Int J Earth Sci (Geol Rundsch) (2018) 107:1407–1429 DOI 10.1007/s00531-017-1545-y ORIGINAL PAPER Fluvial responses to the Weichselian ice sheet advances and retreats: implications for understanding river paleohydrology and pattern changes in Central Poland Piotr Weckwerth Received: 25 April 2017 / Accepted: 27 September 2017 / Published online: 23 October 2017 © The Author(s) 2017. This article is an open access publication Abstract The evolution of the fluvial systems during the of lithotype B3 and energy of sedimentary environment was Weichselian Pleniglacial in the Toruń Basin (Central Poland) more stable than during the deposition of lithotype B2. was investigated through sedimentological investigation and paleohydraulic analysis. Within the basin, three fluvial Keywords River style changes · Sand-bed rivers · cycles deposited after successive phases of the ice advance Climate oscillation · Weichselian glaciation · Poland which took place 50, 28 and 20 ka ago. Successions of four fluvial lithotypes characterize each fluvial formation, that are related to the paleoenvironmental changes (e.g., climate Introduction instability and changes in the river regime) which affected the channel hydraulics and morphology. The successions Changes in fluvial styles throughout glacial-interglacial comprise river-style metamorphosis between high-energy cycles were caused by combinations of periodic climate sand-bed meandering rivers (lithotype M1), high-energy cooling and warming, ice sheet margin dynamics, fluctua- sand-bed braided rivers (lithotype B1), and medium-energy tions of the sea level and neotectonic movements (Kozarski sand-bed braided rivers with either unit bars (lithotype B2) 1983; Vandenberghe et al. 1994; Kasse et al. 1995; Huijzer or compound bars (lithotype B3) reflects the maturity stage and Vandenberghe 1998; Holbrook and Schumm 1999; Hui- of sand-bed-braided river evolution in the basin. The assess- sink 2000; Vandenberghe 2002, 2003; Van Huissteden and ment of the fluvial sedimentary environments enabled the Kasse 2001; Houben 2003; Busschers et al. 2007; Erkens construction of a quantitative model of the changes in the et al. 2009). The influence of the climate on fluvial processes river channel pattern in relation to the climate oscillation. has frequently been investigated by establishing variations Both the paleohydrological controls and their sedimentary in time of the pattern of river channels (a.o. Vandenberghe consequences are discussed in the article. Lithotypes M1 1993, 1995, 2002, 2003; Mol 1997; Holbrook and Schumm and B1 represent riverbed modeled under supercritical flow 1999; Van Huissteden and Kasse 2001; Busschers et al. condition. Deposition of lithotype B2 corresponded to the 2005, 2007, 2008; Starkel et al. 2007; Janssens et al. 2012; river channel pattern transformation and was manifested Meinsen et al. 2014); and by determining the changes in by decreasing flow velocity (energy losses associated with the textural and structural characteristics of fluvial sedi- bedform roughness and with the transportation of coarser ments (a.o. Krzyszkowski 1996; Holbrook and Schumm particles). The flow velocity was generally greater in rivers 1999; Huisink 2000; Zieliński and Goździk 2001; Zieliński 2007). Changes of the riverbed morphology over the last glacial-interglacial transition have been found to be a direct result of spatial and temporal response to allogenic forc- * Piotr Weckwerth firstname.lastname@example.org ings (Schumm and Brakenridge 1987; Erkens et al. 2009). These forcings include climatic variations which influenced Department of Geomorphology and Palaeogeography fluvial development by river regime changes, modification of Quaternary, Faculty of Earth Sciences, Nicolaus of the quality and extent of the vegetation cover in catch- Copernicus University in Toruń, Lwowska 1, 87-100 Toruń, Poland ment and valley, the amount of the surface runoff and valley Vol.:(0123456789) 1 3 1408 Int J Earth Sci (Geol Rundsch) (2018) 107:1407–1429 slope instability, limited infiltration due to permafrost, and Poland (Krzyszkowski 1996; Wysota et al. 1996; Zieliński of the type and intensity of the sediment supply to the river and Goździk 2001; Wysota 2002; Molewski 2007; Zieliński channel (e.g., aeolian sediments input and mass movements; 2007; Weckwerth 2011). Thus, this paper bridges between a.o. Weckwerth and Pisarska-Jamroży 2015; Zieliński et al. the two types of insights here, sedimentological interpre- 2015). Such climate-induced and vegetation-related trans- tation of buried fluvial sequences, and between the Late formation of river pattern (from a braided to a meander- Weichselian superficial insights about the climate-derived ing system) were noted between the Late Pleniglacial and transformation of river channel pattern. Late glacial periods in the alluvial reaches of mid-latitude As the fluvial succession filling up the Pleistocene Toruń European rivers (e.g., Vandenberghe et al. 1994; Kasse et al. Basin reaches a thickness of up to about 100 m, sufficient 1995, 2005; Mol 1997; Busschers et al. 2007; Erkens et al. data became available for the reconstruction of the changes 2009; Janssens et al. 2012). of the fluvial system that occurred between the successive The principal indicators of the changes in the fluvial style phases of advance of the Scandinavian ice sheet. Such a are the changes in the morphology and the hydraulic param- reconstruction is the main objective of the present study, eters of the river channel (e.g., channel depth, width, river which focuses on the evolution of the climate-induced fluvial gradient), and the current energy (stream power). In the case processes. Consequently, this paper concentrates on the sedi- of the Toruń Basin during the Weichselian glaciation the mentology and paleohydraulic reconstructions, whereas ear- river regime depended on the extent of the nearby southern lier studies have been carried out describing the chronologi- sector of the Scandinavian ice sheet (Weckwerth 2010, 2013; cal information the fluvial formations described at the same Weckwerth and Chabowski 2013; Pisarska-Jamroży 2015; sites (Weckwerth et al. 2011, 2013; Weckwerth 2013). The Pisarska-Jamroży et al. 2015; Woronko et al. 2015) as well study had the following approach: recognition of the fluvial as to semi-cyclic alternations of climatic cooling and warm- lithotypes and interpretation of the depositional mechanisms ing (references in the previous paragraph). involved, assessment of the quantitative and qualitative fea- Recognition of the types of river channel patterns and tures of the fluvial environment and construction of a model their changes requires the reconstruction of qualitative and of the changes in the river channel pattern during the Weich- quantitative characteristics of the fluvial environment. The selian glaciation that shows its controlling paleohydrological existing models of the river channel pattern transformation parameters and the sedimentary consequences. are based mainly on the relationships that may occur between the flow competences (Leopold et