Sedimentary architecture and depositional controls of a ... - Pride

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Sedimentary architecture and depositional

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controls of a Pliocene river-dominated delta

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in the semi-isolated Dacian Basin, Black Sea

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Elisabeth L. Jorissen a*, Arjan de Leeuw b,c, Christiaan G.C.

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van Baak a,c, Oleg Mandic d, Marius Stoica e, Hemmo A.

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Abels f, Wout Krijgsman a

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a Palaeomagnetic Laboratory ‘Fort Hoofddijk’, Faculty of Geosciences, Utrecht University, Budapestlaan 17, 3584 CD, Utrecht, The Netherlands, [email protected]; [email protected] b Université Grenoble Alpes, Institut des Sciences de la Terre, 38000 Grenoble, France, [email protected] c CASP, West Building, Madingley Rise, Madingley Road, Cambridge, CB3 0UD, United Kingdom, [email protected] d Geological-palaeontological Department, Natural History Museum Vienna, Burgring 7, 1010 Vienna, Austria, [email protected] e Department of Palaeontology, Faculty of Geology and Geophysics, University of Bucharest, Bălcescu Bd. 1, 010041, Romania, [email protected] f Department of Geosciences and Engineering, Delft University of Technology, Stevinweg 1, 2628 CN, Delft, The Netherlands, [email protected]

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Abstract Sedimentological facies models for (semi-)isolated basins are

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less well developed than those for marine environments, but are

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critical for our understanding of both present-day and ancient deltaic

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sediment records in restricted depositional environments. This study

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considers an 835 m thick sedimentary succession of mid-Pliocene

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age, which accumulated in the Dacian Basin, a former embayment of

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the Black Sea. Detailed sedimentological and palaeontological

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analyses reveal a regression from distal prodelta deposits with

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brackish water faunas to delta-top deposits with freshwater faunas.

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Sediments contain frequent hyperpycnal plumes and an enrichment in

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terrestrial organic material, ichnofossils and in situ brackish and

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freshwater faunas. Deltaic progradation created thin, sharply-based

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sand bodies formed by multiple terminal distributary channels,

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covering a wide depositional area. The system experienced frequent

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delta-lobe switching, resulting in numerous thin parasequences.

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Parasequences are overlain by erosive reddish oxidized sand beds,

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enriched in broken, abraded brackish and freshwater shells. These

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beds were formed after sediment starvation, on top of abandoned

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delta lobes during each flooding event. A robust

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magnetostratigraphic time frame allowed for comparison between the

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observed sedimentary cyclicity and the amplitude and frequency of

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astronomical forcing cycles. Our results indicate that parasequence

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frequencies are significantly higher than the number of time

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equivalent astronomical cycles. This suggests that delta-lobe

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switching was due to autogenic processes. We consider the observed

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facies architecture typical for a delta prograding on a low-gradient

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slope into a shallow, brackish, protected, semi-isolated basin.

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Furthermore, in the absence of significant wave and tidal influence,

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sediment progradation in such a protected depositional setting shaped

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a delta, strongly river-dominated.

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Key-words: Paratethys, isolated basin, river-dominated delta,

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regressive parasequences, autogenic forcing, flooding surface

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1. Introduction

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Restricted sedimentary basins constitute complex depositional

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environments. They form semi-isolated basins when they display

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limited water connections to the marine realm, or isolated basins

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when all connections are interrupted. The sedimentary infill of (semi-

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)isolated basins is mostly controlled by the relative importance of

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accommodation space, water supply and sediment supply (Bohacs et

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al., 2003). Restricted sedimentary basins record generally limited

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accommodation space compared to the open ocean (Carroll and

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Bohacs, 1999). An important control on the limited accommodation

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space is the position and relative height of the spill point of the basins

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(Leever et al., 2010; Yanina, 2014; Fongngern et al., 2016), which

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may at the same time affect their connectivity to the marine realm

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(Leever et al., 2011; Ter Borgh et al., 2014). Restricted connectivity

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with the open ocean results in the prevalence of brackish to

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freshwater depositional environments and is usually accompanied by

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the development of faunas endemic to the basin (e.g., Jones and

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Simmons, 1996; Rögl, 1998; Wesselingh et al., 2006). Limited water

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interchanges between (semi-)isolated basins and the open ocean lead

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moreover to low energy depositional environments. Therefore, (semi-

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)isolated basins generally lack tidal influence and often show reduced

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wave interference (Medvedev et al., 2016). The isolated nature of

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restricted basins also enhances their sensitivity to external forcing

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factors, such as changes in precipitation, runoff and (cyclic)

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astronomical controls (e.g., Müller et al., 2001; Abels et al., 2009;

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Litt and Anselmetti, 2014; Wagner et al., 2014; Constantinescu et al.,

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2015). If properly understood, (semi-)isolated sedimentary basins

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may provide valuable environmental, climatic and biological

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archives. Understanding depositional processes operating in these

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basins may provide spatial and temporal insights into the distribution

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of sedimentary facies and fauna, and could elucidate potential climate

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forcing.

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Although the understanding of (semi-)isolated basins has

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significantly increased in the past decades through various deep-

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drilling projects (e.g., Ross, 1978; Francke et al., 2016), depositional

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processes occurring in restricted basins and their impacts on

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sedimentary facies are still not well known. The sedimentary infill of

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restricted basins is likely controlled by forcing factors that are

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different to the open ocean. When (semi-)isolated basins are infilled

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by prograding deltas, the resulting deltaic architectures are likely

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influenced by both external and internal forcing factors specific to

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restricted depositional settings. The impact of these forcing factors

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may however be difficult to identify in deltaic sedimentary records.

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In particular, the relative importance of allogenic forcing factors

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driving accommodation space, water supply and sediment supply

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acting of the deltaic progradation compared to autogenic deltaic

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avulsion processes remain poorly studied in restricted basins.

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As a result, deltaic facies models for (semi-)isolated basins are

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less well developed than those for open marine environments

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(Andrews et al., 2016; Nutz et al., 2017). This is remarkable as some

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of the present-day (semi-)isolated basins are infilled by major deltas.

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The restricted Black Sea and isolated Caspian Sea accommodate for

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instance the Danube and Volga deltas, the two longest rivers in

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Europe (Overeem et al., 2003; Giosan et al., 2005). Unfortunately,

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continuous exposures recording long periods of deltaic deposition in

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(semi-)isolated basins are relatively rare.

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In this study, we investigate the mid-Pliocene sedimentary

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architecture of a river-dominated delta entering the semi-isolated

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Dacian Basin in Romania. The Dacian Basin formed at that time a

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brackish embayment of the ancient Black Sea. Deltaic and alluvial

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sediments prograded on the northern margin of this restricted basin.

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We investigated an 835 m thick, continuous sedimentary section of

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fossil-rich sediments, cropping out along the Slănicul de Buzău

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River, in Romania (Andreescu et al., 2011; Van Baak et al., 2015). In

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this article we combine detailed analyses of sedimentary facies, with

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a description of the accompanying biofacies. The quality and

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continuity of the exposure allows for establishing a detailed

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sedimentological and sequence-stratigraphic framework, which

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permits investigating the drivers of the internal deltaic architecture.

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Moreover, because of available magnetostratigraphic time constraints

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(Van Baak et al., 2015), the impact of autogenic versus allogenic

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forcing factors on the deltaic sedimentary architecture can be

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discussed. Facies models developed in this paper may form

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analogues for more poorly exposed or subsurface deltaic successions

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in (semi-)isolated basin elsewhere.

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2. Geological background

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The Paratethys Sea formed one of the largest intercontinental

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seas that ever existed (Rögl, 1998; Popov et al., 2006). During the

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Oligocene, convergence between Africa and Eurasia generated a

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topographical barrier, which isolated the Paratethys Sea from the

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Tethys Ocean (Allen and Armstrong, 2008; Schmid et al., 2008).

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Oligocene and Miocene tectonic activity produced numerous

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mountain belts (Vincent et al., 2007, 2016; Schmid et al., 2008),

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which further fragmented the Paratethys Sea into several semi-

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isolated basins (Popov et al., 2006). From west to east, the four major

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sub-basins were the Pannonian, Dacian, Euxinian (Black Sea) and

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Caspian basins (Fig. 1a).

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This paper focuses on the semi-isolated Dacian Basin, a former

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embayment of the Black Sea. This basin represents the late Miocene

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to present-day foreland basin in the eastern and southern parts of the

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Carpathians (Matenco and Bertotti, 2000; Cloetingh et al., 2004;

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Panaiotu et al., 2007; Jipa, 2015). Following mountain belt uplift, a

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deep depocentre was formed in front of the Southeast Carpathians

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(Bertotti et al., 2003; Tărăpoancă et al., 2003). The depression was

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progressively filled with the erosional products of the uplifted

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mountains (Jipa, 1997; Sanders et al., 1999; Tărăpoancă et al., 2003;

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Panaiotu et al., 2007) (Fig. 1b). Open water deposits with brackish

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water faunas, which accumulated in the basin during the late

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Miocene-early Pliocene (Pontian regional stage - Stoica et al., 2013),

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were gradually replaced by alluvial deposits with freshwater faunas

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towards the late Pliocene (Romanian regional stage - Van Baak et al.,

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2015). This transition of depositional environments occurred during

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the intermediate Dacian regional stage, which lasted from 4.8 to 4.2

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Ma (Vasiliev et al., 2005; Van Baak et al., 2015) (Fig. 2a). At that

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time, a major delta prograded east of the Carpathians towards the

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northeastern margin of the Dacian Basin (Jipa, 1997; Jipa and Olariu,

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2009; Fongngern et al., 2016; Matoshko et al., 2016), whereas

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sediments shed from the Southern Carpathians mainly accumulated 6

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in the western Dacian Basin (Jipa and Olariu, 2009; Jipa et al., 2011;

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Ter Borgh et al., 2014; Fongngern et al., 2017). Palaeogeographic

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and provenance data indicate that the basin was eventually entirely

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filled during the late Pliocene to early Pleistocene, when sediments

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started to overspill into the Black Sea (De Leeuw et al., 2018; Olariu

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et al., 2018).

