1
Sedimentary architecture and depositional
2
controls of a Pliocene river-dominated delta
3
in the semi-isolated Dacian Basin, Black Sea
4 5
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
8 9 10 11 12 13 14 15 16 17 18 19 20
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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22 23
Abstract Sedimentological facies models for (semi-)isolated basins are
24
less well developed than those for marine environments, but are
25
critical for our understanding of both present-day and ancient deltaic
26
sediment records in restricted depositional environments. This study
27
considers an 835 m thick sedimentary succession of mid-Pliocene
28
age, which accumulated in the Dacian Basin, a former embayment of
29
the Black Sea. Detailed sedimentological and palaeontological
1
30
analyses reveal a regression from distal prodelta deposits with
31
brackish water faunas to delta-top deposits with freshwater faunas.
32
Sediments contain frequent hyperpycnal plumes and an enrichment in
33
terrestrial organic material, ichnofossils and in situ brackish and
34
freshwater faunas. Deltaic progradation created thin, sharply-based
35
sand bodies formed by multiple terminal distributary channels,
36
covering a wide depositional area. The system experienced frequent
37
delta-lobe switching, resulting in numerous thin parasequences.
38
Parasequences are overlain by erosive reddish oxidized sand beds,
39
enriched in broken, abraded brackish and freshwater shells. These
40
beds were formed after sediment starvation, on top of abandoned
41
delta lobes during each flooding event. A robust
42
magnetostratigraphic time frame allowed for comparison between the
43
observed sedimentary cyclicity and the amplitude and frequency of
44
astronomical forcing cycles. Our results indicate that parasequence
45
frequencies are significantly higher than the number of time
46
equivalent astronomical cycles. This suggests that delta-lobe
47
switching was due to autogenic processes. We consider the observed
48
facies architecture typical for a delta prograding on a low-gradient
49
slope into a shallow, brackish, protected, semi-isolated basin.
50
Furthermore, in the absence of significant wave and tidal influence,
51
sediment progradation in such a protected depositional setting shaped
52
a delta, strongly river-dominated.
53 54
Key-words: Paratethys, isolated basin, river-dominated delta,
55
regressive parasequences, autogenic forcing, flooding surface
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2
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1. Introduction
58
Restricted sedimentary basins constitute complex depositional
59
environments. They form semi-isolated basins when they display
60
limited water connections to the marine realm, or isolated basins
61
when all connections are interrupted. The sedimentary infill of (semi-
62
)isolated basins is mostly controlled by the relative importance of
63
accommodation space, water supply and sediment supply (Bohacs et
64
al., 2003). Restricted sedimentary basins record generally limited
65
accommodation space compared to the open ocean (Carroll and
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Bohacs, 1999). An important control on the limited accommodation
67
space is the position and relative height of the spill point of the basins
68
(Leever et al., 2010; Yanina, 2014; Fongngern et al., 2016), which
69
may at the same time affect their connectivity to the marine realm
70
(Leever et al., 2011; Ter Borgh et al., 2014). Restricted connectivity
71
with the open ocean results in the prevalence of brackish to
72
freshwater depositional environments and is usually accompanied by
73
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
78
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.,
83
2015). If properly understood, (semi-)isolated sedimentary basins
3
84
may provide valuable environmental, climatic and biological
85
archives. Understanding depositional processes operating in these
86
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-
91
drilling projects (e.g., Ross, 1978; Francke et al., 2016), depositional
92
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,
4
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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
114
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
120
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
127
(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).
144
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
147
Carpathians (Matenco and Bertotti, 2000; Cloetingh et al., 2004;
148
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
150
(Bertotti et al., 2003; Tărăpoancă et al., 2003). The depression was
151
progressively filled with the erosional products of the uplifted
152
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
160
Ma (Vasiliev et al., 2005; Van Baak et al., 2015) (Fig. 2a). At that
161
time, a major delta prograded east of the Carpathians towards the
162
northeastern margin of the Dacian Basin (Jipa, 1997; Jipa and Olariu,
163
2009; Fongngern et al., 2016; Matoshko et al., 2016), whereas
164
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
168
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
170
et al., 2018).