al. 1964; Schumm 1977, 1985) and relationships between the stream power and the Geological setting particle sizes of the transported sediments (Van den Berg 1995; Thorne 1997; Van den Berg and Bledsoe 2003; Anisi- The Toruń Basin, located in the north-central part of Poland, mov et al. 2008). It is quite well understood that the hydrau- belongs to the eastern part of the Toruń-Eberswalde ice- lic parameters of the rivers (flow depth, hydraulic gradient, marginal valley (Fig. 1). During the Weichselian glaciation, channel morphology, water flow velocity and stream power) this basin was fed by precipitation-dependent (periglacial) affect the development of bedforms, to which fluvial litho- rivers flowing from the south, but just as well by streams facies correspond (Ashley 1990; Southard and Boguchwal flowing from the north that were fed by the meltwater from 1990; Southard 1991). The associations of bedforms built the ice sheet (Weckwerth 2013; Weckwerth and Chabowski fluvial lithotypes, which in turn are relevant to recognizing 2013; Pisarska-Jamroży 2015; Pisarska-Jamroży et al. 2015). and discriminating river channel patterns developed under The position and extent of the basin was determined by val- the equilibrium channel state (Leopold 1964; Langbein and leys and glacial basins which evolved before the Last Gla- Leopold 1966; Schumm 1977, 1985; Knighton 1984; Van cial Maximum (LGM), i.e. before 24–19 ka ago (Makowska den Berg 1995; Kleinhans and Van den Berg 2011), and 1980; Wysota et al. 2009; Weckwerth 2010, 2013; Marks relating them to former discharge regimes. 2012). Several previous geomorphological and geological stud- The oldest valleys known in the area became infilled ies of Weichselian outwash plain fluvial successions have by fluvial sands and gravels which accumulated during had the objective to quantitatively characterize these envi- the Holstenian interglacial (western part of Toruń Basin; ronments, as does this study (a.o. Zieliński and Van Loon Brykczyński 1986; Uniejewska and Nosek 1992; Marks 1999, 2003; Blažauskas et al. 2007; Pisarska-Jamroży 2015). 2005). Palynology records and results of OSL dating These past studies, however, did not consider to analyse the (Nor yśkiewicz 1978; Makowska 1980; Wysota 2002) river regimes and their hydraulic parameters, in addition to indicate that a next generation of river valleys evolved changes in the ice-sheet edge position, also changed coevally during the Eemian interglacial (Figs. 2, 3) when rivers due to climate fluctuations in the case of fluvial succes- bypassed parts of the Mesozoic basement in the area that sions which nowadays are buried under Weichselian tills in had become uplifted by glacio-isostasy (Weckwerth 2013). 1 3 Int J Earth Sci (Geol Rundsch) (2018) 107:1407–1429 1409 Fig. 1 Location of the study area, showing the extent of the Scandinavian ice sheet during the Weichselian glaciation Fig. 2 Relief of the Toruń Basin and the location of research sites and geological cross-sections (research sites: 1—Zielonczyn, 2—Chobielin, 3—Rozwarzyn, 4—Młyniec, 5—Nowe Dąbie, 6—Wypaleniska, 7—Paterek) This glacio-isostatic rebound induced river incision at development of Last-Interglacial valleys in areas of Meso- the end of Saalian glaciation (Termination II; postglacial zoic basement subsidence. transgression and sea-level stabilization; Marks 2005; Weichselian fluvial and glaciofluvial successions are Siddall et al. 2006) and was also a major cause of the the fills of these Termination-II created valleys. They 1 3 1410 Int J Earth Sci (Geol Rundsch) (2018) 107:1407–1429 Fig. 3 Geological cross-sections in the western part of the Toruń Basin (after Weckwerth 2010, 2013, modified) constitute three fluvial formations separated by tills or and 21 ka years ago) and the Noteć Formation, which by glaciolacustrine sediments (Fig. 4). The first fluvial includes deposits of the Toruń-Eberswalde ice-marginal formation is the Rzęczkowo Formation, which is over valley (pradolina) that dates from 17 to 16 ka years ago 29 ka years old (Wysota 2002; Wysota et al. 2009; Weck- on the base of geomorphological records and 14C dating werth et al. 2011). The two younger fluvial formations (Galon 1961; Tomczak 1987; Wiśniewski 1990; Weckw- are the Zielonczyn Formation (deposited between 28 erth et al. 2011; Weckwerth 2013). The lower members 1 3 Int J Earth Sci (Geol Rundsch) (2018) 107:1407–1429 1411 Fig. 4 Synthetic lithological and stratigraphical cross-section of the Toruń Basin and Noteć River valley (after Weckwerth 2013, modified) Fig. 5 Toruń Basin and its fluvial system during the deposition of the Rzęczkowo and Zielonczyn Formations (after Weckwerth and Chabowski 2013, modified) of the Rzęczkowo Fm. and Zielonczyn Fm. were accu- narrow valleys running from the south and tracks of mulated in precipitation-fed rivers (Figs. 4, 5), whereas meltwater streams flowing from the north (Fig. 5). This the sediments of their upper members were deposited by basin was drained to the west of Bydgoszcz through the meltwaters (Wysota et al. 1996; Wysota 2002; Molewski antecedent river gaps which developed near Nakło due 2007; Weckwerth 2013; Weckwerth et al. 2013). All to glacio-isostatic rebound that induced river incision at members of the Rzęczkowo and Zielonczyn Formations the end of Saalian and Weichselian glaciations (Figs. 2, were deposited in the Toruń Basin which was widened 5) (Weckwerth 2013). and deepened by glacial erosion and deformation dur- ing the Late Pleistocene (Weckwerth 2013). This large depression represents a sedimentary basin which merges 1 3 1412 Int J Earth Sci (Geol Rundsch) (2018) 107:1407–1429 Table 1 Symbols of lithofacies and codes for channel bedforms and macroforms (after Miall 1978, 2006; Eyles et al. 1983; Zieliński 1995, 1998; Zieliński and Pisarska-Jamroży 2012; modified by the present author) Lithofacies code Lithogenetic code Textural symbols Structural symbols D—diamicton m—massive GL—glacial deposition G—gravels h—horizontal stratification, lamination GB—gravelly bedforms and bars S—sands l—low-angle cross-stratification (small and medium scale) SP—scour-pool infill, scouring channel F—silt and/or clay i—low-angle inclined stratification (large scale) CC—chute channel t—trough cross-stratification UP—upper plane bed p—planar cross-stratification (small and medium scale) SB—sandy bedforms f—large scale planar cross-stratification FM-u—transverse bars r—ripple cross-lamination FM-c—transverse bars w—wavy lamination LM—side bars, point bars e—scour infill OF—overbank forms: crevasse splays and sand sheets formation which can be divided into members character- Methods ized by specific lithology (Tucker 2003; Marks et al. 2014). The ages of fluvial formation were obtained by OSL dating Lithofacies analysis of fluvial sediments method, applying the single aliquots regenerative (SAR) dose protocol (Murray and Wintle 2000; Wintle and Mur- Field data from seven key sites show that the fluvial depos- ray 2006). The ages dataset was described and discussed in its occur below the Weichselian tills and till-remnants. author’s earlier studies