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Post-collisional shortening affected the Carpathian Foredeep

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during the Quaternary (Necea et al., 2005; Leever et al., 2006;

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Maynard et al., 2012). Faulting and large-scale folding of the

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foreland infill occurred in the Buzău area of Romania. As a result of

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this Quaternary foreland inversion, long and continuous exposures of

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the late Miocene to early Pleistocene foreland infill can be found at

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the surface.

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3. The Slănicul de Buzău section

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The investigated section crops out along the Slănicul de Buzău

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River, between the villages Cernătești and Minzălești (Fig. 2b). The

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river cuts through a 6.4 km thick stratigraphic succession, recording

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folded foreland late Miocene to Pleistocene deposits (Snel et al.,

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2006; Andreescu et al., 2011; Van Baak et al., 2015).

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Our study focuses on the part of the section corresponding to the

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mid-Pliocene Dacian regional stage. The existing age model for this

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part of the valley (Van Baak et al., 2015) shows that this section

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contains two normal magnetozones, interpreted as paleomagnetic

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chrons C3n.2n (Nunivak), aged 4.631-4.493 Ma and C3n.1n

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(Cochiti), aged 4.300-4.187 Ma (absolute ages from Gradstein et al.,

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2012). Including the under- and overlying sediments, the entire

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studied section was therefore deposited between 4.8 Ma and 4.11 Ma.

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These age constraints are in line with studies of the Dacian regional

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stage at other locations in the Dacian Basin (e.g., Vasiliev et al.,

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2004).

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4. Methods

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4.1. Sedimentological data collecting

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The Dacian segment of the Slănicul de Buzău section starts

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under the bridge North of the village Niculești (N45°26’26.87’’,

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E26°44’37.97’’) and continues for 2 km northwards in the river bed

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(N45°27’14.81, E26°44’41.29’’) (Fig. 2c). The section exposes an

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835 m thick stratigraphic succession. Thicknesses were measured in

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the field and later checked by GPS measurements.

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This part of the section had previously been logged at a m-scale

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(Van Baak et al., 2015). This more generalized log unfortunately

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missed a 40 m interval in the middle part of the section, which is now

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included in our work (Fig. 3a). To support our sedimentological

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analyses, the section was analysed in greater detail and several key

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intervals were described at a cm-scale. Variations in lithology, grain

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size and sedimentary structures were recorded in the field. Particular

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attention was paid to sedimentary structures, such as graded bedding,

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laminations, cross-stratification and ichnofossils. Samples were

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collected for sedimentological and petrographic optical microscopic

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descriptions. Thin-sections with a thickness of 30 µm were made

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perpendicular to the sedimentary structures for petrographic

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descriptions.

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Detailed sedimentological observations allowed for several

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typical lithofacies to be established, based on sediment grain size,

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sedimentary structures, ichnofossils and faunal composition.

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Lithofacies repeatedly occurring together along the section were then

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grouped into facies associations, each of which related to a distinct

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depositional environment and an estimated water depth. A depth

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ranking scale was subsequently constructed by attributing a number

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from 0 to 9 to each facies associations, with 0 being the deepest and 9

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the shallowest depositional environment (Table 1). This ranking scale

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permitted the reconstruction of a relative water-level curve and the

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identification of parasequences, including superimposed lower- and

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higher-order sequences (Fig. 3b).

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4.2. Palaeocurrent determination

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Along the sedimentary succession, 41 palaeocurrent directions

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were measured on 3D cross-beds. As this section was affected by

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post-depositional folding, palaeocurrent directions needed correction

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to remove any tectonically-induced rotations (Supplementary

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material 1). Corrections were realized with the help of the available

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palaeomagnetic dataset. We proceeded with a first step of deplunging

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the fold axis, followed by a second step of correcting for the true

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vertical axis rotation.

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To correct for the plunging fold axis, all obtained

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palaeomagnetic directions, bedding planes and their poles were

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plotted in a stereographic projection using Stereonet 9 software

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(Allmendinger et al., 2011; Cardozo and Allmendinger, 2013).

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Bedding planes and palaeomagnetic directions were subsequently

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rotated 19° around a rotation axis with a 121° azimuth and 0° plunge

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to restore the fold axis to horizontal. The corrected palaeomagnetic

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directions were then entered as pre-tilt directions in the statistics

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portal of palaeomagnetism.org (Koymans et al., 2016), together with

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their associated plunge-corrected bedding planes.

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For the second step of unfolding, regular tilt-correction on the

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basis of these bedding planes was applied to place palaeomagnetic

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directions in their correct tectonic reference frame. Subsequent

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regular statistical analysis revealed a plunge-corrected anticline with

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a mean direction of 171°. This implies a 9° counter clockwise

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horizontal plane rotation of the section. These results are in line with

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the previously determined 14° rotation (Slănicul site of Dupont-Nivet

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et al., 2005; Vasiliev et al., 2009; Van Baak et al., 2015). They

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illustrate tectonic rotation in the Buzău area, due to the deformation

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of the Carpathian Bend zone.

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Once the palaeocurrent directions were corrected, their

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statistical distribution was calculated for 12 sectors of 30° and plotted

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on four rose diagrams with a maximum representability of 50%. One

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of the rose diagrams represents the overall flow direction along the

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section. Palaeocurrent directions were additionally plotted on three

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rose diagrams according to their respective depositional

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environments.

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4.3. Cyclostratigraphical analysis

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An age model was constructed using the existing

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magnetostratigraphic timescale of the studied section. For each

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magnetozone, the sedimentation rate was calculated in order to

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evaluate variations of sediment input into the basin through time.

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A cyclostratigraphical analysis was performed on the

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frequencies of the parasequences, low- and high-order sequences, to

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evaluate potential astronomical forcing on sedimentation. Blackman-

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Tuckey spectral analyses were realized using standard settings on an

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equally-spaced data series in the Analyseries 2.0.4b program.

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Analyses were performed at 90% confidence levels. Bandpass filters

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were generated with a bandpass-width defined arbitrary between 50

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and 138 m, 64 and 111 m, 25 and 37 m and 13.6 and 25 m. Filters

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were then plotted against the facies rank data and the astronomical

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target curves.

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4.4. Faunal analyses

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Working in Paratethyan basins may introduce a certain

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ambiguity between palaeontological and sedimentological

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terminology (Matoshko et al., 2016). These basins registered episodic

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periods of connectivity and disconnectivity with the open ocean and

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therefore display lowered salinity environments, where endemic

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faunas developed though time (Marinescu, 1978; Stoica et al., 2013).

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Here, the terms ‘brackish water’ and ‘freshwater’ are used to specify

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water salinity on the basis of palaeontological indicators. The term

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‘open water’ is used in a sedimentological context in order to

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describe offshore to shoreface depositional environments.

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The Slănicul de Buzău section display very rich assemblages of

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endemic mollusc and ostracod faunas. For this study, both groups

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were analysed in order to corroborate environmental reconstructions

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based on sedimentology. Mollusc assemblages were studied from 58

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sediment samples. Samples of 1000-2000 g were taken throughout

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the section (Fig. 3a, SBD16-nF). Sample preparation for molluscs

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was performed at the Natural History Museum in Vienna. Samples

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were cleaned using pneumatic micro-chisels. They were then washed

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through sieves of 1 mm. The general preservation of the shells was

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moderate to poor. Shells were finely cracked due to secondary

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gypsum mineralisation and carbonate crystal growth. The taxonomic

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identifications follow Wenz (1942) and Marinescu and Papaianopol

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(1995). Taxonomic revision incorporates results by Nevesskaya et al.

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(1997, 2001, 2013) and Neubauer et al. (2014) (Supplementary

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material 2).

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Ostracod assemblages were studied from 35 sediment samples

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of 500-1000 g taken throughout the section (Fig. 3a, SBD16-nO).

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Sample preparation for ostracods was carried out at the Department

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of Palaeontology at the University of Bucharest. Samples were dried

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to remove interstitial water from the sediments. Dry samples were

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subsequently boiled for 30-60 minutes with a sodium carbonate

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solution for better disaggregation. Samples were washed through

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several sieves of 63 to 500 μm. The residues were studied under a

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ZEISS–Stemi SV11 microscope. Pictures of microfaunas were taken

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with a NIKON digital camera. The general preservation of the

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ostracods was moderate to poor. Ostracods were often fragmented

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due to strong diagenetic processes. The taxonomic identifications

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follow Hanganu (1976, 1985), Hanganu and Papaianopol (1977),

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Stancheva (1990) and Olteanu (1995) (Supplementary material 2).

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5. Results

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5.1. Sedimentary facies associations

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The 835 m thick succession displays a generally regressive

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trend, superposed by a rhythmic alternation between more distal

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clays and more proximal sands. Our field observations and

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subsequent microscope descriptions were compared to well

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documented sedimentological classifications (Postma, 1990; Miall,

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2006). We distinguished thirteen lithofacies, formed by distinct

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sedimentary processes (Table 2, Figs. 4-7). Lithofacies were grouped

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into eight depositional facies, representing five main facies

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associations, each of which being related to a distinctive depositional

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environment.

