Gilbert-type deltas recording short-term base-level changes: Delta-brink morphodynamics and related...

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This is an Accepted Article that has been peer-reviewed and approved for publication in the Sedimentology, but has yet to undergo copy-editing and proof correction. Please cite this article as an “Accepted Article”; doi: 10.1111/sed.12212

This article is protected by copyright. All rights reserved.

Received Date : 03-Dec-2013

Revised Date : 18-Feb-2015

Accepted Date : 20-Apr-2015

Article type : Original Manuscript

Gilbert-type deltas recording short-term base-level changes: Delta-brink

morphodynamics and related foreset facies

KATARINA GOBO*, MASSIMILIANO GHINASSI† and WOJCIECH NEMEC‡

*Department of Earth Science, University of Bergen, 5007 Bergen, Norway (E-mail: [email protected]); presently: Statoil ASA, Sandsliveien 90, 5254 Sandsli, Norway

†Department of Geosciences, University of Padova, 35131 Padova, Italy

‡Department of Earth Science, University of Bergen, 5007 Bergen, Norway

Associate Editor – J. P. Walsh

Short Title – Delta foreset facies and base-level changes

ABSTRACT

Gilbert-type deltas are sensitive recorders of short-term base-level changes, but the delta-front record

of a base-level rise tends to be erased by fluvial erosion during a subsequent base-level fall, which

renders the bulk record of base-level changes difficult to decipher from the delta-front deposits. The

present detailed study of three large Pleistocene Gilbert-type deltas uplifted on the southern coast of

the Gulf of Corinth, Greece, indicates a genetic link between the delta-front morphodynamic

responses to base-level changes and the delta-slope sedimentation processes. Sigmoidal delta-brink

architecture signifies a base-level rise and is accompanied by a debrite-dominated assemblage of

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delta foreset deposits, thought to form when the aggrading delta front stores sediment and undergoes

discrete gravitational collapses. Oblique delta-brink architecture tends to be accompanied by a

turbidite-dominated assemblage of foreset deposits, which are thought to form when the delta-front

accommodation decreases and the sediment carried by hyperpycnal effluent bypasses the front. This

primary signal of the system response to base-level changes combines further with the secondary

‘noise’ of delta autogenic variation and possible allogenic fluctuations in fluvial discharge due to

regional climatic conditions. Nevertheless, the evidence suggests that the facies trends of delta foreset

deposits may be used to decipher the delta ‘hidden’ record of base-level changes obliterated by fluvial

topset erosion. Early-stage bayhead deltas may be an exception from the hypothetical model,

because their narrow front tends to be swept by river floods irrespective of base-level behaviour and

their subaqueous slope deposits are thus mainly turbidites.

Keywords: Base-level changes, debrites, delta morphodynamics, delta-slope processes, sea-level

changes, turbidites

INTRODUCTION

Gilbert-type deltas (Fig. 1A) — first described by Gilbert (1885) and later named after him — are a

variety of steeply sloping deltas that form where rivers enter a relatively deep body of standing water.

These deltas are particularly common as fjord-head features (Prior & Bornhold, 1988; Syvitski &

Farrow, 1989; Corner et al., 1990) and incised-valley bayhead systems (Postma, 1984; Corner, 2006;

Eilertsen et al., 2006; Garrison & van der Bergh, 2006; Li et al., 2006; Breda et al., 2007; Gobo et al.,

2014a; Leszczyński & Nemec, 2014); their distinctive tripartite architecture (Fig. 1A) comprises a

steeply inclined foreset of subaqueous delta-slope deposits passing into a subhorizontal bottomset of

prodelta deposits and overlain by a horizontal topset of fluvial delta-plain deposits (Barrell, 1912;

Smith & Jol, 1997). The delta brink (Fig. 1A) is a crucial morphodynamic zone for sediment transfer

from the upper delta front, dominated by river effluent regime, to the lower delta front dominated by

basin hydraulic processes and the delta-slope realm dominated by gravitational sediment transport

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(Prior et al., 1981; Massari & Parea, 1990; Prior & Bornhold, 1990; Lønne & Nemec, 2004;

Kleinhans, 2005; Longhitano, 2008).

These deltas have attracted considerable research interest in the last three decades or so,

including modern cases (Prior et al., 1981; Kostaschuk & McCann, 1987; Prior & Bornhold, 1988,

1989, 1990; Corner et al., 1990; Bell, 2009; Saito, 2011) and many ancient examples (Postma, 1984;

Postma & Roep, 1985; Colella, 1988a, 1988b; Postma & Cruickshank, 1988; Colella & Prior, 1990; Ori

et al., 1991; Dart et al., 1994; Dorsey et al., 1995; Massari, 1996; Chough & Hwang, 1997; Sohn et

al., 1997; Nemec et al., 1999; Dorsey & Umhoefer, 2000; Lønne et al., 2001; Lønne & Nemec, 2004;

Ilgar & Nemec, 2005; Mortimer et al., 2005; García-García et al., 2006; Breda et al., 2007; Ford et

al., 2007; Longhitano, 2008; Backert et al., 2010; Eilertsen et al., 2011; Ilgar et al., 2013; Gobo et al.,

2014b), as well as computer modelling (Muto & Steel, 1992; Syvitski & Daughney, 1992; Hardy et

al., 1994; Uličný et al., 2002) and laboratory experiments (Kleinhans, 2005; Rohais et al., 2011; Bijkerk

et al., 2013; Ferrer-Boix et al., 2015). A comprehensive review of the delta-slope processes and facies

was given by Nemec (1990), including the origin of slope chutes, gullies and cross-strata backsets