under the terms of ages overestima- The conditions at which fluvial deposition took place tion and reliability in paleogeographical context (Weckwerth were reconstructed on the basis of the structural and tex- et al. 2011, 2013; Weckwerth 2013). tural features. This made it possible to restrict the genetic origins of the various lithofacies to bedforms of specific Paleohydraulic analysis channel types. In the studied exposures, several lithofacies coexist and form lithofacies associations related to specific A series of paleo-flow parameters were quantified in the fluvial sub-environments (Table 1). In turn, these lithofa- following manner. The river gradient (S) for the fluvial cies associations form groups of macroscale strata sets successions of the Rzęczkowo and Zielonczyn Formations (cf. Bridge 1993), that we have related to the depositional was calculated using the Du Boy equation τ = ρgDS and the sedimentary environments. Schields equation θ = τ/ρgRd (Paola et al. 1992; Julien and Wargadalam 1995; Dade and Friend 1998; Dade 2000; Paola Grain‑size distribution 2000; Frings 2008; Weckwerth 2011), where τ is the bed −1 −2 −3 shear stress (kgm s ), ρ is the water density (kgm ), D The grain-size distribution of the sands and gravels was is the hydraulic radius (mean flow depth for braided rivers) determined by sieving at 1-φ intervals, whereas the silt (m), θ is the dimensionless Shields parameter, g is the gravi- and mud fractions were measured with a laser particle-size −2 tational acceleration (= 9.81 ms ), R is the relative excess analyser (Analysette 22) at 0.25-φ intervals. The sediments density of the sediment particles, and d is the median grain were classified texturally according to the scale proposed diameter (m) (Fig. 6). by Udden and modified by Wentworth (1922). The statisti- The flow depth D (m) is related to the height of the bed- cal parameters of the grain-size distribution (median grain forms and macroforms in a river channel (Simons and Rich- diameter, mean grain-size and sorting, skewness and kurto- ardson 1962), and thus proportional to the thickness of the sis of the grain-size distribution) were assessed on the basis individual lithofacies of planar and trough cross-stratified of the formulas by Folk and Ward (1957). sands (Simons and Richardson 1962; Allen 1982; Weck- werth 2011; Zieliński 2014) (Fig. 6). The fluvial erosion Age control effects were taken into account during the estimation of flow depth (D) because in braided rivers the mean preserved set Fluvial successions that compose lithostratigraphic unit, thickness for bedforms is about 0.7–0.8 (cf. Paola and Borg- identified by lithological characteristics and stratigraphic man 1991; Mohrig et al. 2000). For massive or horizontally position, distinct from other beds, were classified as 1 3 Int J Earth Sci (Geol Rundsch) (2018) 107:1407–1429 1413 Fig. 6 Scheme of paleohydraulic calculation principles for buried flow depth calculated on the basis of transverse bar height, t —thick- sedimentary facies of sand-bed-braided rivers. t —thickness of litho- ness of channel and scour infill, D —mean flow depth calculated on p f facies of planar-stratified sands (medium bedded), h —megaripple the basis of thickness of channel and scour infill, d —95th percentile mp 95 height calculated on the basis of the thickness of lithofacies of planar- of grain-size distribution, D —mean flow depth calculated for plane pb stratified sands, t —thickness of lithofacies of trough-stratified sands bed (after Julien and Raslan 1998), d*—dimensionless particle diam- (medium bedded), h —megaripples height calculated on the basis eter (after Julien and Raslan 1998) [other symbols are explained in mt of the thickness of lithofacies of trough-stratified sands, D —mean the text; h was calculated after Saunderson and Jopling (1980) and mp mp flow depth calculated on the basis of the thickness of lithofacies of Zieliński (2014), h and D after Leclair and Bridge (2001), D mt mt mp planar-stratified sands, D —mean flow depth calculated on the basis after Zieliński (2014), h after Paola and Borgman (1991), D after mt b b of the thickness of lithofacies of trough-stratified sands, t —thick- Mohrig et al. (2000), V and V after Williams (1983), S after Paola b 1 2 ness of planar cross-stratified sands (large and very large bedded) or et al. (1992), Dade and Friend (1998), Dade (2000), Paola (2000) and cosets of cross-stratified sands, h —transverse bar height, D —mean Frings (2008)] b b stratified sediments, the estimation of the flow depth was Sedimentological characteristics of the fluvial based on the median grain diameter (d ) (Williams 1983; deposits Julien and Raslan 1998). The results are consistent with the method that determines the flow depth on the basis of Sediments of the lower members of the Weichselian the lithofacies thickness (Leclair and Bridge 2001; Bridge fluvial formations 2003). The average current velocity was estimated with Man- Fluvial succession at the Chobilelin site 2/3 1/2 −1 ning’s equation, V = (D S )n , where n is Manning’s coefficient (see Julien 2002). The Froude number was Sediments which represent the lower member of the determined to compare with the sedimentological inferences Rzęczkowo Fm. are present at the Chobielin site and and to establish an independent estimation of flow regime. were deposited between 38.3 ± 4.7 and 29.5 ± 3.9 ka This non-dimensional parameter was calculated using the ago (Figs. 4, 7). These are 9 m thick and consist of three −0.5 equation Fr = V(gD) (Dey 2014) which shows the ratio similar lithofacies associations—Si(SFw), Si(Sm,Sr) and of inertia forces to the gravity forces (Fig. 6). The sinuos- Si(Sm,Sr,Src,Se)—in which a large-scale and low-angle ity of the river channel was estimated using the equation inclined stratification of fine-grained sands (Si) dominate −2 SN = 1 − (VR/252) , where SN is the channel sinuosity and (Fig. 7). The laminae of lithofacies Si dip up to 8° to the VR is the maximum angular range of the mean azimuth of south and the south-east. Sands and silts with wavy lamina- the crossbeds (Langbein and Leopold 1966; Miall 1976). tion (SFw) and ripple cross-lamination (Sr) and beds with climbing ripples: (lithofacies Src), scour infillings (Se) and massive sands (Sm) are also present in this fluvial succes- sion. All the mentioned lithofacies form fining-upward sequences Sm→Si→Sr(Src)/Se. 1 3 1414 Int J Earth Sci (Geol Rundsch) (2018) 107:1407–1429 Fig. 7 Lithofacies log of the Zielonczyn and Chobielin sites (explanation of lithofacial and lithogenetic codes in Table 1) 1 3 Int J Earth Sci (Geol Rundsch) (2018) 107:1407–1429 1415 1 3 Fig. 8 Lithofacies log of the Rozwarzyn, Młyniec and Nowe Dąbie sites (for legend, see Fig. 7 and Table 1) 1416 Int J Earth Sci (Geol Rundsch) (2018) 107:1407–1429 The fluvial succession at Chobielin can be divided into Fluvial deposits at the Rozwarzyn, Młyniec and Nowe Dąbie sites two parts on the basis of the grain-size distribution. The lower part occurs at the depth of 9–5.6 m and consists of Fluvial deposits of the lower members of Weichselian fluvial