171
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
175
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
177
the surface.
178
3. The Slănicul de Buzău section
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The investigated section crops out along the Slănicul de Buzău
180
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).
184
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
7
191
studied section was therefore deposited between 4.8 Ma and 4.11 Ma.
192
These age constraints are in line with studies of the Dacian regional
193
stage at other locations in the Dacian Basin (e.g., Vasiliev et al.,
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2004).
195
4. Methods
196
4.1. Sedimentological data collecting
197
The Dacian segment of the Slănicul de Buzău section starts
198
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
205
missed a 40 m interval in the middle part of the section, which is now
206
included in our work (Fig. 3a). To support our sedimentological
207
analyses, the section was analysed in greater detail and several key
208
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
213
descriptions. Thin-sections with a thickness of 30 µm were made
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perpendicular to the sedimentary structures for petrographic
215
descriptions.
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216
Detailed sedimentological observations allowed for several
217
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
221
depositional environment and an estimated water depth. A depth
222
ranking scale was subsequently constructed by attributing a number
223
from 0 to 9 to each facies associations, with 0 being the deepest and 9
224
the shallowest depositional environment (Table 1). This ranking scale
225
permitted the reconstruction of a relative water-level curve and the
226
identification of parasequences, including superimposed lower- and
227
higher-order sequences (Fig. 3b).
228
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
232
to remove any tectonically-induced rotations (Supplementary
233
material 1). Corrections were realized with the help of the available
234
palaeomagnetic dataset. We proceeded with a first step of deplunging
235
the fold axis, followed by a second step of correcting for the true
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vertical axis rotation.
237
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
9
243
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
248
basis of these bedding planes was applied to place palaeomagnetic
249
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
256
of the Carpathian Bend zone.
257
Once the palaeocurrent directions were corrected, their
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statistical distribution was calculated for 12 sectors of 30° and plotted
259
on four rose diagrams with a maximum representability of 50%. One
260
of the rose diagrams represents the overall flow direction along the
261
section. Palaeocurrent directions were additionally plotted on three
262
rose diagrams according to their respective depositional
263
environments.
264
4.3. Cyclostratigraphical analysis
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An age model was constructed using the existing
266
magnetostratigraphic timescale of the studied section. For each
267
magnetozone, the sedimentation rate was calculated in order to
268
evaluate variations of sediment input into the basin through time.
10
269
A cyclostratigraphical analysis was performed on the
270
frequencies of the parasequences, low- and high-order sequences, to
271
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
275
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
277
were then plotted against the facies rank data and the astronomical
278
target curves.
279
4.4. Faunal analyses
280
Working in Paratethyan basins may introduce a certain
281
ambiguity between palaeontological and sedimentological
282
terminology (Matoshko et al., 2016). These basins registered episodic
283
periods of connectivity and disconnectivity with the open ocean and
284
therefore display lowered salinity environments, where endemic
285
faunas developed though time (Marinescu, 1978; Stoica et al., 2013).
286
Here, the terms ‘brackish water’ and ‘freshwater’ are used to specify
287
water salinity on the basis of palaeontological indicators. The term
288
‘open water’ is used in a sedimentological context in order to
289
describe offshore to shoreface depositional environments.
290
The Slănicul de Buzău section display very rich assemblages of
291
endemic mollusc and ostracod faunas. For this study, both groups
292
were analysed in order to corroborate environmental reconstructions
293
based on sedimentology. Mollusc assemblages were studied from 58
294
sediment samples. Samples of 1000-2000 g were taken throughout
295
the section (Fig. 3a, SBD16-nF). Sample preparation for molluscs
11
296
was performed at the Natural History Museum in Vienna. Samples
297
were cleaned using pneumatic micro-chisels. They were then washed
298
through sieves of 1 mm. The general preservation of the shells was
299
moderate to poor. Shells were finely cracked due to secondary
300
gypsum mineralisation and carbonate crystal growth. The taxonomic
301
identifications follow Wenz (1942) and Marinescu and Papaianopol
302
(1995). Taxonomic revision incorporates results by Nevesskaya et al.
