bedform drag (e.g., Bridge, 1993 and references therein). The results of the Svinsaga Formation paleohydrological parameters are presented in Tables 4 and 5. The calculations of stream and water velocity from the channel-fill conglomerate clasts (Eqs. (1), (2) and Table 4) suggest a waning current strength from site A towards sites B and C (Figs. 3 and 9). The discharge values are highest in the Robekkheim member, which could support high-energy flood cycles or sudden channel deepening (Fig. 9). The values obtained from the trough cross-bed sets (Eqs. (1)−(11) and Table 4 indicate that channel depth and width increase from site D towards E and F (Figs. 3 and 10). This result possibly indicates lower flow velocities and paleoslopes in site F and thus increasingly broader braided channels. The channel paleoslope and bankfull water discharge values, which were obtained from trough cross-bedded sets (Eq. (12), Eq. (13), Eq. (14), and Table 5), suggest a fluvial stream deposit rather than alluvial fan sedimentation (Fig. 11). The steeper paleoslope values and higher discharge indicate that site D was possibly closer to the margins of the sedimentary basin and topographic highs. The calculated average values of the water velocity, stream power, mean water depth, channel width and daily discharge in different locations correspond to the sediment supply rates and patterns in different parts of the Svinsaga Formation.

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Table 4 Paleohydrological values calculated from the Svinsaga Formation channel fill conglomerates. Location/Member

di [m]

v [ms− 1]

W [Watts/m2]

Q [m3s−1]

A [m2]

Ch. width [m]

Ch. depth [m]

dm [m]

S [m/m]

A (Robekkheim) A (Robekkheim) A (Robekkheim) A (Robekkheim) A (Robekkheim) Average B (Robekktjørn) B (Robekktjørn) B (Robekktjørn) B (Robekktjørn) B (Robekktjørn) Average C (Robekktjørn) C (Robekktjørn) C (Robekktjørn) C (Robekktjørn) C (Robekktjørn) Average Total average

0.20 0.19 0.23 0.11 0.15 0.18 0.19 0.15 0.18 0.12 0.14 0.16 0.09 0.08 0.11 0.06 0.05 0.08 0.14

0.096 0.094 0.102 0.073 0.084 0.090 0.094 0.084 0.092 0.076 0.082 0.086 0.067 0.063 0.073 0.056 0.051 0.062 0.079

0.0060 0.0055 0.0076 0.0022 0.0037 0.0050 0.0055 0.0037 0.0050 0.0025 0.0033 0.0040 0.0016 0.0013 0.0022 0.0008 0.0006 0.0013 0.0034

0.41 0.32 0.35 0.16 0.28 0.31 0.97 0.60 0.64 0.38 0.21 0.56 0.56 0.16 0.60 0.22 0.21 0.35 0.41

4.41 3.50 4.20 1.90 3.70 3.54 9.50 6.29 6.84 4.50 2.90 6.00 7.65 2.45 9.40 4.00 4.07 5.51 5.02

3.7 3.8 3.8 2.9 3.0 3.4 4.9 4.2 5.0 3.7 3.8 43 4.9 4.7 5.1 4.0 3.7 3,7 3.8

1.5 1.3 1.4 1.4 1.2 1.4 3.0 2.1 2.3 2.1 2.7 2.4 1.5 0.5 2.0 1.0 1.1 1.2 1.8

0.9 0.7 1.0 0.5 1.0 0.8 2.5 1.7 1.8 1.5 1.0 1.7 0.8 0.2 1.2 0.8 0.8 0.8 1.1

0.00000661 0.00000795 0.00000444 0.00000816 0.00000338 0.00000611 0.00000301 0.00000372 0.00000330 0.00000296 0.00000303 0.00000321 0.00000202 0.00000474 0.00000102 0.00000144 0.00000104 0.00000205 0.00000379

di = length of the intermediate axis of largest visible clast; v = water velocity; W = stream power; and Q = stream discharge. A = channel cross-section; dm = channel mean depth; and S = channel paleoslope.

6. Depositional environment All the observed lithofacies and lithofacies associations (channelfill conglomerates, streamflow bedforms, high-energy floods cycles and debris flow deposits) and unimodal paleoflow directions observed in the Svinsaga Formation indicate a braided fluvial environment (cf. Miall, 1977a,b; Galloway, 1981; Ashley, 1990; Bridge and Gabel, 1992; Skelly et al., 2003; Scherer et al., 2007). A schematic paleogeographic model and members distribution of the Svinsaga Formation depositional environment is presented in Fig. 12. The lithofacies assemblages, paleoflow measurements and paleohydrological values indicate that the Svafjell member represents a proximal part of the braided stream passing towards the Meisetjørn and Meien members and again more distal parts of the Robekkheim and Robekktjørn members. The Svafjell member is dominated by massive block accumulation breccias and waning flood-cycle associations (lithofacies association I), which together with the paleohydrological data suggests at least moderate topographic highs.