well-sorted fine-grained sands, in which the average value of median grain diameter (d ) is 0.14 mm. The upper part, formations were also found at the Rozwarzyn, Młyniec and Nowe Dąbie sites (Fig. 8). These sediments are 5 m thick between 5.6 and 1.7 m, is built of moderately sorted fine- grained sands in which the median grain diameter (d ) is and contain three lithofacies associations at the Rozwarzyn site. The first of them [lithofacies association Sh,Sl(Sr,Sp)] about 0.19 mm. is dominated by middle- and small-scale beds of horizon- tally stratified and low-angle cross-stratified sands (median Fluvial succession at the Zielonczyn site grain diameters d are in the range of 0.15–0.16 mm). These sediments are well sorted and have symmetric or positively Sediments of the lower member of Zielonczyn Fm. are exposed at the Zielonczyn site and reach a thickness skewed grain-size distributions. Locally, they are interbed- ded with ripple cross-laminated sands and with small- and of 19.5 m. These were deposited between 26.9 ± 3 and 25.4 ± 3 ka ago (Figs. 4, 7) and consist of six lithofacies middle-scale lithofacies of planar cross-stratified sands (Sp) which are coarser (average d = 0.22 mm) and moderately associations (Fig. 7). The basal part of the first lithofacial association (GSp,Sr) is built of cross-stratified and poorly sorted. Additionally, normally graded, large-scale and low- angle cross-stratified sands (Sl) occur in the upper part of sorted sandy gravels (lithofacies GSp) in which the average median grain diameter (d ) is 1.77 mm. These sediments the lithofacies association Sh,Sl(Sr,Sp). This lithofacies is characterized by sigmoidal crossbeds and varied thicknesses are overlain by ripple cross-laminated and well-sorted fine- grained sands (lithofacies Sr) in which the average median (from 0.7 up to 2.5 m). The lower members of the Weichselian fluvial succes- grain diameter (d ) is 0.15 mm. The grain-size distributions of lithofacies association GSp,Sr are symmetric and meso- sions are represented by lithologically similar deposits at the Młyniec and Nowe Dąbie sites (Fig. 8). Sands showing or platykurtic. The next lithofacies associations [Si,Sm(Se), Si(Sr) and Si(Sr,Se)] form three similar sequences medium-scale and small-scale planar cross-stratification (Sp) and rippled fine-grained sands were encountered in Sm→Si→Sr/Se. Their lower segments are composed of massive fine-grained sands (Sm), in which median grain both sites; scour infillings additionally occur (lithofacies Se). These deposits are moderately well and moderately diameters (d ) vary between 0.15 and 0.16 mm. Middle seg- ments consist of sets of large scale and low-angle inclined well sorted, while their median grain diameter (d ) varies from 0.15 up to 0.29 mm. Moreover, massive and matrix- stratification of fine-grained sands (Si) in which laminae dip at 6°–14° to the northeast, southeast and southwest supported sandy gravels (lithofacies GSm) were recognized at the Młyniec site, whereas the lithofacies association (Fig. 7). Sediments of lithofacies Si are normally graded (the median grain diameters decrease from 0.17 to 0.15 mm) Sr(Src,SFh) was observed at Nowe Dąbie site. Sandy beds of the ripple-drift cross-lamination occur commonly in this and well or moderately sorted. The uppermost segments of each sequences comprise rippled fine-grained sands (Sr) in lithofacies association. Additionally, horizontally laminated silty sands (lithofacies SFh) were identified at a depth of which the median grain diameters differ between 0.13 and 0.16 mm. These segments may contain sandy scour infillings 7.2–7.6 m (Fig. 8). of a depth up to 0.4 m (lithofacies Se, median grain diameter d between 0.16 and 0.18 mm). Sediments of the upper members of the Weichselian fluvial formations The upper part of the fluvial succession at the Zielonczyn site is composed of two lithofacies associations Sr,Sh(Sm) The upper member of the Zielonczyn Formation at the Nowe and Sl(Sh,Sr)(d) (Fig. 7). The first is dominated by rip- ple cross-laminated fine-grained sands (Sr) the thickness Dąbie site contains two lithofacies associations Sp(Sr,Se) and Sf,Sp(Sl,Sr,Sh) (Fig. 8). The first of them is 2.2 m thick of which is 2 m. These sediments lie on the horizontally stratified sands (Sh) or on massive sands (Sm). The second and comprises medium-scale planar cross-stratified sands (lithofacies Sp, profile Nowe Dąbie 2; Fig. 8). These reverse lithofacies association [Sl(Sh,Sr)(d)] comprises lithofacies of medium-scale and low-angle cross-stratified fine-grained graded sediments (median-grain diameter d increases from 0.29 up to 0.34 mm) are moderately well and well sorted sands (Sl), which are interbedded with ripple cross-lami- nated sands (Sr) and with horizontally stratified sands (Sh). and interbedded with scours infillings (lithofacies Se) or with ripple cross-laminated fine-grained sands (Sr). Litho- Their sedimentary structures were deformed by the develop- ment of reverse faults and clastic dykes. facies association Sf,Sp(Sl,Sr,Sh) has a thickness of 4.75 m. Its lower part is dominated by fine- and medium-grained sands of inclined stratification (lithofacies Sf, median-grain 1 3 Int J Earth Sci (Geol Rundsch) (2018) 107:1407–1429 1417 Fig. 9 Lithofacies log of the Wypaleniska and Paterek sites (for legend, see Fig. 7 and Table 1) diameter d from 0.16 up to 0.31 mm) which reach a thick- and moderately well and well sorted, while their grain-size ness up to 2.1 m and are well and moderately sorted (profile distributions are negatively skewed and leptokurtic. Nowe Dąbie 1; Fig. 8). These sediments are overlain by low- Sediments of the Zielonczyn Fm. are represented by two angle cross-stratified sands (Sl), sands of scours infills (Se) lithofacies associations Sp,SGp(Sr,GSm) and Sf(Sr,Se,Sl) at and by small-scale planar cross-stratified sands (Sp) with the Wypaleniska site (Fig. 9). The first lithofacies associa- interbeds of horizontally stratified sands (Sh). Similar large- tion comprises cosets of medium-scale lithofacies of planar scale sandy beds of planar and inclined stratic fi ation (lithofa - cross-stratified fine- and medium-grained sands (median- cies Sp and Sf, respectively) were recognized in the upper grain diameter d changes from 0.17 to 0.39 mm, profile part of the fluvial succession at the Rozwarzyn site (Fig. 8). Wypaleniska 1) or two-part sequences Sp→Sr which consist Fluvial sediments at the Wypaleniska site represent of planar cross-stratified sands (Sp) with cappings of rippled two Weichselian fluvial formations which were deposited and fine-grained sands (Sr) (profile Wypaleniska 2). Sands up to 29 ± 3 ka ago (Rzęczkowo Fm.) and between 24 and of planar cross-cosets are moderately well and moderately 22 ka ago (Zielonczyn Fm.; Weckwerth et al. 2011; Weck- well sorted, whereas their grain-size distributions vary from werth 2013). The first of them (Rzęczkowo Fm.) consists of negatively to very positively skewed. The second lithofacies medium- and fine-grained planar cross-stratified sands with association [Sf(Sp,Sr,Se,Sl)] is dominated by very thickly