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5.1.1. Prodelta facies association

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Description

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The first facies association is generally 1 to 5 m thick. It

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consists of three types of dark-bluish-gray (GLEY2-4/5B) to bluish-

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gray (GLEY2-5/5B) mudstone that occur successively. There are

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massive (Fm), laminated (Fl) and lenticular (Fs) mudstones. Massive

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mudstones (Fm) occur at the base of the prodelta facies association

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strata. They are generally 0.5 to 1 m thick, but sometimes are absent

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from this facies association. They display a dark-bluish-gray color

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(Fig. 4b). They may contain well-preserved in situ brackish water

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molluscs, such as Euxinicardium olivetum, Pontalmyra tohanensis or

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Chartoconcha rumana (Supplementary material 3, Fig. 8). This

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facies is progressively replaced by 1 to 3 m thick, laminated gray

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mudstones (Fl). These mudstones have mm-scale horizontal

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laminations of silt (Fig. 4c). They may also contain cm-scale

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horizontal laminations of silt with mm-scale fragments of terrestrial

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organic material. Upwards, the muddy succession may contain 0.5 to

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1 m thick, gray mudstones with lenticular bedding (Fs). The

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lenticular bedding consists of 1 to 5 cm thick isolated lenses made of

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silt to very fine sand showing trough cross-stratification (Fig. 4d).

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The sandy layers are occasionally affected by cm-scale convolute

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bedding. Throughout this facies association, the laminations and the

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lenses become thicker, more frequent, and composed of coarser

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sediments toward the top. Palaeocurrent directions were measured on

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3D trough-cross stratifications present in lenticular-bedding.

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Measurements in these deposits demonstrate a mean direction of

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225° (n=17; Fig. 7a). They display a wide range of current directions

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from 180° to 270°. In addition, deposits occasionally show

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intercalations of cm-thick beds of gray (GLEY1-6/N) fine to medium

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grained sandstones (Sfr). Their bases form a wavy, sharp surface,

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highly perturbed by vertical burrows 3 to 5 cm wide and 5 to 15 cm

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deep (Fig. 5e). These sandstones are moderately-sorted and are

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composed of subangular quartz grains with high sphericity. They

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contain many abraded or broken, reworked brackish and freshwater

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cardiids, unionids, dreissenids or viviparids.

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Interpretation

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This facies association is interpreted as a prodelta environment

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because of the distal depositional setting and the evidence of distal

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fluvial input. The massive mudstones (Fm) indicate deposition from

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suspension in open water. The progressive transition to mm-scale

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silty laminations (Fl) is related to large fluvial outflows, energetic

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enough to reach the distal part of the basin. The cm-scale, organic-

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rich, silty laminations could be related to hyperpycnal flows, 14

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associated with episodic larger river discharge events (Mulder et al.,

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2003; Bhattacharya and MacEachern, 2009; Lamb and Mohrig,

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2009). Upwards, the gradual occurrence of silty lenticular bedding

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(Fs) is related to wave action and/or winnowing (De Raaf et al.,

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1977). The sandy beds comprising reworked, abraded and broken

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shells (Sfg) are thought to illustrate sporadic higher energetic

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depositional processes occurring in the muddy surrounding

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environment. Coarser structureless sediments were transported into

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the basin over long distances during intermittent sand influxes. As in

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previous studies (Starek et al., 2010; Hampson et al., 2011), these

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sandstones are interpreted as storm deposits. The coarsening

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character within this facies association illustrates an increase in the

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energy of the depositional process and is seen as a shallowing of the

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environment. Following earlier studies of similar muddy facies

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associations (Overeem et al., 2003; Olariu and Bhattacharya, 2006;

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Fielding, 2010), we propose that sediments were deposited in a

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prodelta setting. The large range of palaeocurrent directions

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highlights the development of several delta-lobes, feeding a wide

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prodelta region.

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5.1.2. Distal delta-front facies association

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Description

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The second facies association shows a 0.5 to 5 m thick, regular

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alternation between cm- to dm-thick layers of mudstones and

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sandstones. The mudstones are blueish-gray (GLEY2-5/5B). They

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record mm- to cm-scale horizontal laminations of silts or very fine

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sands draped by silts and fragments of terrestrial organic material

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(Fl), as well as lenticular bedding made of cm-scale lenses with

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trough cross-stratified silt to very fine sand (Fs). They are 15

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intercalated with layers of grayish-brown sands (2.5Y-5/2). These

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sands are very fine to fine grained, moderately sorted, and contain

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low-spherical and subangular grains. The sandstones show three

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types of cm-scale cross-beddings (Sr, Ss, St). They comprise 10 to 50

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cm thick climbing ripples (Fig. 5b), 20 to 100 cm thick sigmoidal

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cross-stratification (Fig. 5c) and 10 to 50 cm thick trough cross-

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stratification (Fig. 5d). The cross-bedding foresets are commonly

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draped by mm-scale laminae of fragments of terrestrial organic

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matter. The sandstones frequently show convolute bedding on a cm-

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to dm-scale. The sandy beds become thicker, more frequent, and are

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composed of coarser sediments toward the top of the facies

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association. As in the previous prodelta facies association, deposits

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are occasionally interrupted by the same cm-scale beds of gray

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(GLEY1-6/N) fine to medium grained sandstones (Sfg). They are

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structureless and composed of well-sorted sediments with highly-

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spherical and subangular quartz grains. Their bases display the same

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wavy, sharp, highly bioturbated surface (Fig. 5e). They also contain

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many reworked, abraded and broken brackish and freshwater

420

molluscs.

421

Interpretation

422

This facies association is interpreted as representing a distal

423

delta-front environment more frequently influenced by fluvial input.

424

The mudstones were deposited from suspension in open waters. The

425

episodic intercalations of sandstones are related to increases of sand

426

input coming from the distal margin of distributary channels. Once

427

these sands reach the distal delta lobe, they record a relative

428

deceleration and form cm- to dm-scale, migrating submarine dunes

429

comprising small-scale cross-bedding. The lenticular bedding (Fs) is 16

430

related to wave action and/or winnowing (De Raaf et al., 1977). The

431

climbing ripples may be related to rapid sedimentation rates and non-

432

uniform flows, due to a loss of flow confinement or a decrease in

433

slope gradient (Jobe et al., 2012). The cm-scale, organic-rich, silty

434

laminations could be related to hyperpycnal flows, occurring during

435

episodic larger river discharge events (Mulder et al., 2003;

436

Bhattacharya and MacEachern, 2009; Lamb and Mohrig, 2009). The

437

convolute bedding could have been created when the sandstones

438

were rapidly deposited on the underlying water-saturated mudstones,

439

causing an expulsion of the fluids contained in the mud (Oliveira et

440

al., 2009). The shell-rich sandy beds (Sfg) sporadically intercalated in

441

this facies association suggest storm deposits (Starek et al., 2010;

442

Hampson et al., 2011). This facies association, showing coarsening

443

up, displays an increase in the energy of the depositional process and

444

illustrates a depositional environment closer to the distributary

445

system. In line with previous studies on similar facies associations

446

(Fielding, 2010; Hampson et al., 2011), the depositional setting was

447

interpreted as a distal delta-front environment.

448

5.1.3. Proximal delta-front association

449

Description

450

The third facies association is marked by the lack of mudstones.

451

It is composed of 0.5 to 2 m thick, grayish-brown (2.5Y-5/2), fine

452

grained sandstones. The sands are moderately sorted, have low

453

sphericity and are subangular. They form dm-scale layers and

454

comprise the same three types of cross-beddings, but at a dm-scale

455

(Sr, Ss, St). Sandstones contain 10 to 50 cm thick sets of climbing

456

ripples (Fig. 5b), 20 to 100 cm thick sigmoidal cross-stratification

457

(Fig. 5c) and 10 to 50 cm thick trough cross-stratification (Fig. 5d). 17

458

Sedimentary structures are draped by mm-scale laminae of fragments

459

of terrestrial organic material. Palaeocurrent directions in this facies

460

association were measured on trough cross-stratification, sigmoidal

461

cross-stratification and climbing ripples. They have a mean direction

462

of 180° (n=19; Fig. 7b), with a range from 90° to 330°.

463

Interpretation

464

This facies association is interpreted as deposited in a proximal

465

delta-front environment, in agreement with comparable studies of

466

similar facies associations (Fielding, 2010; Hampson et al., 2011;

467

Forzoni et al., 2015). Sediments were deposited under higher energy

468

conditions compared to the distal delta-front deposits and are

469

therefore mostly sand-dominated. The thicker sandy beds contain

470

larger-scale sedimentary structures, formed by migration of larger-

471

scale dunes. Previous authors deduced that similar sediments were

472

transported towards the basin by subaqueous terminal distributary

473

channels (Bhattacharya, 2006; Olariu and Bhattacharya, 2006). The

474

wide range of palaeocurrent directions may relate to the

475

multiplication of the active terminal distributary channels due to

476

deltaic progradation, as seen in other deltas (Olariu and Bhattacharya,

477

2006).

478

5.1.4. Delta-top facies association

479

The fourth facies association groups four different facies,

480

deposited under specific sedimentary processes, in the same

481

depositional environment.