The correlation of the Svafjell member with the Meisetjørn member is problematic, but it is possible that the lithofacies Gcm conglomerate beds, in the Svafjell member, may be correlated with the lithofacies Gci conglomerates in the Meisetjørn member (Fig. 5A, B). In the Meisetjørn member, the occurrence of the several-meters-thick debris-flow and hyperconcentrated flow deposits above the several meters thick and approximately 100-m-wide mudstone unit is difficult to interpret. Thus, two possible sedimentation models must be considered in the lower part of this member. Either the debris and hyperconcentrated flows overlying the mudstone deposit indicate lateral channel migration above the ponded-lake sedimentation, or they represent some embankment deposits that have been opened and flooded above the lake. Lateral channel migrations of over 100 km have been reported from modern braided rivers (see Miall, 1977a,b and references therein). The lowermost inversely graded conglomerate bed (lithofacies Gci) in the Meisetjørn member can be correlated with the lowermost clast-supported conglomerates in the Meien member (lithofacies Gcm) (Fig. 5B, C). This correlation and differences between these two

Table 5 Paleohydrological values calculated from the Svinsaga Formation trough cross-bedded sets. Location/Member

h [m]

dm [m]

w [m]

Qm [m3s− 1]

db [m]

wb [m]

Qm (2) [m3s− 1]

S1 [m/m]

S2 [m/m]

Qb [m3s− 1]

Ad [km2]

L [km]

D (Robekktjørn) D (Robekktjørn) D (Robekktjørn) D (Robekktjørn) D (Robekktjørn) Average E (Robekktjørn) E (Robekktjørn) E (Robekktjørn) E (Robekktjørn) E (Robekktjørn) Average F (Robekktjørn) F (Robekktjørn) F (Robekktjørn) F (Robekktjørn) F (Robekktjørn) Average Total average

0.15 0.20 0.42 0.12 0.30 0.24 0.60 0.85 0.74 0.68 0.90 0.75 1.10 0.95 1.00 1.05 0.97 1.01 0.67

1.60 2.03 3.79 1.32 2.86 2.32 5.12 6.86 6.10 5.68 7.19 6.19 8.51 7.53 7.86 8.19 7.66 7.95 5.49

63.8 81.3 151.6 52.9 114.3 92.8 204.7 274.2 244.1 227.3 287.7 247.6 340.6 301.1 314.4 327.5 306.4 318.0 219.5

76.4 123.9 431.2 52.5 245.0 185.8 785.3 1410.1 1117.1 969.1 1552.3 1166.8 2174.9 1700.0 1853.0 2011.4 1760.5 1900.0 1084.2

3.6 4.2 6.0 3.3 5.1 4.4 7.2 8.5 7.9 7.6 8.7 8.0 9.6 9.0 9.2 9.4 9.1 9.3 7.2

54.5 66.3 109.9 46.8 87.4 73.0 140.2 177.9 161.8 152.8 184.9 163.5 212.1 191.9 198.7 205.5 194.6 200.6 145.7

25.1 35.1 83.5 19.4 56.4 43.9 126.6 190.1 161.7 146.5 203.2 165.7 256.9 216.5 229.8 243.3 221.8 233.7 147.7

0.0116 0.0104 0.0079 0.0126 0.0090 0.0103 0.0069 0.0061 0.0064 0.0066 0.0059 0.0064 0.0055 0.0058 0.0057 0.0056 0.0058 0.0057 0.0075

0.0170 0.0134 0.0073 0.0204 0.0096 0.0135 0.0054 0.0041 0.0046 0.0049 0.0039 0.0046 0.0033 0.0037 0.0036 0.0034 0.0037 0.0035 0.0072

734.6 1048.9 2642.1 557.8 1736.1 1343.9 4130.8 6401.9 5376.7 4834.2 6880.3 5524.8 8863.8 7365.9 7858.5 8357.9 7562.1 8001.6 4956.8

6629 10,657 36,527 4591 20,866 15,854 66,280 118,873 94,194 81,741 130,860 98,389 183,441 143,317 156,238 169,614 148,430 160,208 91,484

275 365 765 220 547 434 1094 1553 1351 1240 1645 1377 2015 1737 1830 1922 1774 1856 1222

h = mean trough cross-bedded set thickness; dm = mean water depth; w = channel width; Qm = average daily discharge; and db = bankfull channel depth. wb = mean bankfull channel width; Qm(2) re-calculated average daily discharge; and S1 = channel paleoslope (Schumm, 1968). S2 = channel paleoslope (Schumm, 1972); Qb = bankfull water discharge (using values S1 + S2 derived by two); Ad = drainage area; and L = stream length.

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Fig. 9. Histograms of the paleohydrological values estimated from the channel-fill conglomerates. See the measured locations from Fig. 2 (A–C). The plotted bars represent the average values from each location. Avrg = the total average. See the text for discussion.

Fig. 10. Diagrams of the paleohydrological values estimated from the trough cross-bedded sets. See the measured locations from Fig. 3 (D–F). The plotted bars represent the average values from each location. Avrg = the total average. See the text for discussion.