(Fig. 1A). Particular attention has been given to the sigmoidal and oblique toplap geometries (sensu

Mitchum et al., 1977) of the delta foreset/topset relationship (Fig. 1A). These geometries reflect the

rising or subhorizontal to falling time-distance trajectory (sensu Helland-Hansen & Martinsen, 1996) of

the delta brink during progradation and are attributed to the short-term changes in base level caused

by tectonics and/or fifth-order to sixth-order eustatic cycles, possibly combined with changes in

sediment supply rate (Colella, 1984; Gawthorpe & Colella, 1990; Dart et al., 1994; Massari, 1996;

Soria et al., 2003). Therefore, Gilbert-type deltas are widely considered to be sensitive recorders of

short-term base-level changes (Massari & Colella, 1988; Colella & Prior, 1990; Breda et al., 2007;

Gobo et al., 2014b), with a full awareness that the record may be highly incomplete, because the

sigmoidal toplap signature of base-level rise may be erased during a subsequent base-level fall

(Fig. 1B).

However, little attempt has thus far been made to study the relationship between the delta toplap

geometries and the corresponding foreset facies (see relevant discussion by Helland-Hansen

& Hampson, 2009). The deltaic foresets — with a few notable exceptions (Mastalerz, 1990; Sohn et

al., 1997; Nemec et al., 1999; Lønne et al., 2001; Lønne & Nemec, 2004; Gobo et al., 2014b) — had

seldom been studied systematically in detail on a bed by bed basis and their deposits instead were

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simply lumped together as products of unspecified 'slope avalanches'. Importantly, if there is a

recognizable relationship between the delta-brink trajectory and foreset facies, then the delta foreset

deposits alone might possibly be used to decipher the ‘hidden’ record of base-level changes (Fig. 1B).

The objective of the present study is to address empirically this intriguing issue on the basis of a

number of Pleistocene Gilbert-type deltas exposed on the southern coast of the Gulf of Corinth,

Greece. Segments of delta longitudinal outcrop sections with alternating sigmoidal and oblique toplap

geometries have been studied, with roughly horizontal logs used to analyse the corresponding foreset

deposits. The variation in foreset facies, assessed in statistical terms of their relative thickness

frequency, indicates a diagnostic link with the toplap geometry. A delta strike section is used to

demonstrate lateral uniformity of the architecture and facies anatomy of a delta-brink zone with

sigmoidal toplap, and thereby to justify the use of a single longitudinal section as representative for

delta subaqueous processes. The study as a whole gives new insights into the morphodynamics of

Gilbert-type deltas and reveals a genetic relationship between delta-brink trajectory and foreset facies.

GEOLOGICAL SETTING

The Gilbert-type deltas selected for this study belong to a Plio-Pleistocene syn-rift sedimentary

succession, nearly 3 km thick, uplifted in the footwalls of fault blocks along the southern margin of the

Corinth Rift in central Greece (Fig. 2A). The Gulf of Corinth, ca 105 km long and 30 km wide, is the

submerged part of this active intracontinental rift trending WNW–ESE, whose development

commenced in the Pliocene (Ori, 1989; Briole et al., 2000; Doutsos & Kokkalas, 2001; Leeder et

al., 2008). The rift present-day extension rate is 11 to 16 mm/yr (Briole et al., 2000; Ford et al., 2013),

with an uplift rate of the southern margin of up to 1.5 mm/yr and subsidence rate in the central to

northern part of 2.5 to 3.6 mm/yr (Tselentis & Markopoulos, 1986; Doutsos & Piper, 1990; Collier &

Dart, 1991; Briole et al., 2000; McNeill & Collier, 2004; Lykousis et al., 2007). Fault activity at the

southern margin has been shifting progressively towards the rift axis (Ori, 1989; Sorel, 2000; Rohais et

al., 2007a, 2007b) and is presently localised along the gulf southern coast. The north-dipping faults

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divide the rift margin into blocks, 5 to 10 km long and 5 to 8 km wide, tilted southward at 25 to 30°

(Dart et al., 1994).

The syn-rift sedimentary succession comprises three main lithostratigraphic units (Fig. 2A and B;

Ghisetti & Vezzani, 2004, 2005; Rohais et al., 2007a, 2007b): (i) the Lower Group, composed of fluvio-

lacustrine deposits dated to between 3.6 Ma and 1.5 Ma; (ii) the Middle Group, with deposits of

northward-built giant Gilbert-type deltas and offshore fine-grained turbidites dated to between 1.5 Ma

and 0.7 Ma; and (iii) the Upper Group, represented by onshore colluvial deposits, coastal raised

marine terraces and smaller Gilbert-type deltas dated to be younger than 0.4 Ma.

The spectacular Gilbert-type deltas of the Middle Group have been studied extensively from the

point of view of their spatial-stratigraphic stacking pattern and relationship to the rift-margin faults (e.g.

Ori, 1989; Ori et al., 1991; Seger & Alexander, 1993; Dart et al., 1994; Ford et al., 2007; Rohais et al.,

2008; Backert et al., 2010; Ford et al., 2013), but relatively little high-resolution sedimentological

analysis of these deltas (Ford et al., 2007; Backert et al., 2010; Gobo et al., 2014b) and the Upper

Group deltas (Gobo et al., 2014a) has been conducted. The present detailed study focuses on the

brink-zone architecture and corresponding foreset facies in three selected deltas: the Evrostini and

Ilias deltas of the Middle Group (Fig. 2D) and the Akrata delta of the Upper Group (Fig. 2C).