18

482

5.1.4.1. Interdistributary bay facies

483

Description

484

The sediments deposited in the first facies consist of 1 to 5 m

485

thick sandstones, directly overlying prodelta facies. The greenish-

486

gray (GLEY1-5/5GY), moderately sorted, sandstones (Sm) have low

487

sphericity and are subangular. The sandstones form m-thick

488

continuous layers with a diffuse base (Fig. 4e). The layers coarsen

489

upwards from very fine to medium grain-sizes towards the top of the

490

sandy beds. The sandstones are massive and structureless. A few,

491

dm-scale troughs occur in these beds. The troughs are infilled with

492

mm-scale fragments of terrestrial organic material. The sandstones

493

include some well-preserved in situ freshwater molluscs, such as

494

unionids or viviparids (Supplementary material 3, Fig. 8). They also

495

contain many vertical and horizontal burrows 0.5 to 1 cm wide and 5

496

to 10 cm deep, made by Cruziana ichnofossils, such as

497

Cylindrichnus.

498

Interpretation

499

This facies was deposited from suspension under low-energy

500

and low salinity conditions. Environmental conditions are

501

corroborated by the presence of freshwater molluscs, burrows, and

502

terrestrial organic material. In line with previous studies (Elliott,

503

1974; Overeem et al., 2003), this facies is interpreted to be deposited

504

in an interdistributary bay environment, between distributary

505

channels. Sand-laden currents entered and progressive encroached

506

the interdistributary bay, producing a coarsening upwards succession.

507

Finer-scale sedimentary structures and thin intervening bay

508

mudstones and sandstones were probably erased by intensive

509

bioturbation. 19

510

5.1.4.2. Distributary mouth bar facies

511

Description

512

The second facies forms 2 to 5 m thick sandstone beds. The

513

grayish-brown (2.5Y-5/2), fine to medium grained sandstones are

514

moderately-sorted and comprise low-sphericity and subangular

515

quartz grains. They form m-scale beds with weak inverse grading

516

from the base to the center of the beds. Sand beds may also

517

occasionally record weak normal grading from the middle to the top

518

of the beds. The sandstones contain four types of dm- to m-scale

519

cross-bedding (Sr, Ss, Sl, Sh). They display 10 to 50 cm thick

520

climbing ripples at the bases and/or at the tops of the sandy beds

521

(Fig. 5b), 50 to 100 cm thick sigmoidal cross-stratification (Fig. 5c),

522

50 to 200 cm thick low-angle cross-stratification (Fig. 4f) and mm-

523

scale horizontal laminations (Fig. 4g). The foresets of the cross-

524

beddings are draped by mm-scale laminae of fragments of terrestrial

525

organic material (Fig. 6a, 6b). Some cm-scale clay pebbles are

526

sometimes found at the base of this facies.

527

Interpretation

528

This facies is interpreted to be formed in distributary mouth

529

bars, under high-energy depositional processes. The climbing ripples

530

at the bases and tops of these beds were formed by migration of

531

small-scale current ripples, whereas the larger-scale cross-bedding in

532

the middle parts were formed by migration of large-scale dunes. The

533

clay pebbles at the base of the sand beds were likely formed by

534

erosion of the underlying muddy substratum. Scouring occurred due

535

to relatively high-energy and concentrated density flows (Mulder and

536

Alexander, 2001). In comparison to similar studies based on similar

537

facies (Allen, 1983; Olariu and Bhattacharya, 2006; Forzoni et al., 20

538

2015), we interpreted these deposits as being formed in shallow

539

channelized channels, by lateral and longitudinal accretion of

540

distributary mouth bars.

541

5.1.4.3. Channel fill facies

542

Description

543

The third facies displays 2 to 3 m thick sandstones forming

544

several dm-thick layers. The grayish-brown (2.5Y-5/2), fine grained

545

sandstones are moderately sorted, have low-sphericity and are

546

subangular. The sandstones show a haphazard succession of various

547

cm-scale cross-bedding types (Sc, Sh, Sr, Sl). They contain

548

asymmetrical current ripples with an amplitude of 3 to 5 cm and a

549

wavelength of 7 to 10 cm (Fig. 5a), mm-scale horizontal laminations

550

(Fig. 4g), 10 to 50 cm thick climbing ripples (Fig. 5b) and

551

occasionally 30 to 50 cm thick low-angle cross-stratification (Fig.

552

4f). The foresets of the cross-beds and the horizontal laminations are

553

draped by mm-scale laminae of fragments of terrestrial organic

554

material. Palaeocurrent directions were measured on trough cross-

555

stratification, sigmoidal cross-stratification, climbing ripples, low-

556

angle cross-stratification and asymmetrical current ripples, recorded

557

within the interdistributary bay, distributary mouth bar and channel

558

fill deposits. The measurements display a mean direction of 165°

559

(n=5; Fig. 7c) and a range from 0° to 210°. They highlight a very

560

wide range of flow directions. Unfortunately, the available amount of

561

data is insufficient to extract any other useful information.

562

Interpretation

563

These sandstones were gradually deposited on top of the

564

distributary mouth-bar deposits. They display various sedimentary

565

structures that are formed by migration of small-scale current ripples 21

566

or by deposition of fine sediment from suspension. The enrichment in

567

terrestrial organic material draping the sedimentary structures is

568

indicative of waxing and waning of fluvial flow. Sediments were

569

deposited under fluctuating fluvial current velocities. According to

570

work on similar facies (Elliott, 1974; Fielding, 1986; Bhattacharya,

571

2006), this facies was interpreted as the infill of a channel,

572

progressively affected by avulsion.

573

5.1.4.4. Coastal plain facies

574

Description

575

The last facies includes 0.2 to 0.5 m thick structureless clays

576

(C). The clays are very-dark-gray (GLEY1-3/N) and are rich in

577

dispersed mm- to dm-scale fragments of terrestrial organic material

578

(Fig. 4a). The clays contain some well-preserved in situ freshwater

579

molluscs, such as unionids (Rumanunio rumanus) or pachychilids

580

(Tinnyea abchasica) (Supplementary material 3, Fig. 8). The top of

581

the clay beds are occasionally showing cm-scale ichnofossils, such as

582

Planolites, forming horizontal burrows 0.5 to 1 cm wide and 1 to 3

583

cm deep. More rarely, the top of these clay beds display vertical roots

584

0.5 to 1 cm wide and 5 to 10 cm deep. The upper 5 to 10 cm of this

585

facies is occasionally indurated.

586

Interpretation

587

The organic-rich mudstones are interpreted as having been

588

deposited from suspension, in low-energy coastal plain mires. The

589

upper parts of this facies, affected by burrows and roots, point to

590

sporadic subaerial exposure of the environments. More prolonged

591

subaerial exposure may have indurated the upper parts of these

592

deposits. In line with previous studies of similar facies (Fielding,

593

2010; Hampson, 2010; Forzoni et al., 2015), we interpreted the 22

594

organic-rich layers to be deposited on coastal plains during fluvial

595

flooding.

596

5.1.5. Hardground

597

Description

598

The last facies association consists of 0.2 to 0.4 m thick

599

sandstones. The sandstones (Sfr) form dm-thick layers with wavy,

600

sharp, erosive bases (Fig. 5f). Sediments are quartz-rich with highly

601

spherical and subangular grains. They are fine to medium grained

602

and well-sorted sandstones. The weathered surfaces of these

603

sandstones have a noticeable reddish-brown color (2.5YR-4/4),

604

whereas the fresh surface is more grayish (GLEY1-4/N). The

605

sandstones are mostly structureless, but occasionally show cm- to

606

dm-scale low-angle cross-stratification. Microscopic observations

607

realized on thin-sections show enrichment in subangular glauconite

608

grains (Fig. 6c, 6d). The sandstones also contain high concentrations

609

of shells. The shells are often abraded or broken, and composed of a

610

mix between brackish and freshwater molluscs (Supplementary

611

material 3). The sandstones display iron cement, that is post-

612

diagenetically oxidized, distributed throughout the entire sand bed,

613

leading to the formation of indurated layers.

614

Interpretation

615

These sandstones form hardgrounds, created under high-energy

616

depositional processes. The formation of erosive sand beds,

617

comprising mature sands and many abraded and reworked shells,

618

requires erosion and sediment reworking along the shoreface

619

(Weimer, 1988; Scarponi et al., 2013). The formation of glauconite

620

necessitates slow sedimentation rates down to slight erosion (Cloud,

621

1955). Subsequent winnowing processes may have diminished the 23

622

sedimentation rate and caused episodic sediment starvation (Kidwell

623

and Aigner, 1985; Brett, 1995), leading to the formation of

624

condensed layers (Kidwell, 1989; Abbott and Carter, 1994; Brett,

625

1995; Scarponi et al., 2013). Similar to previous interpretations

626

(Nummedal and Swift, 1987; Weimer, 1988; Murakoshi and Masuda,

627

1992; Cattaneo and Steel, 2003; Hurd et al., 2014), we interpret the

628

formation of such oxidized shell-rich hardgrounds to occur during

629

relative water-level rises and therefore represent flooding surfaces.

630

Along our section, these deposits display a red weathering color and

631

are cemented, which is probably the result of post-diagenetic

632

oxidation during subaerial exposure.

633

5.2. Fauna assemblages

634

The Slănicul de Buzău section contains very rich mollusc and

635

ostracod assemblages. We identified 25 ostracod species (Figs. 9-12)

636

and 47 mollusc species (Fig. 8). Molluscs comprise about 70%

637

bivalve and about 30% gastropod species.

638

5.2.1. Biofacies

639 640 641

Based on macrofaunal observations, three major biofacies were identified within the studied deltaic sedimentary succession. The first biofacies comprised an autochthonous assemblage of

642

several cardiid species, such as Euxinicardium olivetum, Pontalmyra

643

tohanensis and Chartoconcha rumana (Fig. 8). They were

644

preferentially found in clayish prodelta deposits and occasionally in

645

clayish distal delta-front sediments (Supplementary material 3).