303
(1997, 2001, 2013) and Neubauer et al. (2014) (Supplementary
304
material 2).
305
Ostracod assemblages were studied from 35 sediment samples
306
of 500-1000 g taken throughout the section (Fig. 3a, SBD16-nO).
307
Sample preparation for ostracods was carried out at the Department
308
of Palaeontology at the University of Bucharest. Samples were dried
309
to remove interstitial water from the sediments. Dry samples were
310
subsequently boiled for 30-60 minutes with a sodium carbonate
311
solution for better disaggregation. Samples were washed through
312
several sieves of 63 to 500 μm. The residues were studied under a
313
ZEISS–Stemi SV11 microscope. Pictures of microfaunas were taken
314
with a NIKON digital camera. The general preservation of the
315
ostracods was moderate to poor. Ostracods were often fragmented
316
due to strong diagenetic processes. The taxonomic identifications
317
follow Hanganu (1976, 1985), Hanganu and Papaianopol (1977),
318
Stancheva (1990) and Olteanu (1995) (Supplementary material 2).
12
319
5. Results
320
5.1. Sedimentary facies associations
321
The 835 m thick succession displays a generally regressive
322
trend, superposed by a rhythmic alternation between more distal
323
clays and more proximal sands. Our field observations and
324
subsequent microscope descriptions were compared to well
325
documented sedimentological classifications (Postma, 1990; Miall,
326
2006). We distinguished thirteen lithofacies, formed by distinct
327
sedimentary processes (Table 2, Figs. 4-7). Lithofacies were grouped
328
into eight depositional facies, representing five main facies
329
associations, each of which being related to a distinctive depositional
330
environment.
331
5.1.1. Prodelta facies association
332
Description
333
The first facies association is generally 1 to 5 m thick. It
334
consists of three types of dark-bluish-gray (GLEY2-4/5B) to bluish-
335
gray (GLEY2-5/5B) mudstone that occur successively. There are
336
massive (Fm), laminated (Fl) and lenticular (Fs) mudstones. Massive
337
mudstones (Fm) occur at the base of the prodelta facies association
338
strata. They are generally 0.5 to 1 m thick, but sometimes are absent
339
from this facies association. They display a dark-bluish-gray color
340
(Fig. 4b). They may contain well-preserved in situ brackish water
341
molluscs, such as Euxinicardium olivetum, Pontalmyra tohanensis or
342
Chartoconcha rumana (Supplementary material 3, Fig. 8). This
343
facies is progressively replaced by 1 to 3 m thick, laminated gray
344
mudstones (Fl). These mudstones have mm-scale horizontal
345
laminations of silt (Fig. 4c). They may also contain cm-scale
13
346
horizontal laminations of silt with mm-scale fragments of terrestrial
347
organic material. Upwards, the muddy succession may contain 0.5 to
348
1 m thick, gray mudstones with lenticular bedding (Fs). The
349
lenticular bedding consists of 1 to 5 cm thick isolated lenses made of
350
silt to very fine sand showing trough cross-stratification (Fig. 4d).
351
The sandy layers are occasionally affected by cm-scale convolute
352
bedding. Throughout this facies association, the laminations and the
353
lenses become thicker, more frequent, and composed of coarser
354
sediments toward the top. Palaeocurrent directions were measured on
355
3D trough-cross stratifications present in lenticular-bedding.