an admixture of granules and pebbles [lithofacies associa- planar cross-bedded fine-grained sands (median-grain tion Sp(SGp), Fig. 9]. These deposits are normally graded diameters vary between 0.21 and 0.32 mm; Fig. 9). These (median-grain diameter d decreases from 0.40 to 0.17 mm) sediments are overlain by medium- and small-scale planar 1 3 1418 Int J Earth Sci (Geol Rundsch) (2018) 107:1407–1429 Fig. 10 Typical Weichselian fluvial lithofacies associations and fluvial styles in relation to the flow regime conditions (changes of Froude number) and sediments size (d variations) reflecting river channel pattern transformations (explanation of lithofacial and lithogenetic codes in Table 1) cross-stratified sands (Sp), sandy scour infillings (Se), rip - rippled sands (Sr) and sandy scour infillings (Se) form two- pled sands (Sr) and low-angle cross-stratified sands (Sl). part sequences Sp→Sr and Sf→Se/Sp. Their lower seg- A similar u fl vial succession was recognized at the Paterek ments (lithofacies Sp and Sf) comprise medium-grained and site, where the sandy cosets of mid- and large scale pla- moderately well-sorted sands (median-grain diameters d nar cross-stratification (lithofacies Sp and Sf, respectively), are in the range 0.40–0.54 mm) which are characterized by positively and very positively skewed grain-size distribu- tion. The upper segments consist of small- and medium- scale lithofacies Sp, Se and Sr in which fine-grained sands predominate (median-grain diameters d vary from 0.15 up to 0.40 mm). The upper member of Zielonczyn Fm. at the Paterek site is built of lithofacies association Sh(GSm,Sp). It’s basal part contains massive pebbles with a sandy matrix (lithofa- cies GSm, average median-grain-diameter d = 8.15 mm) (Fig. 9). These sediments are overlain by horizontally stratified fine-grained sands (Sh) which are normally graded (median-grain-diameter d decreases from 0.24 to 0.17 mm) and moderately sorted. These are interbed- ded with small-scale planar cross-stratified and medium- grained sands (lithofacies Sp, average median grain diameter d = 0.41 mm). Fig. 11 Channel depth in relation to water flow velocity and flow regime (Froude number) for Weichselian fluvial lithotypes 1 3 Int J Earth Sci (Geol Rundsch) (2018) 107:1407–1429 1419 Fig. 12 Grain-size compositions of Weichselian fluvial lithotypes in the Toruń Basin shallow channels (depth up to 1 m) and at 8°–14° if channels Fluvial lithotypes, riverbed morphology and channel patterns reached a depth of 1–3.5 m. These low inclined point-bar surfaces developed in channels of 20–30 m wide, the sinuos- The sedimentological results of Weichselian fluvial succes- ity of which varied from 1.25 up to 1.5 (see Leeder 1973). The deposition of fluvial lithotype M1 took place mostly sions in the Toruń Basin allowed for four fluvial lithotypes to be recognized: one of a meandering-style sand-bed river under transitional water flow conditions from subcritical to supercritical or under supercritical flow. These condi- type (M1), and three of braided-style sand-bed river styles (B1, B2 and B3). These lithotypes relate to the sand-bed tions were related especially to the high water flow veloc- −1 ity (up to 1.7 ms ), whereas the Froude number reached rivers of different fluvial style. a maximum only of 0.6 (Fig. 7,11). Similar low values of the Froude number (0.16–0.68) for supercritical flow were Lithotype of a sand‑bed, high‑energy meandering river (M1) observed in modern sand-bed braided and meandering riv- ers (Leopold et al. 1964; Karim 1995; Julien and Raslan The lithotype M1 is represented by lithofacies association 1998; Prent and Hickin 2001). Moreover, the experimental results (Cheel 1990b; Julien and Raslan 1998; Best 2005; Si(Sm,Sr,Se) (Chobielin and Zielonczyn sites; Figs. 7, 10) which are arranged into stacked sequences Sm→Si→Sr(Se). Fourrière et al. 2010) reveal that the supercritical flow may occur in a wide Froude number range 0.21–1.6. These true The indicative large-scale lithofacies of low-angle inclined stratification of fine-grained sands (Si) capped by rippled low values of the Froude number, experimentally obtained for a sequence of which the sedimentology indicates con- sands (Sr) and sandy scours infillings (Se), represent point- bar sequences (Tunbridge 1981, 1984; Stear 1983; Miall ditions of supercritical flow, may reflect the high rate of fine-grained sands suspended in a water column (Bridge 1985, 2006; Olsen 1988, 1989; Julien 2002; Bridge 2003). The surfaces of point-bars sloped at up to 8° in the case of a and Best 1988; Baas and Best 2002; Best 2005). Moreover, 1 3 1420 Int J Earth Sci (Geol Rundsch) (2018) 107:1407–1429 the average median grain diameter (d ) of fluvial lithotype (lithofacies Sh) (Rozwarzyn site and the upper member of M1 is 0.18 mm (Fig. 12). Low-angle cross-stratification the Zielonczyn Fm. at the Paterek site; Figs. 9, 10). These and parallel lamination (inclined or horizontal) commonly lithofacies represent sheet-like beds veneering the floor of a occur in these fine-grained sands, as was reported usually sand-bed-braided river under the condition of supercritical −1 in sands characterized by median grain diameter in the flow (flow velocity up to 1.4 ms ; Figs. 9, 11). The depo- range 0.1–0.3 mm and transported under the condition of sition within shallow streams was related to a wide sheet the Froude number lower than 0.84 (Ashley 1990; South- flow rather than channelised flow (flow depth 0.4–0.7 m) and ard and Boguchwal 1990; Southard 1991). took place at the low Froude number as a consequence of The high transport rate of fine-grained sands of the flu- the high concentration of suspended sediments (Bridge and vial lithotype M1 caused suppression of turbulence in the Best 1988; Cheel 1990a; Julien and Raslan 1998; Best 2005; moving bedload dispersion (see Allen and Leeder 1980). Fourrière et al. 2010). Thus, the association of sediments Thus, the stable upper-plane bed and inclined parallel in suspension with supercritical flow conditions implies stratification of point-bar sequences (lithofacies Si) were abundant horizontally stratified sands (see Bridge and Best formed due to the migration of low-relief bed waves of up 1988; Paola et al. 1989; Southard 1991; Best and Bridge to 20–25 mm high and 1.3 m long (see Bridge and Best 1992; Alexander et al. 2001; Fourrière et al. 2010). These 1988; Paola et al. 1989; Best and Bridge 1992). Similar lithofacies were formed by the migration of antidunes and supercritical water flow conditions were responsible for low-relief bed waves. the deposition of massive sands (lithofacies Sm; Chobielin The upper segment of the lithotype B1 (large-scale and site; Figs. 7, 10) (see Alexander et al. 2001; Duller et al. low angle crossbeds Sl with sigmoidal laminae) reflects 2008) as a result of scour-pool infilling (Gorrell and Shaw lateral aggradation of humpback dunes (Kostaschuk and 1991; Russell et al. 2007) in a river thalweg. Moreover, Villard 1996; Mohrig and Smith 1996; Kostaschuk 2000; two-dimensional dunes migrated and sandy-gravelly upper Winsemann et al. 2011). The