The Evrostini and Ilias deltas

The Evrostini and Ilias Gilbert-type deltas are uplifted and exposed in a rift-margin block bounded by

the Valimi, Evrostini and Xylokastro faults to the south and the Pirgaki-Mamoussia, Akrata and

Derveni faults to the north (Fig. 2A); their topsets reach an altitude of 1200 m and 700 m, respectively,

and their remarkably thick foreset units indicate northward delta progradation in rift-margin waters up

to 500 m deep. Ori (1989) and Seger & Alexander (1993) originally suggested that the Ilias delta was

younger than the Evrostini delta, but this interpretation was subsequently revised by Rohais et al.

(2007a, 2008), who established that it was the latter delta that prograded directly over the former. The

Olvios River (Fig. 2A) fed these successive deltas by cutting into the pre-rift bedrock composed of

limestone, chert, greenstone and quartzite, but tectonic uplift and block back-tilting eventually caused

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its drainage reversal and delta abandonment. Therefore, there are no major younger deltas in the

Derveni area (Fig. 2A; Seger & Alexander, 1993; Dart et al., 1994).

The studied outcrop section of the Ilias delta is a north-trending cliff in its western part (see locality

7E in Fig. 2D). The outcrop section of the Evrostini delta is in the north-eastern flank of a north-west

trending perched dry valley, ca 2.2 km long and 80 m deep, incised axially in the delta (see locality 7D

in Fig. 2D).

The Akrata delta

The Akrata delta filled a palaeovalley incised in the Platanos delta of the Middle Group (Fig. 2A; Gobo

et al., 2014a). The relay-ramp zone hosting the delta is bounded by the Akrata and Derveni faults to

the south and the East Helike Fault to the north (Fig. 2A). The delta topset reaches an altitude of

about 180 m and the delta foreset is up to 80 m thick, although its transition to bottomset is poorly

exposed. The delta was formed by the antecedent Krathis River, which has presently incised it axially

and is building a modern delta at the coastline (Fig. 2C). The flanking cliffs of the modern river valley

afford good longitudinal outcrop sections of the Akrata delta (see localities 7A to 7C in Fig. 2C). In

addition, a slightly strike-oblique section of the delta brink zone has been studied in the coastal cliff

(see locality 4B in Fig. 2C).

TERMINOLOGY AND METHODS

The terminology for Gilbert-type delta architecture, including toplap geometry, is summarized in

Fig. 1A. The term base level denotes the relative water level of the host basin. The descriptive

sedimentological terminology, including clast fabric notation, is after Harms et al. (1975, 1982) and

Collinson et al. (2006). Following Blikra & Nemec (1998, 2000), the terms ‘debrisflow’ and ‘debrisfall’

are written as single words, in analogy to terms such as ‘mudflow’ and ‘rockfall’. Sedimentary facies

associations are defined as groups of spatially and genetically related facies representing particular

sub-environments of a deltaic system; they are for simplicity labelled with interpretive genetic names,

but their descriptions are separated from interpretations in the text. In the analysis of delta foreset

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deposits, the term ‘facies assemblage’ is used to denote packages of foreset beds that differ in their

facies composition from the adjacent bed packages.

Conventional field methods of sedimentological analysis were used, with detailed bed by bed

logging and an overlay line-drawing of bedding on enlarged outcrop photomosaics. Quasi-horizontal

logs of delta foresets were made around 20 to 50 m below the topset, with the local logging route

dependent on outcrop accessibility. The line drawings, made directly at the outcrops, were the basis

for correlating particular toplap geometries with the corresponding packages of foreset beds. Special

care was taken to mark the boundaries of successive bed packages in the log, with the binocular-

equipped drawing person leading the logging team by means of walkie-talkie communication. The line-

drawing technique combined with vertical logging was used also to study the strike section of delta-

brink deposits.

The facies composition of foreset bed packages in the logs was determined and compared in both

qualitative (graphical plots) and quantitative terms (frequency percentage). The percentage of facies

relative thickness, rather than bed number, has been used, because it is often difficult to distinguish

between amalgamated beds of same facies.

DELTA FACIES ASSOCIATIONS

The deltaic deposits studied are mainly conglomeratic, weakly to moderately cemented, with a

variable amount of sandy matrix. Gravel is polymictic, rich in clasts of limestone, chert, vein quartz and

greenstone. The Ilias delta is the coarsest-grained, with 99 vol % of its foreset composed of

conglomerates and the rest of pebbly sandstones. The Evrostini and Akrata deltas are comparable, as

about 65 vol % of their foresets consists of conglomerates and 35 vol % of pebbly sandstones and

minor sandstones. The detailed logging of delta deposits revealed 15 component facies (Table 1),

which form three main facies associations – representing the delta topset, delta front and steep

subaqueous foreset (Fig. 1A).