646

These species demonstrate brackish water environments (Nevesskaya

647

et al., 2001).

24

648

The second biofacies included an autochthonous assemblage of

649

unionid, pachychilid and viviparid species, such as Rumanunio

650

rumanus, Tinnyea abchasica and Viviparus rumanus (Fig. 8). They

651

were mostly recorded in clayish delta-top environments, such as

652

coastal plains, and at times in proximal delta-front sediments

653

(Supplementary material 3). These species indicate fresher water

654

conditions (Mandic et al., 2015; Rundić et al., 2016).

655

The third biofacies consisted of a mixture of broken and abraded

656

shell and shell fragments, recorded in erosive sandstone beds.

657

Molluscs were transported post mortem from proximal to more distal

658

deltaic environments. They were commonly deposited within storm

659

events or flooding surfaces (Supplementary material 3).

660

5.2.2. Biostratigraphy

661

The evolution of mollusc and ostracod assemblages was

662

analyzed throughout the studied section, in order to identify the

663

stratigraphic position of the boundary between the Lower and Upper

664

Dacian regional substages, as defined by Marinescu and Papaianopol

665

(1995).

666

In the lower part of the investigated section, we found several

667

index mollusc species of the Lower Dacian, such as Stylodacna

668

heberti, Pachydacna (Parapachydacna) serena, Psilodon munieri,

669

Zamphiridacna orientalis and Viviparus argesiensis (Fig. 8). They all

670

display their latest occurrences around 445-503 m, except for

671

Stylodacna heberti which extends to 621 m (Supplementary material

672

4). The Lower Dacian is similarly marked by several characteristic

673

ostracod species. The most common is Cyprideis ex gr. torosa,

674

associated with Candona neglecta, Caspicypris alta, Camptocypria

675

balcanica, Pontoniella ex gr. quadrata, Scottia dacica, Amplocypris 25

676

sp. and Cytherissa boghatschovi (Figs. 9, 10). Beside these species,

677

we also noted in this interval the presence of Amnicythere

678

multituberculata, A. andrusovi, A. ex gr. cymbula, Loxoconcha

679

schweyeri and L. babazananica.

680

The two major index mollusc species for the Upper Dacian

681

encountered in the investigated section are Psilodon haueri and

682

Zamphiridacna zamphiri (Fig. 8). Their first occurrence is around

683

445-561 m (Supplementary material 4). We also found several

684

characteristic ostracod species for the Upper Dacian. In this part,

685

Cyprideis ex gr. torosa becomes more abundant. It is associated with

686

several other species, such as Cytherissa bogathschovi, C. lacustris,

687

Caspiocypris ornatus, Cyprinotus sp., Amplocypris sp., Scottia

688

kempfi and S. bonnei (Figs. 11, 12). We also noticed low abundance

689

of several additional species like Pontoniella ex gr. quadrata,

690

Ilyocypris bradyi, I. gibba, Darwinula stevensoni and Cyclocypris

691

laevis.

692

On the basis of these observations, the stratigraphic position of

693

the boundary between the Lower and Upper Dacian regional

694

substages was identified around 445-503 m in the Slănicul de Buzău

695

section.

696

5.3. Deltaic stratigraphy

697

5.3.1. Regressive parasequences

698

The litho- and biofacies form facies associations which tend to

699

appear in the same stratigraphic order throughout the entire section.

700

They generally form sedimentary successions of about 15 m thick

701

(Fig. 13), which may occasionally extend to a maximum thickness of

702

about 40 m. These sedimentary successions begin with 1 to 13 m

26

703

thick massive or laminated prodelta mudstones with autochthonous

704

cardiid species from biofacies 1 (Fig. 13, logs A-F). The prodelta

705

deposits are overlain by 1 to 5 m thick distal delta-front mudstones

706

with thin sandy intercalations (Fig. 13, logs A, C-F). The transition

707

from distal delta-front to proximal delta-front is marked by a

708

progressive coarsening-up and the deposition of 0.5 to 3 m thick

709

small-scale cross-bedded sandstones with thin muddy intercalations

710

(Fig. 13, logs C-F). Successions continue upwards with several delta-

711

top deposits, which occasionally correspond to 1 to 5 m thick

712

massive sandstones deposited in interdistributary bays (Fig. 13, log

713

B). These sandstones are deposited directly on top of the prodelta and

714

delta-front sediments. They mark the transition from distal to

715

shallower and more restricted depositional environments, without

716

recording any deltaic sandy input. On other occasions, the deltaic

717

input is recorded and prodelta and delta-front sediments are overlain

718

by 2 to 5 m thick distributary mouth bars, forming large-scale cross-

719

bedded sandstones (Fig. 13, logs C, E-F). Distributary mouth bars

720

sometimes erode the underlying proximal delta-front and are directly

721

deposited on top of distal delta-front deposits (Fig. 13, log A). The

722

distributary mouth bars are infrequently overlain by 1 to 3 m thick

723

channel fill deposits, creating small-scale cross-bedded organic-rich

724

sandstones (Fig. 13, log C) or 0.2 to 0.5 m thick coastal plain

725

deposits, with bioturbated organic-rich clays and an autochthonous

726

faunal assemblage of unionid, pachychilid and viviparid species from

727

biofacies 2 (Fig. 13, log E).

728

Each sedimentary succession displays a shallowing-upward

729

trend, regressing from deeper open water towards shallower fluvial

730

environments. Regressive successions are bounded by oxidized, 27

731

shell- and glauconite-rich flooding surfaces with broken and abraded

732

shell assemblage from biofacies 3 (Fig. 13, logs A-C, E-F). Flooding

733

events produced basal erosional unconformities. These surfaces were

734

formed during relative water-level transgressions, corresponding to

735

delta-lobe switching, which formed in total 64 shallowing-upwards

736

successions, defined in the literature as parasequences (Catuneanu et

737

al., 2011). Parasequences are illustrated in a relative water-level

738

curve, based on attributing a depth rank to the facies associations

739

(Table 1), which highlights numerous relative water-level variations

740

of low magnitude (Fig. 3b).

741

5.3.2. Regressive sequences

742

In addition to water-level variations of low magnitude, the

743

relative water-level curve displays variations of higher magnitude.

744

The 64 parasequences can be stacked in larger-scale regressive

745

events. In other well documented cases, larger-scale events are

746

bounded by major unconformities usually correlated throughout the

747

entire basin and defined in the literature as sequences (Catuneanu et

748

al., 2011).

749

In our section, the parasequences can be grouped into nine low-

750

order regressive sequences (Fig. 3b), each of them enclosing between

751

17 to 29 regressive parasequences. The low-order regressive

752

sequences are delimited by well-developed delta-top facies, such as

753

m-thick distributary mouth bars or channel fill deposits. The low-

754

order sequences can themselves be stacked into three high-order

755

regressive sequences (Fig. 3b). Each high-order regressive sequence

756

encloses three regressive low-order sequences. The high-order

757

regressive sequences are bounded by even shallower delta-top facies,

758

such as m-thick channel fills or dm- to m-thick coastal plain deposits, 28

759

marked by enrichment in terrestrial organic material, ichnofossils and

760

freshwater faunas. These low- and high-order sequences highlight

761

larger scale regressive events, related to larger relative water-level

762

variations in the basin.

763

5.3.3. General regressive trend

764

The section records numerous water-level variations of various

765

amplitudes, which are superposed onto a general regressive trend

766

seen on the scale of the entire sedimentary succession (Fig. 3b). The

767

base of the section is mostly mud-dominated and comprises 15 m

768

thick parasequences, which are on average composed of 77%

769

mudstones (Fig. 14). The muddy regressive parasequences start with

770

m-thick prodelta deposits, showing numerous density-driven and

771

hyperpycnal flows (Fig. 14a). Prodelta deposits are overlain by m-

772

thick distal delta-front deposits, often disturbed by convolute bedding

773

(Fig. 14b). The succession regresses up to dm-thick proximal delta-

774

front deposits, forming thin sandy beds with small-scale sedimentary

775

structures (Fig. 14c). Delta-top deposits hardly occur in the basal part

776

of the section. Furthermore, at the base, parasequences are grouped in

777

94 to 137 m thick low-order sequences and in 350 m thick high-order

778

sequences (Fig. 3b).

779

Towards the top of the section, the sedimentary succession

780

becomes sand-dominated. Parasequences are on average composed of

781

56% mudstones, whereas the amount of sand has doubled compared

782

to the base of the section (Fig. 15). In the sandy regressive

783

parasequences, the prodelta deposits are only dm- to m-thick, or are

784

absent. They are overlain by dm- to m-thick distal delta-front

785

deposits, showing frequent density-driven flows (Fig. 15a). On top of

786

the distal delta-front deposits, m-thick proximal delta-front deposits 29

787

are deposited and formed by m-thick sandstone layers. These sand

788

beds contain various large-scale cross-beds. The sandy parasequences

789

commonly regress up to m-thick erosional distributary mouth bar

790

deposits (Fig. 15b), or more rarely into m-thick channel fill deposits.

791

They are occasionally capped by cm- to dm-thick coastal plain

792

deposits, showing organic-rich sediments with roots and burrows

793

(Fig. 15c). The thickness of the parasequences decreases to 9 m

794

towards the top of the section. These thinner parasequences can be

795

grouped into thinner low- and high-order sequences. The low-order

796

sequence thicknesses decrease to 65 to 121 m and the high-order

797

sequence thicknesse decrease to 208 m (Fig. 3b).

798

5.4. Autogenic delta-lobe switching

799

Deltaic progradation formed numerous regressive

800

parasequences and sequences, generated by frequent delta-lobe

801

switching. Thanks to the robust time frame available for this section,

802

it is possible to test if the sediment rhythmicity of this delta is

803

autogenic or allogenic (Fig. 17).