356
Measurements in these deposits demonstrate a mean direction of
357
225° (n=17; Fig. 7a). They display a wide range of current directions
358
from 180° to 270°. In addition, deposits occasionally show
359
intercalations of cm-thick beds of gray (GLEY1-6/N) fine to medium
360
grained sandstones (Sfr). Their bases form a wavy, sharp surface,
361
highly perturbed by vertical burrows 3 to 5 cm wide and 5 to 15 cm
362
deep (Fig. 5e). These sandstones are moderately-sorted and are
363
composed of subangular quartz grains with high sphericity. They
364
contain many abraded or broken, reworked brackish and freshwater
365
cardiids, unionids, dreissenids or viviparids.
366
Interpretation
367
This facies association is interpreted as a prodelta environment
368
because of the distal depositional setting and the evidence of distal
369
fluvial input. The massive mudstones (Fm) indicate deposition from
370
suspension in open water. The progressive transition to mm-scale
371
silty laminations (Fl) is related to large fluvial outflows, energetic
372
enough to reach the distal part of the basin. The cm-scale, organic-
373
rich, silty laminations could be related to hyperpycnal flows, 14
374
associated with episodic larger river discharge events (Mulder et al.,
375
2003; Bhattacharya and MacEachern, 2009; Lamb and Mohrig,
376
2009). Upwards, the gradual occurrence of silty lenticular bedding
377
(Fs) is related to wave action and/or winnowing (De Raaf et al.,
378
1977). The sandy beds comprising reworked, abraded and broken
379
shells (Sfg) are thought to illustrate sporadic higher energetic
380
depositional processes occurring in the muddy surrounding
381
environment. Coarser structureless sediments were transported into
382
the basin over long distances during intermittent sand influxes. As in
383
previous studies (Starek et al., 2010; Hampson et al., 2011), these
384
sandstones are interpreted as storm deposits. The coarsening
385
character within this facies association illustrates an increase in the
386
energy of the depositional process and is seen as a shallowing of the
387
environment. Following earlier studies of similar muddy facies
388
associations (Overeem et al., 2003; Olariu and Bhattacharya, 2006;
389
Fielding, 2010), we propose that sediments were deposited in a
390
prodelta setting. The large range of palaeocurrent directions
391
highlights the development of several delta-lobes, feeding a wide
392
prodelta region.
393
5.1.2. Distal delta-front facies association
394
Description
395
The second facies association shows a 0.5 to 5 m thick, regular
396
alternation between cm- to dm-thick layers of mudstones and
397
sandstones. The mudstones are blueish-gray (GLEY2-5/5B). They
398
record mm- to cm-scale horizontal laminations of silts or very fine
399
sands draped by silts and fragments of terrestrial organic material
400
(Fl), as well as lenticular bedding made of cm-scale lenses with
401
trough cross-stratified silt to very fine sand (Fs). They are 15
402
intercalated with layers of grayish-brown sands (2.5Y-5/2). These
403
sands are very fine to fine grained, moderately sorted, and contain
404
low-spherical and subangular grains. The sandstones show three
405
types of cm-scale cross-beddings (Sr, Ss, St). They comprise 10 to 50
406
cm thick climbing ripples (Fig. 5b), 20 to 100 cm thick sigmoidal
407
cross-stratification (Fig. 5c) and 10 to 50 cm thick trough cross-
408
stratification (Fig. 5d). The cross-bedding foresets are commonly
409
draped by mm-scale laminae of fragments of terrestrial organic
410
matter. The sandstones frequently show convolute bedding on a cm-
411
to dm-scale. The sandy beds become thicker, more frequent, and are
412
composed of coarser sediments toward the top of the facies
413
association. As in the previous prodelta facies association, deposits
414
are occasionally interrupted by the same cm-scale beds of gray
415
(GLEY1-6/N) fine to medium grained sandstones (Sfg). They are
416
structureless and composed of well-sorted sediments with highly-
417
spherical and subangular quartz grains. Their bases display the same
418
wavy, sharp, highly bioturbated surface (Fig. 5e). They also contain
419
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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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.
Figure 1 Click here to download high resolution image
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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