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Delta topset facies association

The horizontally bedded topsets of the deltas range in thickness from 15 to 45 m and are visually

similar, but with access for logging in only one outcrop of the Akrata delta (locality 3 in Fig. 2C). The

topset there is up to 25 m thick and consists of lenticular conglomeratic units stacked laterally and

vertically upon one another (Fig. 3C). These units are up to 2.5 m thick and a few tens of metres wide,

and most of them show a fining-upward trend (Fig. 3B to D). Their concave-upward erosional bases

are paved with clast-supported, well-rounded, massive coarse gravel (clast sizes ≤45 cm), locally

reaching 70 cm in thickness and showing a 'rolling' a(t)b(i) fabric indicative of general north-eastward

transport direction (facies Gm in Table 1; Fig. 3C). This facies reoccurs vertically and is overlain by

parallel-stratified, finer-grained sandy conglomerate or pebbly sandstone, with the stratification varying

from subhorizontal to gently inclined (10 to 15°) downstream, sideways or occasionally upstream

(facies Gs in Table 1; Fig. 3B and C). The strata sets are a few decimetres thick, commonly show

upward coarsening or fining, and are wedging out against each other or superimposed upon one

another (Fig. 3B) in a compensational manner (sensu Straub et al., 2009). The stratified gravelly

facies Gs is sporadically overlain by erosional relics of a mottled, faintly laminated to massive sandy

mudstone up to 20 cm thick (facies Fl in Table 1; Fig. 3A and C).

The characteristics of this facies association and its occurrence directly above the delta foreset

indicate deposition in shallow braided-stream distributary channels (Miall, 1985, 1996) of the delta

plain. Facies Gm is interpreted to be channel-floor lags and facies Gs to represent low-relief, mid-

channel longitudinal bars (Boothroyd & Ashley, 1975; Nemec & Postma, 1993) or 'sheet bars' (sensu

Boothroyd, 1972). Some of these bars seem to have been channel bank-attached as side bars. The

coarsening-upward trend of a bar is the signature of frontal progradation, whereas a fining-upward

trend reflects lateral accretion. The dimensions and cross-cutting relationship of channel-fill bodies

indicate laterally-shifting fluvial conduits a few tens of metres wide and mainly less than 2 m deep. The

muddy facies Fl (Fig. 3A) is interpreted to be a slack-water deposit capping an abandoned channel.

Such channel-fill caps were probably more common than presently observed (Fig. 3C), because their

preservation potential in a system of laterally shifting braided channels was very low (see Miall, 1996;

Bridge, 2003).

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Delta-front facies association

The delta-front facies association of the topset/foreset transition is best developed where the toplap

geometry is sigmoidal (Fig. 4A). These deposits have been studied in detail in a coastal strike section

of the Akrata delta (Fig. 4B; see locality 4B in Fig. 2C) in an abandoned quarry at an altitude of about

130 m. The east–west outcrop section is ca 50 m wide and 7 m high, slightly oblique to the mean

direction of delta progradation (Fig. 2C). This facies association is characterized by an overall upward

coarsening (Fig. 4B), with an upward decrease in seaward bedding inclination reflecting sigmoidal

toplap geometry (Fig. 4A). Several minor post-depositional faults occur in the outcrop eastern part

(Fig. 4C).The lower delta-front deposits are dominated by disorderly ‘scour and fill’ features, whereas

the upper delta-front deposits consist mainly of mounded units stacked upon one another in a shingled

manner (Fig. 4B and C). Worth noting is the lateral persistence of these characteristics, which means

that they should be equally recognizable in any random longitudinal dip section of the delta.

Upper delta-front deposits

The upper part of delta-front facies association in the studied case reaches about 5 m in thickness and

consists of interbedded conglomerates and gravelly sandstones (Fig. 4B and C; log 2d in Fig. 4I). In

longitudinal outcrop section, the deposits form tabular or mounded bed packages 1 m to 2 m thick,

inclined seaward at 4 to 10° (Fig. 4A). In transverse section, the packages have a lateral extent of

several tens of metres and form lenticular mounds stacked upon one another erosionally or non-

erosionally in a compensational manner (Fig. 4B and C). The conglomerate beds of facies Giw

(Table 1, Fig. 4D) are up to 30 cm thick, have a clast-supported, sand-filled to openwork texture and

show distinct layers rich in either subspherical or seaward-dipping flatter clasts. The associated

gravelly sandstone beds of facies Ssw (Table 1, Fig. 4D) are 20 to 50 cm thick, showing plane-parallel

stratification with scattered pebbles and fine pebble stringers. In the uppermost part, below the

overlying topset, many of the conglomerate beds have concave-upward erosional bases, are coarser-

grained and poorer stratified and sorted; the associated sandstone beds are also coarser and richer in

gravel.

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The depositional mounds are interpreted to be mouth bars built by the frictional effluent of delta

distributary channels (Wright, 1977). Their compensational stacking is attributed to the autogenic

switching of delta distributary channels (Massari & Parea, 1990; Kleinhans, 2005; Longhitano, 2008;

Saito, 2011), with the erosional contacts representing extreme river floods or mouth-bar gravitational

collapses. The mouth bars are thought to have been heavily wave-worked, as their component facies

resemble beachface deposits (cf. Bluck, 1967, 2010; Dunne & Hampton, 1984; Kleinspehn et

al., 1984). The erosional, coarser and poorer-sorted conglomerate beds in the uppermost part of delta

front are considered to be river-lain deposits that filled the shallow outlets of bar cross-cutting feeder

channels (see Wright, 1977) and suffered minimal reworking by sea waves.

Lower delta-front deposits

The lower part of delta-front facies association, up to about 6 m thick in the outcrop (Fig. 4B and C; log

2bc in Fig. 4I), shows a wider range of facies and an architecture dominated by scour and fill features.