804

On a small-scale, we observed 64 parasequences, when there are

805

only 27 precession cycles in the corresponding time-interval.

806

Parasequences have a frequency of 12 ± 9 kyr. The filters with 13.6-

807

25 m and 25-37 m bandpass-width are not in tune with the 23 kyr

808

precession cycle. Parasequences repeat too frequently to be coeval

809

with any astronomical cycles. On a larger scale, the nine low-order

810

sequences display a frequency of 81 ± 44 kyr. The 64-111 m and 50-

811

138 m filters do not correlate with the 40 kyr obliquity cycle or with

812

the 100 kyr eccentricity cycle. The low-order sequences can likewise

813

not be reliably linked to astronomical cycles. On an even larger scale,

30

814

the three high-order sequences have a frequency of 243 ± 61 kyr.

815

Similarly, they occur too often compare to the 400 kyr eccentricity

816

cycle. In summary, it appears that neither the parasequences, nor the

817

sequences reflect astronomical climatic forcing.

818

The absence of correlation between the

819

sequences/parasequences and the astronomical cycles may be due to

820

condensed intervals and minor hiatuses recorded in the sedimentary

821

succession. These events might have affected the time frame of

822

deltaic progradation. Each parasequences is topped by shell-rich

823

oxidized layers, marked by a basal erosional unconformity, formed

824

during flooding events. Each of these events generated a minor

825

hiatus, due to sediment starvation and winnowing occurring during

826

relative water-level rises. Furthermore, at about 250 m in the section,

827

the boundary between two low-order sequences is marked by a series

828

of four erosional shell-rich oxidized layers, stacked together in a 5 m

829

thick interval. Each of these beds represents a full parasequence of

830

about 1 m thick. As parasequences are on average 15 m thick

831

elsewhere in the section, we may estimate that about 55 m sediments

832

have been eroded or were not deposited. This resulted in a dramatic

833

decrease in sedimentation rate from 144 cm/kyr to 65 cm/kyr (Fig.

834

16). The sedimentary succession seems therefore to have recorded a

835

major hiatus in this part of the section, which may have impacted its

836

time frame and possible correlation with astronomical cycles.

837

However, even if the section displays some hiatuses, the

838

parasequences and sequences are most likely the result of autogenic

839

relative water-level variations.

31

840

6. Discussion

841

6.1. Palaeoenvironmental evolution of the mid-Pliocene

842

eastern Dacian Basin

843

During the mid-Pliocene Dacian stage, the Dacian Basin

844

received erosion products of the uplifting Carpathians (Fig. 1a, 1b).

845

The eastern part of the Carpathians was drained by a river running

846

parallel to the mountain belt (Jipa, 1997; Popov et al., 2006; Jipa and

847

Olariu, 2009; Leever et al., 2010; Stoica et al., 2013; Fongngern et

848

al., 2016; Matoshko et al., 2016). The Dacian alluvial and deltaic

849

river system prograded southwards, with a mean palaeocurrent

850

direction of 195° (Fig. 7d). The system prograded on the northern

851

margin of the Dacian Basin and progressively infilled the deep

852

southern foreland depression (Jipa, 1997; Sanders et al., 1999;

853

Tărăpoancă et al., 2003; Panaiotu et al., 2007). The basin became

854

overfilled during the Romanian regional stage (Jipa and Olariu,

855

2009). At that time, the Dacian deltaic system merged with the

856

Danube system and sediments started to overspill into the Black Sea

857

at around 4 Ma (De Leeuw et al., 2018; Olariu et al., 2018).

858

The Dacian deltaic system remained relatively stable during the

859

entire Dacian regional stage. The long-term stability of the sediment

860

system was ensured by an equilibrium between subsidence and

861

sedimentation rates (Bertotti et al., 2003). The regional subsidence

862

rate of 90 cm/kyr (Tărăpoancă et al., 2003) was balanced in the

863

eastern Dacian Basin by average sedimentation rates of 90 cm/kyr

864

along the more northern Râmnicu Sărat section (Fig. 2b) (Vasiliev et

865

al., 2004) and 139 cm/kyr along the more southern Slănicul de Buzău

866

section (Fig. 17). This balance permitted a major storage of 32

867

sediments within only 0.6 Ma. About 1300 m of deltaic sediments

868

were recorded along the Râmnicu Sărat section (Vasiliev et al., 2004)

869

and about 835 m along the Slănicul de Buzău section. These two

870

sections recorded the southern progradation of the Dacian deltaic

871

system through the eastern Dacian Basin. Deltaic progradation

872

generated a north to south decrease of sediment grain size and a

873

thinning of the deltaic and alluvial sand bodies. In the northern

874

Râmnicu Sărat section, sand bodies are about 2 to 3 m thick (Jipa and

875

Olariu, 2009) and composed of medium grained sediments (Vasiliev

876

et al., 2004). In the southern Slănicul de Buzău section, sand bodies

877

are only 1 to 2 m thick and are composed of fine grained sediments.

878

The depositional environment thus became progressively more distal

879

towards the northern margin of the Dacian Basin.

880

The Slănicul de Buzău section records progradation of the entire

881

Dacian deltaic system though time. The section registers a gradual

882

coarsening-upward trend, coeval with progressive thinning of the

883

regressive parasequences and sequences (Fig. 3b). These trends

884

demonstrate an increase in energy of depositional processes and a

885

decrease in accommodation space. The sedimentary succession

886

records in parallel two major changes in faunal assemblages. The

887

first change marks the boundary between the Lower and Upper

888

Dacian regional substages at about 445-503 m (Supplementary

889

material 4). This change occurred gradually, as the transitional

890

interval extends to 621 m for some of the species. Importantly, the

891

transition occurred independently from changes in the depositional

892

environment. This suggests that the boundary between the Lower and

893

Upper Dacian regional substages might be synchronous throughout

894

the entire Dacian Basin. On the basis of the present age model, we 33

895

estimate this boundary to be within chron C3n.1r at an age around

896

4.42 Ma (Fig. 17). The second major change in faunal assemblage

897

corresponds to the boundary between the Upper Dacian and the

898

Romanian regional stages. The boundary, marked by a relatively

899

abrupt transition from brackish water to freshwater faunas, is located

900

at about 835 m in the sedimentary succession and was dated at

901

around 4.2 Ma (Van Baak et al., 2015). The transition in faunal

902

assemblages is covalent to the first coal layers deposited in delta-top

903

environments, which are observed in the upper-most part of the

904

section (Fig. 3a). As the boundary between the Upper Dacian and the

905

Romanian stages is linked to the depositional setting, it might be

906

diachronous throughout the Dacian Basin.

907

6.2. Distinctive deltaic features in semi-isolated basins

908

The Slănicul de Buzău section documents the mid-Pliocene

909

infill of the semi-isolated Dacian Basin by a substantial prograding

910

deltaic system. This enclosed depositional environment seems to

911

have influenced the sedimentary facies and internal architecture of

912

this delta.

913

The restricted basin formed a protected depositional

914

environment, with limited wave and tide activity. Deposits are only

915

occasionally disturbed by minor wave and storm action, creating

916

small-scale lenticular bedding and thin shell-rich storm deposits.

917

There is very little evidence of sediment reworking, which we relate

918

to low-energy in this protected environment. However, the absence

919

of indications of wave or tide influence could also be related to the

920

very strong river input entering the basin.

34

921

The isolated nature of the basin caused lowered waters salinities

922

(Popov et al., 2006; Leever et al., 2010; Stoica et al., 2013). Waters

923

with lowered salinities present lowered water densities, which means

924

that hyperpycnal flows are more likely to occur than in regular sea

925

waters (Sturm and Matter, 1978; Mulder et al., 2003). Fine grained

926

organic-rich sediments deposited by hyperpycnal plumes are very

927

common in our section. The terrestrial organic material, occasionally

928

found on foresets of cross-bedded sandstones, confirms a proximal

929

low-energy depositional setting. Furthermore, the section records

930

numerous ichnofossils, in particular in interdistributary bay and

931

coastal plain deposits, which might be favored by low salinity and

932

low-energy settings. Moreover, this depositional environment

933

typically encouraged enrichment in situ mollusc fauna.

934

The delta prograded into a shallow depositional environment.

935

Deltaic progradation formed thin sand bodies with an average

936

thickness of 1 to 2 m, whereas they can be more than 10 m thick in

937

the open ocean (e.g., Olariu and Olariu, 2015). Sediments were

938

deposited in sand bodies with an erosive base and dm- to m-thick

939

cross-bedded strata. Sharply-based sand bodies with relatively small-

940

scale cross-bedding are often formed in shallow depositional settings

941

(Fielding, 2010; Vincent et al., 2010). Due to reduced water depths,

942

deltaic progradation generated numerous thin regressive

943

parasequences. They are on average only 13.5 m thick in our section.

944

Parasequences are known to be relatively thin in shallow

945

environments (Bohacs et al., 2000; Sztanó et al., 2013), whereas they

946

can become hundreds of m-thick in the open ocean (e.g., Olariu and

947

Olariu, 2015). Moreover, the more distal parasequences at the base of

948

our section are on average about 5 m thicker than the proximal ones 35

949

at the top of the section. This upwards thinning highlights a gradual

950

decrease in the rate of accommodation space available through time.

951

Sediment progradation occurred on a low-gradient slope.