morphology of these bedforms flow regime sheets developed in this riverbed section was changeable due to a variable rate of suspended sedi- (lithofacies GSp and SGm, respectively; Zielonczyn site; ment concentration (Kostaschuk and Villard 1996; Mohrig Fig. 7) (Fielding 2006). The rippled sands (Sr) and sandy and Smith 1996; Kostaschuk 2000; Lang and Winsemann chute cut-off infillings (Se) which form the upper segments 2013). Therefore, the deposition of fine-grained sands from of point-bar sequences, are indicators of waning flow con- the suspended load close to the bed controls the lee side of ditions at the end of flood events and reflect episodes of the slope angle (Kostaschuk et al. 2009). Additionally, the sedimentation and erosion (Farrell 1987; Zieliński 1992, migration of humpback dunes that were formed around the 1993; Miall 2006). transition to a supercritical flow could be considered as the The depositional environment of the fluvial lithotype M1 initial phase of the riverbed transformation from upper-plane and its high-energy nature show similarities to the sedimen- beds to the dunes of the lower flow regime (Fielding 2006). tary environment of flashy, ephemeral, sheet flood, sand- Sedimentological features of the fluvial lithotype B1 are bed braided river (Biju Creek type; see McKee et al. 1967; related to the Bijou Creek fluvial lithotype (McKee et al. Tunbridge 1981, 1984; Miall 1985, 2006). However, the 1967; Williams 1971; Tunbridge 1984) and to the model common occurrence of the low-angle inclined stratification of the flashy, ephemeral, sheet flood, sand-bed river (distal, of fine-grained sands (large scale lithofacies Si) indicates sheet flood braided model of Miall 1985). A similar fluvial lateral accretion of point-bars developed in an ephemeral succession was described by Wysota (2002) based on his stream (Stear 1983). Thus, the fluvial lithotype M1 recog- studies of the Rzęczkowo Formation in the area of the Lower nized in the Toruń Basin corresponds to the sedimentary Vistula Valley (in the northern vicinity of the Toruń Basin). environment of an ephemeral sandy meandering river (Miall 2006). Lithotype of a sand‑bed braided river with unit bars (B2) Lithotype of a shallow, high‑energy sand‑bed braided river (B1) The third fluvial lithotype B2 contains lithofacies asso- ciation Sp(Sr,Se,Sl,Sh) which is dominated by planar or Sediments of the fluvial lithotype B1 are similar to the trough cross-stratified sands (lithofacies Sp and St, respec- lithotype M1 in terms the textual features (Fig. 12). These tively; Figs. 9, 10). Typical lithofacies (Sp) are interpreted deposits consist mainly of fine-grained sands (61–62%), as related to the unit transverse bars and numerous sandy while the average value of the median grain diameter is megaripples which covered the bed of a braided river (Smith 0.17–0.18 mm. The key indicator of the lithotype B1 is 1978; Miall 1985; Bridge 1993, 2003). Sedimentation took the dominance of horizontally stratified fine-grained sands place under the transitional conditions between the lower 1 3 Int J Earth Sci (Geol Rundsch) (2018) 107:1407–1429 1421 and upper parts of the lower flow regime (Froude number Fr water flow velocity caused by the decrease of the flow 0.3–0.4; Figs. 9, 10, 11). The surfaces of the unit bars were depth (see Ashley 1990; Southard 1991). Thus, variations dissected by small cross-bar channels with sandy infillings of hydraulic parameters from the lower to upper flow regime (Se) and covered by sands of diminished dunes (Sl). During tended to change the morphology of the compound bar sur- the waning flood stages, fine-grained sands were accumu- faces in abundant areas of sand flats in sand-bed-braided lated due to the tractional deposition under the condition of river. Moreover, the compound bars were dissected by small a ripple-phase bed (lithofacies Sr). cross-bar channels filled with sands (lithofacies Se) during The u fl vial lithotype B2 was deposited from a low-energy the falling flood stage (Bridge and Lunt 2006). flow characterized by a wide range of flow depth (0.3–2.8 m) The secondary lithofacies association Sr(Src,SFh) of the −1 and velocity of up to 1.6 ms (Figs. 9, 11). Its maximum fluvial lithotype B3 represents the overbank sub-environ - value occurred during the deposition of massive gravelly ment where floodplain fines were accumulated as a result sandy sheet-like beds (GSm). The dominance of crossbeds of overbank splay development and sheet flows (Fig. 10) (lithofacies Sp) and minor sequences of laminated sand (Bridge 2003; Bridge and Lunt 2006; Miall 2006). sheets, rippled sands and infillings of cross-bar channels Sediments of the fluvial lithotype B3 were deposited are the common features of a “Plate-type” sand-bed braided predominately under transitional conditions between the river (shallow perennial braided; Miall 1985, 2006). lower and upper parts of the lower flow regime (Fr ≈ 0.3), although, there were transitional conditions from subcriti- Lithotype of a sand‑bed‑braided river with compound cal to supercritical flow in the zones of upper stage plane bars (B3) beds and in cross-bar channels (Figs. 10, 11). The water flow −1 velocity was variable (up to 1.6 ms ). The lithotype B3 contains two lithofacies associations rep- Qualitative features of the sedimentary environment of resenting two fluvial sub-environments (Fig. 10). The main the fluvial lithotype B3 are similar to those in the South association Sf(Sp,Sl,Se,Sr) was deposited within the channel Saskatchewan River (Cant and Walker 1978) and are com- of sand-bed-braided river and is dominated by large-scale parable with the depositional environment of deep perennial planar cross-stratified sands (Sf) formed by the downstream sand-bed braided river (Miall 1985, 1993, 1994, 2006) and migration of macroforms lee side (Nowe Dąbie, Wypaleni- with the lithotype D-3 described by Zieliński (1993) as sedi- ska and Paterek sites; Figs. 8, 9, 10) (Williams 1971; Cant ments of aggrading-braided channels dominated by foreset and Walker 1978; Miall 1978, 1985; Miall and Jones 2003). bar accumulation. This type of lithofacies predominantly occurred at the depths of 1.5–2.5 m below the surface of compound bars, which corresponds to those recognized in the South Saskatchewan The relations between fluvial deposits River (Cant and Walker 1978; Sambrook Smith et al. 2006; and hydraulic conditions of the bedform Ashworth et al. 2011). These macroforms develop in a sand- development bed-braided river due to the unit transverse bar accretion and were also observed in the Platte (Blodgett and Stanley Hydraulic parameters of the rivers (e.g., water flow veloc- 1980; Allen 1983; Crowley 1983; Bridge 1993, 2003), Vis- ity) affect the development of bedforms, to which fluvial tula (Babiński 1992), Yamuna (Best et al. 2003) and Niob- lithofacies correspond. Due to this the relationships between rara (Skelly et al. 2003) rivers. Moreover, lithofacies Sf may properties of fluvial sediments (structural and textural) and well be interpreted as an effect of alternate bar migration flow competence