The bedding in longitudinal section is generally well-defined, with a seaward inclination of 10 to 20°,

common up-dip pinch-outs and prominent scours up to 4 m deep (Fig. 4A). The scours in transverse

section are up to 20 m wide (Fig. 4C) and filled with the gravelly sandstones and sand-supported

conglomerates of facies Gms (Table 1; Fig. 4E and G), commonly alternating with the stratified pebbly

sandstones and sandy conglomerates of facies Gsp, 6 to 80 cm thick (Table 1; Fig. 4E and H).

Smaller scours, up to 60 cm deep and 1 m wide, are filled with facies Gms and/or Gsp (Fig. 4H) and

occur locally also within the infill of large scours. Soft-sediment deformation is common, mainly due to

loading, and some of the thick massive beds of facies Gms show internal listric shear bands (Fig. 4E;

log 2b in Fig. 4I). The lenticular conglomeratic beds are commonly enveloped by fine-grained silty

sandstones of facies Ssp (Table 1), up to 60 cm thick, with seaward ripple cross-lamination and minor

plane-parallel stratification (Fig. 4F).

These lower delta-front deposits are attributed to an alternation of high-stage and low-stage

frictional river effluent (facies Gsp and Ssp, respectively) with episodic debrisflows due to delta-front

collapses (facies Gms; cf. Chough & Hwang, 1997). The river bedload accumulated as mouth bars

clearly tended to be re-mobilized by gravitational slumping and be transported further downslope by

sediment gravity flows (Prior & Bornhold, 1988, 1990; Massari & Parea, 1990; Nemec, 1990;

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Postma, 1990). The delta-brink gravitational collapses would occur whenever a mouth-bar slope

exceeded ca 35° (see Syvitski & Farrow, 1989; Nemec, 1990; Hardy et al., 1994), and would probably

be enhanced by such factors as the loading by storm waves, seismic shaking and the internal

streaming of river effluent and wave backwash through the porous mouth-bar sediment (Sims, 1973;

Hempton & Dewey, 1983; Nemec, 1995). Most scours are thought to be the head parts of gullies

formed by localized collapses and downslope escape of sediment (Postma, 1983; Nemec, 1990), but

some may be chutes cut by the stream-flood effluent turning into a hyperpycnal flow (Prior et al., 1981;

Postma, 1984; Kostaschuk & McCann, 1987; Postma & Cruickshank, 1988; Prior & Bornhold, 1988,

1990; Nemec, 1990; Saito, 2011). The sediment-gravity flows released from these numerous scours

resulted in deposition on the delta slope, whereas the scours themselves were gradually plugged by

debrisflows or buried by stratified mouth-bar deposits.

Delta foreset facies association

None of the three deltas has its bottomset exposed, and hence the full thicknesses of their foresets

cannot be measured. The exposed foreset thicknesses in the studied outcrops range from ca 50 m in

the Akrata delta to ca 80 m in the Evrostini delta and ca 100 m in the Ilias delta. The foreset beds are

inclined seaward at 20 to 30° and their facies (Table 1, Fig. 5) have been grouped into three main

generic categories: debrisflow deposits (facies Gms and Sm), turbidites (facies Gsa, Smg and Ssr)

and debrisfall deposits (facies Go). Debrisflow deposits are the most abundant facies in the studied

foreset sections, except for the Akrata-3 section (Fig. 7C) which represents an early bayhead stage of

the Akrata delta (Gobo et al., 2014) and is dominated by turbidites (Fig. 6). Facies Sxu and Sru

(Table 1) are rare and volumetrically insignificant (less than 1 vol %); they are described and

interpreted here, but are disregarded in subsequent quantitative analysis.

Debrisflow deposits

These deposits, referred to also as debrites, are mainly massive, non-graded or coarse-tail inversely

graded, matrix-supported to clast-supported conglomerate beds 10 cm to 120 cm thick (facies Gms,

Fig. 5), occasionally amalgamated into composite beds up to 850 cm in thickness. Non-graded beds

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tend to be thicker than the graded ones, and sand-supported conglomerate beds are most common.

The gravel fraction is moderately sorted, comprising subrounded to rounded pebbles and small

cobbles, whereas matrix is a mixture of medium-grained to very coarse-grained sand and granules.

Beds are tabular or mounded in shape, with sharp non-erosional bases and also sharp tops,

commonly erosional (Fig. 5). Most beds show flow-parallel alignment of large clasts. Some of the

thickest beds show diffuse listric shear-banding accentuated by clast a-axis alignment (Fig. 5A). The

subordinate massive, non-graded beds of facies Sm (Table 1, Fig. 5) are 5 cm to 70 cm thick and

commonly overlain by sandstone with plane-parallel stratification.

Facies Gms is attributed to cohesionless debrisflows (sensu Nemec & Steel, 1984; Blikra &

Nemec, 2000) generated by delta-front collapses and characterized by a low to moderate rate of

internal shear strain (frictional shear regime, sensu Drake, 1990). The coarse-tail inverse grading is

due to a vertically differential rate of shear strain and the en route loss of coarsest clasts by their

settling out from the strongest-sheared lower part of the flow (Nemec & Postma, 1991). The listric

shear-banding indicates internal thrusting and signifies debrisflows whose frontal braking was still

accompanied by the flow-body movement (Massari, 1984; Nemec, 1990). Facies Sm is attributed to

similarly cohesionless, high-viscosity sandy debrisflows, with the stratified sandy caps indicating an

accompanying or closely following low-density turbidity current (sensu Lowe, 1982), possibly resulting

from the dilution of debrisflow top by water entrainment (Nemec, 1990; Falk & Dorsey, 1998). Some

beds of facies Sm, especially thinner ones, may be attributed to co-genetic debrisflows spawned by

the collapsing of high-density turbidity currents (see Postma et al., 1988; Vrolijk & Southard, 1997;

Mulder & Alexander, 2001).