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strong during the deposition of the lower part of the Robekktjørn member but became weaker during deposition of the upper parts of the member (i.e., towards the distal parts of the formation). Lack of a deep-marine basin-fill, evidence of contemporaneous tectonics, the abrupt lateral lithofacies variations, coarse-grained alluvial fan sedimentation, intrabasinal unconformities, bi-modal volcanism, and basin margin faults strongly suggest strike−slip influence for the early Sveconorwegian units and thus for the Svinsaga Formation (cf. Nilsen and McLaughlin, 1986; Mueller and Corcoran 1998). Köykkä and Laajoki (2009) suggested that the periglacial climate and block accumulations played an important role during the genesis of the Svinsaga Formation breccias and possibly during deposition of the fluvial sediments. Thus, it is possible that the occurrence of abundant flood flow cycle sediments and hyperconcentrated and debris flows deposits indicate the influence of periglacial melt waters and topographic highs during the evolution of the basin margin strike−slip system. Fig. 11. An estimation of the channel paleoslope values measured from the trough cross-bedded sets. According to Blair and McPherson (1984), the maximum gradient for rivers is 0.007 m/m and the minimum gradient for alluvial fans is 0.026 m/m. Locations of the measurements are marked in Fig. 3 (D–F).

members indicate that they were possibly deposited in different parts of the channel systems, but closely in space and time. The mass flow dominated Meien member is characterized by lithofacies association II. The Meien member gradually passes to the Robekkheim member, but the contact is not clearly exposed (Fig. 5C, D). The entire Robekkheim member records high-energy streamflow and flood cycles that deposited upper flow regime horizontal-laminated bedforms (lithofacies Sh), imbricated channel-fill conglomerates (lithofacies Gcim), and minor laminated mudstone interbeds (lithofacies Fl) and, thus, lithofacies associations II and IV. Stratigraphically upwards, the conglomerate beds become thinner, and horizontal lamination is replaced by trough cross-bedded sandstone (lithofacies St) and channel lag conglomerate beds (lithofacies Gcim) that form lithofacies association V (Fig. 5D, E). This change marks the boundary between the Robekkheim and Robekktjørn members. The calculated paleohydrological data indicate that flows were relatively

7. Discussion Comparing the estimated paleohydrological paleoslope and discharge values (Fig. 11) with studies from ancient alluvial fan deposits (e.g. Köykkä, 2011), indicates that paleohydrological investigations are important in interpreting ancient fluvial systems. Values obtained from the Mesoproterozoic alluvial fan Heddersvatnet Formation (see Köykkä, 2011) indicate much steeper paleoslopes than those of the Svinsaga Formation. Generally, the greater paleoslope values obtained from the Heddersvatnet Formation clearly indicate relics of steep margin channels, whereas the Svinsaga Formation braided system shows a more gentle paleoslope. The question of “natural depositional gap” between gradient limits of fans and streams (Fig. 11) has also been noted in the previous studies on ancient alluvial deposits (Van der Neut and Eriksson, 1999; Eriksson et al., 2006, 2008; Köykkä, 2011). According to Blair and McPherson (1994), this paleoslope “gap” is between 0.007 m/m (maximum for rivers) and 0.026 m/m (minimum for alluvial fans). Van der Neut and Eriksson (1999) concluded that a unique combination of weathering, lack of vegetation, and steep basin-margins could explain these unique values, whereas Eriksson et al. (2008) concluded that the values could be explained by sedimentation processes and the

Fig. 12. A schematic paleogeographic model and members distribution of the Svinsaga Formation depositional environment (modified from Köykkä and Laajoki, 2009). See the text for discussion. Figure not to scale.

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environment itself. Köykkä (2011) suggested that values plotting within the depositional gap are due to the tectonic–sedimentary evolution of the basin itself. However, as pointed out by Eriksson et al. (2006), it is possible that this “gap” was much smaller during the Precambrian. In the case of the Svinsaga Formation, the plotting of some values to the “depositional gap” may indicate that the proximal part of the formation had been located in topographic highs. This result would correspond with the earlier interpretation presented by Köykkä and Laajoki (2009) that at least the early stage of the Svinsaga Formation sedimentation was influenced by periglacial climates and blockfield or blockslope accumulations. If so, it is possible that the periglacial climate has also affected the numerous flood-flash cycles present in the Svinsaga Formation by generating high-discharge melt water runoffs, at least during the later sedimentation, as suggested in the depositional model earlier in the text. 8. Conclusions The following conclusions can be drawn from this study: (1) The Svinsaga Formation records a relatively high-energy braided stream depositional environment in a strike–slip basin. (2) The formation is dominated by flood flow cycle sediments, hyperconcentrated debris flow deposits and shallow braided channel bedforms with ponded-lake sediments. (3) The calculated paleohydrological data from the Svinsaga Formation correspond with the lateral distribution of the sedimentation patterns and lithofacies assemblages. (4) The combination of sedimentary lithofacies analysis and paleohydrology provides better understanding of ancient fluvial systems and sedimentation patterns. Thus, the paleohydrological estimations are important in comparison of different fluvial styles/systems, in different settings, and in recognizing possible unique fluvial styles during the Precambrian.

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