reflect hydrodynamic conditions of the flu- (McCabe 1977; Crowley 1983), the crest of which was dis- vial depositional environment. For this reason, quantitative sected by chute channels (lithofacies Se). Alternate bars may interpretation of sedimentary structures and textural proper- be subsequently subjected to transformation into midchannel ties of fluvial lithotypes is fundamental. compound bars (see Bridge 1985, 2003; Bridge et al. 1986; Lithotypes M1 and B1 are characterized by the most sim- Babiński 1992). ilar grain-size distributions (Fig. 12) which are generally The upper part of the compound bar succession was symmetrical or—locally—show a negative skewness asso- accumulated during the flood as a result of two-dimensional ciated with an admixture of fractions coarser than average. dunes and ripple migration (lithofacies Sp, Sr, respectively) Fine-grained sands represent 61–62% on average, while the which constituted sand shoals on the compound bar surface average values of the median grain size distributions are 0.17 (Sambrook Smith et al. 2006; Ashworth et al. 2000, 2011). and 0.18 mm in the case of lithotypes M1 and B1, respec- These bedforms may change rapidly into the diminished tively. Although these lithotypes represent different riverbed dunes (lithofacies Sl) or into upper plane beds (lithofacies patterns, their riverbed was in each stage modeled under Sh) due to the predominance of fine-grained sands trans- supercritical flow regime or in transitional circumstances ported within the stream under the condition of increasing towards such flow conditions (Fig. 10). 1 3 1422 Int J Earth Sci (Geol Rundsch) (2018) 107:1407–1429 Fig. 13 Types of channel deposits in relation to the flow velocity, channel bedforms and Weichselian fluvial lithotypes Unit bars and dunes, characteristic of the lithotype B2, are composed mainly of sands with a different gravel admixtures (Fig. 12). As a result, these sediments vary from fine- and medium-grained sands (characterized by symmetrical grain- size distributions) to coarse- and very coarse-grained sands, or even gravels (very positive and very negative skewed) (Fig. 12). The average value of median grain diameters is 0.45 mm. The lithofacies of fluvial lithotype B3 comprise texturally homogeneous sands which contain of about 52% medium- grained sands. The average value of median grain diameters of these deposits is 0.26 mm (Fig. 12). These deposits are better sorted than sands of the lithotype B2, show symmetri- cal grain-size distributions and in few cases slightly positive or negative skewness. Sediments of the lithotype B3 have a negligible contents of coarse-grained sands and gravels and for these reasons, the average median grain-size diameter is lower than for the lithotype B2 and is 0.26 mm (Fig. 12). These data indicate greater stability of hydrodynamic condi- tions of the depositional environment of the lithotype B3 in comparison with the lithotype B2. An analysis of the relationship between texture charac- teristics of the Weichselian fluvial successions in relation to the conditions for the development of the river bedforms, allowed for the delimitation of the three types of fluvial deposits. The first type (“I” in Fig. 13) is represented by the sand-bed-braided river with bar-related sedimentation (“b” Fig. 14 The model of river channel transformation during the Weich- in Fig. 13), transported mainly under condition of lower flow selian glaciation as an effect of climate oscillation 1 3 Int J Earth Sci (Geol Rundsch) (2018) 107:1407–1429 1423 regime (lithotypes B2 and B3). The sorting of these deposits becomes worse for sediments characterized by median grain diameter (d ) in the range of 0.25–0.5 mm. The second type of the river channel deposits (“II” in Fig. 13) contains sands in which sorting is more poor with the decrease in median grain diameter from 0.25 to 0.15 mm. Moreover, these sedi- ments represent gradual changes from the lithotypes related to sand-bed-braided rivers with unit and compound bars (“b” in Fig. 13) to the lithotypes of meandering rivers and shal- low braided rivers with upper plane bed (“a” in Fig. 13), and hence changeable conditions from lower to the upper flow regime. The third type of the channel sediments has an average grain-diameter (d ) lower than 0.15 mm (“III” in Fig. 13). These sands represent high-energy, sand-bed and shallow- braided rivers (lithotype B1) or ephemeral meandering rivers (lithotype M1). The sorting of these deposits gets worse with the increase in size of median grain diameter. As is apparent from the relationship between the median grain diameter (d ) and the velocity of water flow (V ), the same fraction of the deposits may be transported in differ - ent hydraulic conditions (see Ashley 1990; Southard and Boguchwal 1990; Southard 1991). This applies mainly to the deposits of an average diameter (d ) smaller than 2.5 mm (types “II” and “III” in Fig. 13). Regardless of this, the median grain diameter of 0.15 mm may be the bound- ary between the channel sediments transported primarily in the conditions of the upper flow regime (type III of flu- vial deposits) and deposits of sand-bed-braided rivers with transverse bars and dunes (types “I” and “II”; Fig. 13) (see Fig. 15 The relation between lithostratigraphic units and Weich- Sly et al. 1983). Nevertheless, during the transportation selian lithotype succession in the Toruń Basin [M—litothotype of of fine-grained sands (d lower than 0.15 mm), there is a eemian meandering river after Makowska (1980); SS—lithotype of potential for ripples or an upper plane bed to develop (type proglacial ice-dammed lake after Wysota (2002); M1, B1–B3—flu- III of channel deposits; Fig. 13), which refers flow compe- vial litothotypes deposited during Weichselian glaciation (analyzed in this study); B4—fluvial lithotype of high-energy-braided river tence to sediments transporting under conditions of different with sandy gravelly three-dimensional dunes after Weckwerth (2010, (lower or upper) flow regimes (see Ashley 1990; Southard 2013); p1–p3—pradolinas 1–3] and Boguchwal 1990; Southard 1991). For these reasons, the textual criterions (d = 0.25 and d = 0.15 mm) cannot 50 50 be unambiguous in hydrodynamic interpretation of fluvial deposits. M1/B1 in Fig. 14) were characterized by channel depth up to 1 m, a high Froude number (greater than 0.4) and high −1 flow velocity (greater than 0.9 ms ). Apart from this, condi- Climate‑derived transformation of riverbed tions of shallow water flow in meandering channels (litho- morphology and flow conditions type M1) become similar to the conditions characteristic during Weichselian glaciation of shallow-braided rivers with the dominance of the upper plane bed (lithotype B1; cf. Dreyer 1993). Moreover, the Meandering and shallow‑braided rivers increasing flow depth in meandering rivers, together with increasing their sinuosity index (SN from 1.25 up to 1.5), Lithotypes of the high-energy rivers, forming the lower were responsible for decreasing the Froude number from members of the Weichselian fluvial formations, were depos- 0.5–0.6 