952

Numerous thin regressive parasequences were formed by frequent

953

migration of multiple small distributary channels, covering a wide

954

range of palaeocurrent directions. The formation of a wider

955

distributary area can be enhanced by low-gradient slopes

956

(Bhattacharya, 2006; Olariu and Bhattacharya, 2006). Distributary

957

channels were affected by repeated delta-lobe switching, occurring

958

due to recurrent avulsion. Delta-lobe switching seems to have been

959

strictly controlled by autogenic processes. Neither the frequency of

960

the parasequences, low-order sequences nor high-order sequences is

961

in tune with astronomical climatic forcing (Fig. 17). Conversely,

962

farther to the north, in the Râmnicu Sărat sections, more proximal

963

deltaic deposits seem to have been affected by astronomical

964

precession cycles (Vasiliev et al., 2004). This suggests Milankovitch

965

cycles can be registered in river-dominated deltas (Sacchi and

966

Müller, 2004; Li and Bhattacharya, 2013), but that climate forcing of

967

river-dominated deltas may, in some cases, become overridden by

968

frequent autogenic delta-lobe switching in more distal environments

969

(Castelltort and Van Den Driessche, 2003).

970

As autogenic delta-lobe switching occur, relative water-level

971

rise was noted on top of each abandoned delta lobe. The upper

972

surface of abandoned lobes consequently recorded sediment

973

starvation and winnowing, creating hardgrounds, enriched in shell

974

fragments and glauconite. Such layers are often found in restricted

975

basins (Cloud, 1955; Cattaneo and Steel, 2003). Along the studied

976

section, these layers highlight post-diagenetic induration and 36

977

oxidization only during the sedimentary interval corresponding to the

978

mid-Pliocene Dacian deltaic system. High quantities of glauconite

979

and organic material found in the deposits suggest increased runoff,

980

which may have occurred in response to higher global temperatures

981

during the Pliocene (Fedorov et al., 2013). Similar iron-rich

982

sediments were formed in other locations around the Black Sea

983

during the Pliocene (Nevesskaya et al., 2003; Krijgsman et al., 2010).

984

Increased temperatures during the Pliocene could also have caused

985

increased weathering of adjacent land areas, leading to increased iron

986

content of the waters in the basins involved (Muratov, 1964). The

987

depositional controls of these layers are still not well understood and

988

additional research is needed.

989

6.3. Typical example of a river-dominated delta?

990

The studied deltaic system was interpreted, according to the

991

classification of Galloway (1975), as a river-dominated delta. The

992

sediments display a strong river influence, with a sedimentation rate

993

of 152 cm/kyr on average. In the older part of the section, sediments

994

were preferentially transported into the basin through density-driven

995

and hyperpycnal flows, whereas towards the younger part, they were

996

progressively transported through fluvial channels. The sediment

997

supply fed a multitude of deltaic lobes and distributary channels,

998

covering a wide distributary area. The deltaic system does not display

999

any evidence for tide interference and reveals only minor wave

1000

activity. There is no evidence for sediment reworking except during

1001

flooding events. Additionally, no sand spit formation, symmetrical

1002

wave structures or clay draping on foresets were observed along the

37

1003

section. The absence of these features suggests a strictly river-

1004

dominated delta.

1005

However, in enclosed basins where wave and tide interferences

1006

are weak or even absent, this classification might not be the most

1007

appropriate. A more relevant classification could be the one proposed

1008

by Postma (1990), who considers the sedimentary architecture of

1009

river-dominated deltas depending on water depth and gradient of the

1010

basin. These two forcing factors seem to have played an important

1011

role on the sedimentary architecture of the studied delta. With

1012

regards to this classification, we could interpret it as a mouth bar-type

1013

delta with a Gilbert-type profile. Nevertheless, this classification does

1014

not include all the basin characteristics perceived along the studied

1015

section. Classifications established on deltas evolving in the open

1016

ocean might therefore not be applicable for deltaic system prograding

1017

into (semi-)isolated basins.

1018

Conclusions

1019

During the mid-Pliocene, the Dacian Basin formed an

1020

embayment of the Black Sea. The northern margin of this semi-

1021

isolated basin was supplied by a river-dominated delta prograding

1022

southwards east of the Carpathians. Deltaic progradation gradually

1023

infilled the basin throughout the Dacian regional stage (4.8-4.2 Ma),

1024

before being replaced by a predominantly fluvial environment at the

1025

onset of the Romanian regional stage (4.2-1.8 Ma).

1026

The delta prograded into the restricted Dacian Basin, which

1027

formed a protected, brackish water, shallow depositional

1028

environment, with a low-gradient slope. This atypical depositional

38

1029

setting strongly influenced the sedimentary architecture of the deltaic

1030

system, which differs from a typical open ocean delta. By contrary to

1031

well-known river-dominated deltas, this deltaic system shaped a

1032

larger number of small terminal distributary channels, experiencing

1033

frequent delta-lobe switching. As a result, numerous thin regressive

1034

parasequences were created, each of them overlain by oxidized shell-

1035

rich and glauconite-rich flooding surfaces. Detailed

1036

cyclostratigraphic analyses were conducted using the

1037

magnetostratigraphic time-frame for the section. They reveal that

1038

neither the 64 observed parasequences, nor the higher order

1039

sequences in which they recombine, are governed by precession,

1040

obliquity or eccentricity astronomical cycles. We thus infer that

1041

delta-lobe switching was mainly controlled by autogenic processes.

1042

If this study provides valuable information in term of deltaic

1043

sedimentary architecture and depositional processes in a restricted

1044

depositional setting, further research is still necessary to improve our

1045

understanding of the drivers acting on various (semi-)isolated basins.

1046

Acknowledgements

1047

This research was supported by the project PRIDE

1048

(Pontocaspian RIse and DEmise), which has received funding from

1049

the European Union's Horizon 2020 research and innovation

1050

program, under the Marie Sklodowska-Curie [grant agreement No

1051

642973]. Part of this research was conducted during a secondment at

1052

CASP, one of the partner institutes of PRIDE, which we thank for its

1053

cooperation and hospitality. Finally, we thank the reviewers and

39

1054

Jasper Knight, editor of Sedimentary Geology, for their constructive

1055

comments of the manuscript.

1056

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56

Highlights (for review)

    

Facies model for delta prograding into a protected, shallow, brackish basin. Low-angle slope induced multitude of terminal distributary channels. Frequent delta-lobe switching formed numerous thin parasequences. Parasequences are overlain by indurated and oxidized flooding surfaces. Autogenic deltaic avulsion was not governed by astronomical climate forcing.

Figure captions

Fig. 1. (a) Palaeogeographic map of the Paratethyan Basins during the early Pliocene (adapted from Popov et al., 2006). (b) Enlarged palaeogeographic map of the semi-isolated Dacian Basin during the mid-Pliocene, with associated drainage systems and water connections to adjacent basins (adapted from Popov et al., 2006). The studied section is marked by a red star.

Fig. 2. (a) International stratigraphic nomenclature during the Plio-Pleistocene (Cohen et al., 2013, updated), correlated with the regional stages of the Dacian Basin (adapted from Vasiliev et al., 2005; Krijgsman et al., 2010; van Baak et al., 2015). (b) Detailed geological map of the eastern part of the Carpathians foreland basin affected by large-scale folding (adapted from Motas et al., 1966; Dumitrescu et al., 1970; van Baak et al., 2015). (c) Enlarged geological map of the mid-Pliocene interval of the Slănicul de Buzău section, running through the top of a plunging anticline. Fig. 3. (a) Sedimentological log of the Slănicul de Buzău section with measured magnetostratigraphic time frame. In orange, stratigraphic positions of the detailed sedimentological logs of some of the parasequences shown in Figure 13. In red, stratigraphic positions of the clayish and sandy parasequences illustrated in Figure 14 and 15. In black, stratigraphic positions of the samples analysed for faunal content. (b) Relative water-level curve of the Dacian interval based on facies depth ranks, where 0 is deep, 9 is shallow (Table 1), and interpreted low-order sequences (dark gray) and highorder sequences (light gray).

Fig. 4. Photos of the different lithofacies identified along the section, facies codes in brackets (Part I). (a) Organic-rich clays C. (b) Massive mudstone Fm. (c) Laminated mudstone Fl. (d) Lenticular mudstone Fs. (e) Massive sandstone Sm. (f) Low-angle cross-stratified sandstone Sl. (g) Horizontally laminated sandstone Sh.

Fig. 5. Photos of the different lithofacies identified along the section, facies codes in brackets (Part II). (a) Current rippled sandstone Sc. (b) Climbing rippled sandstone Sr. (c) Sigmoidal cross-stratified sandstone Ss. (d) Trough cross-stratified sandstone St. (e) Grayish shelly sandstone Sfg. (f) Reddish shell-rich sandstone Sfr.

Fig. 6. Thin-section photograph of a sigmoidal cross-stratified sandstone realised (a) under polarized light and (b) under polarized and analysed light. Thin-section photograph of oxidized shell-rich sandstones (c) under polarized light and (d) under polarized and analysed light. Some quartz grains (Qtz), glauconite grains (Gl), organic material fragments (OM) and shells are highlighted on the pictures.

Fig. 7. Palaeocurrent directions measured along the studied section. (a) Direction of progradation measured for the pro-delta sediments. (b) Direction of progradation measured for delta-front sediments. (c) Direction of progradation measured for delta-top sediments, including mouth bar, interdistridutary bay and channel fill deposits. (d) Direction of progradation measured for the entire Dacian stage interval.