Turbidites

This category of foreset deposits is represented by facies Gsa, Smg and Ssr (Table 1). The sandy

conglomerate and pebbly sandstone beds of facies Gsa are the most common, showing sharp bases

and crude to distinct plane-parallel stratification with vertical grain-size fluctuations (Fig. 5). Strata sets

range from 3 cm to 70 cm in thickness and are commonly amalgamated into cosets up to nearly 6 m

thick. The sandstone facies Smg (Table 1) is relatively rare and its massive beds, 18 to 70 cm thick,

show normal grading with scattered pebbles in the basal part and vague plane-parallel stratification at

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the top (Fig. 5C). Even less common are the silty sandstones of facies Ssr (Table 1), whose cross-

laminated layers drape coarser-grained facies (Fig. 5F) and are mainly 3 to 12 cm thick, but

sporadically up to 160 cm; they occasionally also show local plane-parallel stratification and contain

isolated pebbles or small cobbles. Ripple cross-lamination indicates downslope sediment transport

direction.

These foreset facies represent tractional deposition by turbulent sediment-gravity flows, and hence

are interpreted as turbidites. The considerable thicknesses and fluctuating grain size of facies Gsa

beds suggest sustained (long-duration) pulsating currents, considered to be river flood-generated

hyperpycnal flows (Nemec, 1990, 1995; Mulder & Alexander, 2001). The graded massive beds of

facies Smg, with stratified top parts, suggest surge-type high-density turbidity currents (sensu

Lowe, 1982), possibly formed by the scouring action of river hyperpycnal effluent or a rapid turbulent

dilution of delta-front debrisflows (Falk & Dorsey, 1998). Facies Ssr indicates deposition by low-density

turbidity currents (sensu Lowe, 1982), probably small hyperpycnal flows. Solitary large clasts would

roll down easily on a steep sandy substrate even under a relatively weak current (Allen, 1983; Isla,

1993), which may explain the occurrence of such isolated clasts in this facies. Some spheroidal large

clasts may have rolled down on the steep delta slope due to the sheer pull of gravity (debrisfall, sensu

Holmes, 1965; Nemec, 1990; see the next section).

Debrisfall deposits

These are openwork coarse conglomerates (facies Go in Table 1), found only in the middle to lower

part of delta foreset. Their lenticular beds are 5 cm to 70 cm thick, composed of well-rounded

subspherical pebbles and cobbles up to 23 cm in size (Fig. 5A and F). Cobbles tend to be

concentrated in the bed downslope part, whereas pebbles dominate in the thinner upslope ‘tail’ of the

deposit, which is also commonly fining both upwards and in the upslope direction. The beds are non-

erosional and their shape is generally adjusted to an uneven substrate. Many isolated or clustered

outsized pebbles and cobbles entrapped between the beds of other facies (Fig. 5B) are also

considered to represent this category of deposits.

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These openwork gravel deposits are interpreted to be products of debrisfall processes

(Holmes, 1965; Nemec, 1990). The large clasts are thought to have moved rapidly downslope under

the force of gravity, in small groups or in isolation, by freely rolling, sliding and bouncing against one

another. The coarse gravel must have been derived from the delta front, where it accumulated by

wave action and as bedload lag at channel outlets to be released for freefall from head-scarps formed

by mouth-bar collapses (Nemec, 1990; Sohn et al., 1997; Nemec et al., 1999).

Subordinate other facies

Beds of the subordinate facies Sxu (Table 1, Fig. 5E) are backsets of sandy to gravelly cross-strata

dipping upslope at up to 30° relative to the delta foreset bedding surfaces. These sporadic backsets

are up to 5 m thick and mainly isolated (see log horizon B in Fig. 7E), but occasionally multiple (see

log interval E–F in Fig. 7A). They form the infill of downslope-trending trough-shaped scours, where

the backset in its downslope part typically abuts against the rear relief of a debrisflow mound of facies

Gms. The backsets are interpreted to be slope chute-fill deposits formed by a supercritical low-density

turbidity current subject to hydraulic jump against the topographic relief of debrisflow mound (Postma,

1984; Nemec, 1990; Massari, 1996; Nemec et al., 1999).

Rare is also facies Sru (Table 1, Fig. 5D), which forms isolated fine-grained silty sandstone layers

up to 6 cm thick, locally bioturbated, showing ripple cross-lamination indicative of an upslope current.

The sporadic occurrences of this facies were found only in the outcrop section Akrata-2, where the

logging route climbed relatively high, to less than 30 m below the delta topset (Fig. 7B). The origin of

this facies is attributed to a weak action of tidal flood currents, whose flow power in microtidal settings

could be amplified by topographic confinement (e.g. Longhitano & Nemec, 2005; Corner, 2006;

Longhitano et al., 2012). Facies Sru would appear to be the only recognizable record of tidal influence

in the studied deltas.

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DELTA-BRINK TRAJECTORY AND FORESET FACIES CHANGES

The foreset–topset contact in the selected longitudinal sections of Gilbert-type deltas alternates

between sigmoidal and oblique (Fig. 7), which indicates short-term changes in the delta-brink

trajectory (see Helland-Hansen & Martinsen, 1996). The targeted detailed logging of delta foresets

allowed packages of foreset beds to be correlated with a particular type of delta-brink trajectory. Two

varieties of foreset bed packages are distinguishable: a debrite-dominated facies assemblage (DFA;

Fig. 8A and B) and a turbidite-dominated facies assemblage (TFA; Fig. 8C and D). Their relationship

to the delta toplap geometry of delta foreset is shown in Fig. 9 and described below.