to 0.3–0.4 (Figs. 10, 11, 14). Increasing median ited as a result of the activities of precipitation-dependent grain diameter of the channel deposits in the case of the rivers (Weckwerth and Chabowski 2013). Typical water flow lithotype M1 (e.g., fluvial succession at the Chobielin conditions of meandering and shallow-braided rivers (area site) reflects cumulative effects over successive cycles of 1 3 1424 Int J Earth Sci (Geol Rundsch) (2018) 107:1407–1429 climate changes (cf. Busschers et al. 2007; Van Balen et al. shallow-braided- (lithotype B1) into the sand-bed-braided 2010), between Dansgaard-Oeschger cycles D/O8 and D/ river dominated by unit bars at the beginning of H3 and O5 (Rahmstorf 2003). In turn, this may have resulted in the H2 events (cycles M1→B2 and B1→B2, respectively; seasonal increase in the channel aggradation rate of coarser Figs. 14, 15). These transformations were manifested by sediments (Van Huissteden and Kasse 2001; Van Huissteden decreasing flow velocity (settled the low Froude num- et al. 2001). Similar periodic floods are typical of modern ber), linked to the energy losses associated with bedform rivers in the polar and sub-polar regions (Vandenberghe roughness and with the transportation of coarser particles 2001). supplied to the Toruń Basin by sediment-laden meltwater Based on the thickness of the lithotypes M1 and B1 in from the north (Figs. 10, 11, 12, 13, 14). relation to the time of their accumulation in the TSB, it can The beginning of the second stage of the ice-marginal be assumed that the rate of channel deposition occurring valley development (during H2 event, pradolina 2) coin- between 28 and 24 ka ago (between D/O4 and D/O2) was cided with unit bar amalgamation which resulted in the higher than 40–30 ka ago (between D/O8 and D/O5). Its development of compound bars typical for the lithotype B3 cause, despite the lack of meltwater feeding, may have been (Figs. 14, 15). These macroforms constituted sand shoals an increase in the efficiency of surface runoff affected by within braided river channels. The water flows showed a the expansion of continuous permafrost (Renssen and Van- similar depth range during the deposition of the lithotypes denberghe 2003; Buschers et al. 2005, 2007). These climate B2 and B3, although the flow velocity was generally greater conditions in the area of the Toruń Basin is proven by wedge in rivers of lithotype B3 (Figs. 11, 14), for which the energy casts filled with well-sorted sands at 27 ka ago (Drozdowski of sedimentary environment was more stable than during and Fedorowicz 1987; Drozdowski and Szupryczyński the deposition of lithotype B2. Therefore, the lithotype B2 1987). represents a transitional phase in the evolution of the sand- Repetitive climate cooling at the beginning of H3 event bed-braided river morphology characterized by two- and (~ 30 ka ago) and H2 event (~ 24 ka ago) (Bond et al. 1999; three-dimensional dunes and unit bar development. Hemming 2004; Peck et al. 2007) caused changes in the Deposition of lithotypes B2 and B3 (which was related paleoenvironmental conditions of river catchments. The to the increase of the channel aggradation rate) took place hydraulic parameters of meandering rivers (lithotype M1) 1–2 thousand years before every ice sheet advance into the and shallow-braided rivers (lithotype B1) got similar to the Toruń Basin, during the development of pradolinas 1 and 2 conditions enabling threshold changes in the riverbed mor- (Figs. 14, 15). Fluvial lithotypes succession (B2→B3) was phology (a.o. Schumm 1979). They consisted in decreas- deposited during the middle Weichselian and indicates the ing the water flow velocity, the reduction of both hydraulic relatively stable conditions of water flow only before the gradient and the Froude number, as well as in increasing the LGM (area B3 in Figs. 14, 15). Therefore, the development diameter of transported particles and the channel aggrada- of compound bars is a mature and finale stage of the sand- tion rate (Figs. 13, 14). The consequence was a river pat- bed-braided river evolution in fluvial cycle. Fluvial succes- tern transformation from meandering and shallow-braided- sions accumulated in the Toruń-Eberswalde ice-marginal (lithotypes M1 and B1) to a braided river style with unit valley (pradolina 3) during Pomeranian phase indicate bars (lithotype B2; Figs. 14, 15). This process was initiated braided river environment which varied under the terms of by humpback and two-dimensional dune development (e.g., morphology and flow energy (Fig. 15; Weckwerth 2010, Rozwarzyn site; Fig. 8). 2013; Pisarska-Jamroży 2015). River channel pattern transformations which correspond Braided rivers with unit and compound bars to the fluvial lithotype sequences M1/B1→ B2 and B2→B3, occurred 30 and 24 ka ago in the area of the Toruń Basin. The lithotypes B2 and B3 form the upper members of Fluvial transformation proceeded in Central Poland and the Weichselian fluvial formations. These lithotypes are in the Southern North-Sea Basin at roughly the same time related to the streams which were fed by meltwaters flow - (Rotnicki 1987; Turkowska 1995; Huisink 1997, 2000; ing from the north and precipitation-dependent rivers Mol 1997; Buschers et al. 2005, 2007; Cordier et al. 2006; flowing from the south (Weckwerth and Chabowski 2013). Zieliński 2007). Their causes were related to the glacio-isos- Thus, lithotypes B2 and B3 represent two stages of the ice- tacy and climate cooling, which resulted in the development marginal valley development (pradolinas 1 and 2; Fig. 15) of continuous permafrost, in conditions of which increased during H3 and H2 events (pradolinas 1 and 2, respectively) the sediment supply into rivers in coincidence with particle which differ in relation to the lithological features. coarseness in fluvial succession (cf. Huijzer and Vanden - The first stage began with the deposition of the litho- berghe 1998; Renssen and Vandenberghe 2003; Van Huisst- type B2. This corresponded to the transformations of the eden et al. 2001). Moreover, one of the main reasons for the river channel pattern from meandering (lithotype M1) and transformation of the river channel morphology in the Toruń 1 3 Int J Earth Sci (Geol Rundsch) (2018) 107:1407–1429 1425 creativecommons.org/licenses/by/4.0/), which permits unrestricted use, Basin was river regime changes from precipitation-related to distribution, and reproduction in any medium, provided you give appro- glacial which was associated with the two ice-marginal val- priate credit to the original author(s) and the source, provide a link to ley development before LGM (Weckwerth and Chabowski the Creative Commons license, and indicate if changes were made. 2013). In these valleys, the high rate of fluvial accumulation (deposition of lithotypes B2 and B3) was dependent on the significant sediment load in meltwater streams which fed the Toruń Basin. 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International Journal of Earth Sciences – Springer Journals
Published: Oct 23, 2017
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