Fig. 8. Cardiids (1-26), unionids (27-28), dreissenids (29-32) and viviparids (33-34) present in the Dacian interval of the Slănicul de Buzău section. Please see online supplementary material 2 for mollusc species authorities. 1. Euxinicardium olivetum; 2. Tauricardium olteniae; 3. Dacicardium rumanum; 4. Phyllocardium planum; 5. Pontalmyra tohanensis; 6. Pontalmyra conversa; 7. Pseudocatillus dacianus; 8. Chartoconcha bayerni; 9. Chartoconcha ovata; 10. Chartoconcha rumana; 11. Caladacna steindacheri; 12. Stylodacna heberti; 13. Pachydacna (Parapachydacna) orbiculata; 14. Pachydacna (Parapachydacna) cobalcescui; 15. Pachydacna (Parapachydacna) serena; 16. Prosodacnomya sturi; 17. Prosodacna semisulacata; 18. Prosodacna minima; 19. Prosodacna obovata; 20. Psilodon haueri; 21. Psilodon munieri; 22-23. Psilodon neumayri; 24. Zamphiridacna motasi; 25. Zamphiridacna orientalis; 26. Zamphiridacna zamphiri; 27. Rumanunio rumanus; 28. Hyriopsis krejcii; 29. Dreissena polymorpha; 30. Dreissena rimestiensis; 31. Dreissena rostriformis; 32. Andrusoviconcha botenica; 33. Viviparus argosiensis; 34. Viviparus rumanus. The scale bar represents 1 cm. Fig. 9. Most common ostracods present in the Lower Dacian along the Slănicul de Buzău section (Part I, 1-22). Please see online supplementary material 2 for ostracod species authorities. 1-2. Amplocypris dorsobrevis; 3-8. Pontoniella ex gr. quadrata; 9-10. Candona (Camptocypria) ex gr. balcanica;

11-12.

Candona

(Caspiocypris)

alta;

13.

Bakunella

dorsoarcuata;

14-15.

Fabaeoformiscandona sp.; 16-17. Candona (Camptocypria) sp.; 18. Tyrrhenocythere sp.; 19. Loxoconcha babazananica; 20. Amnicythere andrusovi; 21-22. Amnicythere ex gr. cymbula. Fig. 10. Most common ostracods present in the Lower Dacian along the Slănicul de Buzău section (Part II, 1-20). 1-8. Cytherissa bogatschovi; 9-20. Cyprideis ex gr. torosa. Fig. 11. Most common ostracods present in the Upper Dacian along the Slănicul de Buzău section (Part I, 1-22). 1-6. Candona (Caspiocypris) alta; 7. Candona (Caspiocypris) ornata; 8-9. Candona (Caspiocypris) ex gr. neglecta; 10-13. Candona (Camptocypria) ex gr. balcanica; 14-15. Scottia kempfi; 16-17. Cyclocypris laevis; 18. Amplocypris sp.; 19-22. Pontoniella ex gr. quadrata.

Fig. 12. Most common ostracods present along the Slănicul de Buzău section in the Upper Dacian (Part II, 1-22). 1-8. Cyprideis ex gr. torosa; 9-18. Cytherissa bogatschovi; 19-22. Loxoconcha ex gr. schweyeri; 23-24. Ilyocypris bradyi.

Fig. 13. Sedimentological logs of some of the most representative regressive parasequences observed along the section.

Fig. 14. Photos and corresponding sedimentological log from a succession of several thick clayish parasequences. Illustrations of some typical distal sedimentary deposits: (a) hyperpycnal flow deposit, (b) convolute bedding and (c) small-scale cross-stratifications in sandstone.

Fig. 15. Photos and corresponding sedimentological log from a succession of several thin sandy parasequences. Illustrations of some typical proximal sedimentary deposits: (a) frequent densitydriven flow deposits, (b) m-thick erosive distributary mouth bar deposit and (c) indurated and burrowed coastal-plain deposit.

Fig. 16. Sedimentation rates calculated along the section. Comparison with the low-order sequences (dark gray) and the high-order sequences (light gray). Fig. 17. Age model proposed for the Slănicul de Buzău section by correlating the measured magnetostratigraphic time frame with the international magnetostratigraphic timescale from 4.8 to 4.11 Ma (Gradstein et al., 2012). Plotting of the selected bandpass filters against the Milankovitch astronomical target curves: eccentricity, obliquity and precession (Laskar et al., 2011).

Table 1. Facies depth ranks used to create the relative water-level curve of the section.

Table 2. Description of the main lithofacies characteristics and associated sedimentary processes identified along the section.

Supplementary material 1. Stereonet projection of the measured bedding planes along the Slănicul de Buzău section and their poles in blue. Deduced fold axis with its corresponding pole in red. The corrected pole of the fold axis is marked by a green point.

Supplementary material 2. List with mollusc and ostracod species found in the section and references to the authors first describing these species.

Supplementary material 3. Distribution of the mollusc fauna according to depositional environment. In dark red, the marker species for brackish water environments. In light red, the species found mostly in brackish water environments. In dark green, the marker species for fresh water environments. In light green, the species found mostly in fresh water environments.

Supplementary material 4. Distribution of the mollusc fauna according to stratigraphic position. The red frame indicates the stratigraphic position of the transition between the marker species for the Lower Dacian (dark gray) and the marker species for the Upper Dacian (light gray), used to place the boundary between the Lower and Upper Dacian regional substages.

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Table 1 Click here to download Table: Jorissen et al_Table 1_Facies depth ranks.xlsx

Index Depositional environment 0 1 2 3 4 5 6 7 8 9

Open water Distal prodelta Proximal prodelta Distal delta-front Proximal delta-front Distributary mouth bar/Interdistributary bay Hardground Channel fill Coastal plain Peat

Sedimentary structures Mudstones, structureless, in situ brackish water faunas Mudstones, mm-scale planar laminations of silts Mudstones, cm-scale lenticular bedding of silts and very fine sandstones Mudstones, cm- to dm-scale cross-bedding of very fine sandstones Fine grained sandstones, dm-scale cross-bedding of very fine sandstones Medium grained sandstones, dm to m-scale cross-bedding, erosive base, reworked freshwater faunas Medium grained sandstones, erosive base, glauconite, reworked brackish and freshwater faunas Fine grained sandstones, cm-scale cross-bedding, organic material fragments, ichnofossils Clays, structureless, organic material fragments, ichnofossils, in situ freshwater faunas Clays, structureless, cm-scale coal layers, roots

Table 2 Click here to download Table: Jorissen et al_Table 2_Lithofacies characteristics.xlsx

Code Grain size

Sedimentary structures

Sfr

Fine to medium Hints of low-angle cross-stratification, erosive grained sandstone base, induration, reddish oxidation

Sfg

Fine to medium Structureless, erosive base grained sandstone

St

Very fine to medium grained sandstone

Ss

Sr

Trough cross-stratification, preserved cross-set thickness of 10-50 cm, possible erosive base, unidirectional Sigmoidal cross-stratification, preserved crossFine to medium set thickness of 20-100 cm, possible erosive grained sandstone base, possible inverse grading, unidirectional Very fine to Climbing ripple, preserved cross-set thickness of medium grained 10-50 cm, unidirectional sandstone

Inclusions

Processes

Very high energy flow, sediment Reworked brackish and reworking, post-diagenetic induration freshwater shells, glauconite and oxidation Reworked brackish and Very high energy flow, sediment freshwater shells, reworking ichnofossils

Fig. 10f

10e

Organic material fragments along foresets

Moderate energy flow, migration of sinuous crested dunes

10d

Organic material fragments along foresets

Moderate to high energy flow, migration of straight crested dunes

10c

Organic material fragments along foresets

Moderate energy flow, migration of curved crested ripples, abundant suspended material

10b

Moderate energy flow, migration of ripples, abundant suspended material

10a

Sc

Fine to medium Asymmetrical current ripple, amplitude of 3-5 grained sandstone cm and wavelength of 7-10 cm, unidirectional

Organic material fragments along foresets

Sh

Fine to medium Horizontal lamination, possible normal and grained sandstone inverse grading, possible erosive base

Possible organic material Low to high energy flow, plane-bed fragments along laminations flow

9g

Sl

Low-angle cross-stratification, preserved crossFine to medium set thickness of 30-200 cm, possible inverse grained sandstone grading, possible erosive base, unidirectional

Organic material fragments along foresets

High energy flow, migration of low relief dunes

9f

Sm

Fine to medium Structureless, inverse grading, dm-scale troughs grained sandstone infilled with organic material fragments

Reworked and in situ freshwater shells, ichnofossils

Low energy flow

9e

Fs

Mudstone

Lenticular bedding made of trough crossOrganic material fragments stratified silt and very-fine sand, preserved crossalong foresets set thickness of 1-5 cm

9d

Fl

Mudstone

Horizontal lamination

9c

Fm

Mudstone

Structureless

In situ brackish water shells

Very low energy, deposition from suspension

9b

Clay

Structureless, possible coal layers, possible induration

In situ freshwater shells, ichnofossils, roots, organic material fragments

Very low energy, deposition from suspension, possible post-diagenetic induration

9a

C

Fluctuation between very low and moderate energy flow, deposition from suspension and distal outflows Fluctuation between very low and Silts and organic material moderate energy flow, deposition fragments along laminations from suspension and distal outflows

Supplementary material 1 Click here to download Supplementary material for on-line publication only: Jorissen et al_Supplementary material 1_Paleocurre

Supplementary material 2 Click here to download Supplementary material for on-line publication only: Jorissen et al_Supplementary material 2_Species au

Supplementary material 3 Click here to download Supplementary material for on-line publication only: Jorissen et al_Supplementary material 3_Biofacies.x

Supplementary material 4 Click here to download Supplementary material for on-line publication only: Jorissen et al_Supplementary material 4_Biostratigr