Sigmoidal toplap geometry reflects an ascending, normal–regressive delta-brink trajectory (see

Helland-Hansen & Martinsen, 1996; Helland-Hansen & Hampson, 2009), rising at up to 24° in the

present cases (Figs 7E and 9). The seaward shoreline shifts recorded by sigmoidal delta-brink

deposits are in the range of 20 to 45 m in the Akrata delta (Fig. 7A to C), 35 to 50 m in the Evrostini

delta (Fig. 7D) and 120 to 140 m in the Ilias delta (Fig. 7E). The corresponding upward climb of delta

shoreline is in the range of 2 to 6 m, 7 to 10 m and 7 to 15 m, respectively, reflecting the magnitude of

short-term base-level rises. The foreset bed packages associated with sigmoidal toplap are generally

facies assemblages DFA, dominated by debrisflow deposits (Figs 8 to 10). It is also worth noting that

the relatively thick debrisflow beds tend to be associated with a gentler-rising brink trajectory, whereas

thinner debrisflow beds seem to dominate in cases of a steeper-rising trajectory (Fig. 9).

An oblique toplap geometry of the delta foreset/topset contact indicates a stationary or

descending, forced-regressive brink trajectory (see Helland-Hansen & Martinsen, 1996; Helland-

Hansen & Hampson, 2009), falling in the present cases at an angle of up to 9° (Fig. 7). The seaward

shoreline shifts associated with an oblique toplap are in the range of 40 to 230 m in the Akrata delta

(Fig. 7A to C), 60 to 75 m in the Evrostini delta (Fig. 7D) and 100 to more than 120 m in the Ilias delta

(Fig. 7E). The corresponding drop of the delta shoreline is, respectively, in the range of 2 to 8 m, 1 to

2 m and 10 to 15 m. The foreset bed packages associated with oblique toplap are mainly facies

assemblages TFA, dominated by turbidites (Figs 8 to 10). It is only the Akrata-3 section of an early-

stage bayhead delta (Gobo et al., 2014a) which – despite its sigmoidal toplap geometry – shows an

irregular alternation of foreset bed assemblages DFA and TFA (Fig. 9) and an overall high proportion

of turbidites (Fig. 10).

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DISCUSSION

Gilbert-type deltas are an extreme case of clinoform depositional system with a well-defined narrow

brink zone, which renders them some of the most sensitive coastal recorders of short-term base-level

changes (Massari & Colella, 1988; Colella & Prior, 1990; Lønne & Nemec, 2004; Breda et al., 2007;

Gobo et al., 2014b). The base-level rise in such advancing systems is signified by a sigmoidal toplap

geometry due to the rising brink trajectory, whereas the base-level stillstand or fall are recorded as an

erosional oblique toplap due to the horizontal or descending brink trajectory (Fig. 1A; see Mitchum et

al., 1977; Helland-Hansen & Martinsen, 1996; Massari, 1996; Uličný et al., 2002; Soria et al., 2003;

Helland-Hansen & Hampson, 2009). Consequently, the sigmoidal brink-zone record of a base-level

rise can readily be erased by fluvial incision during a subsequent base-level fall and be

unrecognizable (Fig. 1B). The key issue addressed by the present study was whether this obliterated

record of base-level changes can possibly be deciphered from the delta foreset facies.

The detailed study of five suitable sections of three large Gilbert-type deltas at the southern coast

of the Corinth Rift indicates that the sigmoidal delta-brink architecture is associated with debrite-

dominated foreset facies assemblages DFA, whereas the oblique delta-brink architecture tends to be

associated with turbidite-dominated foreset facies assemblages TFA (Figs 8 and 9). This former

relationship is attributed to the increased accommodation and storage of sediment at the delta brink

during a base-level rise, leading to frequent discrete collapses in the form of debrisflows (Fig. 4E to H),

whereas the latter relationship can be attributed to the deficit of delta-brink accommodation and an

increased bypass of sediment across the delta front by means of turbidity currents (mainly

hyperpycnal flows) during a base-level stillstand or fall. The base-level changes in the present case

were probably due to a combination of Pleistocene eustasy and rift-margin tectonics (Rohais et al.,

2011). This interpretation concurs with most of the previous interpretive notions of the

morphodynamics of Gilbert-type delta front (Colella, 1984, 1988b; Syvitski & Farrow, 1989; Massari &

Parea, 1990; Nemec, 1990; Massari, 1996; Uličný et al., 2002; Mortimer et al., 2005; Breda et al.,

2007; Longhitano, 2008). The coarse-grained deltas are characteristically devoid of mud, because this

finest-grained sediment fraction is persistently winnowed from the narrow delta-front zone by waves

and tends to be expelled seaward as a hypopycnal suspension plume (Syvitski & Farrow, 1983;

Syvitski, 1989; Nemec, 1990, 1995). This is why the accumulation rate of prodeltaic mud in estuarine

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basins with coarse-grained bayhead deltaic systems is generally very high (Syvitski et al., 1985,

1987).

In the studied Corinthian deltas, the foreset bed packages with a sigmoidal toplap tend to be

dominated by debrites (DFA), whereas some of those with an oblique toplap appear to include both

facies assemblages TFA and DFA (Fig. 9). It is suggested that these mixed oblique-topped bed

packages may in reality include erosionally-truncated primary sigmoidal packages DFA (see Fig. 1B).

Notably, the oblique bed packages that directly precede the sigmoidal ones — and hence are most

certain to have formed during a base-level fall or stillstand — are facies assemblages TFA (Figs 9 and

10). The high proportion of foreset turbidites can be attributed to a fairly persistent sediment bypass of

the delta front due to a deficit of accommodation, with a highly limited transient storage of sediment at

the delta front and hence less frequent collapses in the form of debrisflows. The notion of a river

bedload bypassing the delta front is evidenced by the abundance of turbidites (hyperpycnites) and the

associated slope chutes filled with facies Sxu (cf. Prior et al., 1981; Nemec et al., 1999; Lønne et

al., 2001; Gobo et al., 2014b).

If this reasoning is correct, a debrite-dominated foreset bed package with an oblique toplap (Fig. 9)

may be an original sigmoidal package that formed during a base-level rise and was erosionally

truncated during the subsequent base-level fall (Fig. 1B). Such foreset bed packages might thus be

regarded as the delta ‘hidden’ record of base-level rises. On this interpretive premise, it is suggested

that the debrite-dominated and turbidite-dominated facies assemblages of delta foreset deposits

(Fig. 8) can possibly be used to infer short-term base-level changes in the development history of an

ancient Gilbert-type delta (Fig. 11).

The association of relatively thick foreset debrisflow deposits with gently rising (up to 2°) delta-

brink trajectories and the thinner debrisflow deposits with steeper trajectories (Fig. 9) suggests that the

rate of base-level rise coupled with the rate of sediment supply may play a significant role. A high rate

of delta-front aggradation would appear to favour sediment sloughing by frequent small-volume

collapses, whereas a low rate of aggradation to favour occasional large-volume collapses.

Accordingly, it is hypothesized that the thicker debrisflow deposits may indicate volumetrically larger

but less frequent delta-front collapses.

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The studied deltas differ considerably in thickness (host-water depth), which suggests that the

inferred physical coupling of delta-slope sedimentation with base-level changes may be scale-

indifferent. However, the evidence from an early-stage Akrata bayhead delta (section Akrata-3 in Figs

7C, 9 and 10; Gobo et al., 2014a) suggests that, for a given range of river discharges, the delta front

excessively confined by valley tends to be swept by hyperpycnal river-flood effluent irrespective of the

base-level behaviour, whereby the delta slope is dominated by turbidites. The narrow, early-stage

bayhead Gilbert-type deltas would thus appear to be considerably less sensitive to low-magnitude

base-level changes when it comes to the mode of delta-slope processes.

The suggested facies model (Fig. 11) focuses on the role of base-level changes in controlling the

delta-brink morphodynamics and the corresponding prevalent mode of subaqueous sediment

transport on a Gilbert-type delta slope. However, the changes in base level are obviously not the only

factor controlling the hydraulic regime and morphodynamics of a delta front. The variable proportion of

component facies in the foreset bed assemblages DFA and TFA (Fig. 9) and the fluctuating grain size

of deposits (Fig. 8) are thought to reflect short-term spatial-temporal variation in fluvial sediment

delivery to the delta front due to the system autogenic variability (Kleinhans, 2005; Longhitano, 2008;

Saito, 2011) and possibly also such other allogenic factors as regional climate. The coarse-grained

deltas lack microfossils, but the deposition time span of such individual systems is anyway far beyond

biostratigraphic resolution (e.g. see Leszczyński & Nemec, 2014). For example, the Pleistocene

Middle Group in the study area shows the record of several successive Gilbert-type deltas with

intervening episodes of fluvial valley incision, but its bulk time span is only 800 ka. The time span of a

single delta may be in the order of 100 to 200 ka, and its studied short segment would then represent

only a small fraction of this time. The short-term changes in delta foreset facies thus cannot be

attributed to the Pleistocene long-term regional climatic changes (cf. Bottema & van Zeist, 1981;

Nemec & Postma, 1993; Roberts & Wright, 1993; Landman et al., 1996; Pope et al., 2008), although

they may possibly reflect climate seasonality (even though this effect has thus far been little evidenced

in the region). Because of the relatively short time scale of a delta growth, it is also impossible to

determine as to which changes in the delta-brink architecture were forced by high-order eustatic

cycles and which by the rift-margin tectonics, or by a combination of these allogenic factors.

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Because no debrisflow or turbidity current is generated along the entire delta front, it is

questionable as to how representative a single dip section is for a Gilbert-type delta system. The

example strike section (Fig. 4) shows that the delta-brink architecture, although varied on a local

scale, is laterally uniform, showing the same range of gravitational collapses and scour and fill

processes of down-slope sediment transfer. This evidence implies that, for a given delta-brink

configuration, also the range of the corresponding delta-slope processes will be laterally similar. In

other words, the rate of the short-term autogenic changes along the delta front far exceeds the

effective rate of delta progradation and hence any longitudinal cross-section of the delta can be

expected to be equally representative of the foreset facies composition. The sample datasets collected

in the present study are thus considered to be representative for the deltas. However, future research

should verify whether the link between delta-brink trajectory and foreset facies indicated by this study

is generic or specific only to the three Corinthian deltas.

CONCLUSIONS

Gilbert-type deltas are widely regarded as some of the most sensitive coastal recorders of short-term

base-level changes, with the base-level rise reflected in the delta-brink rising trajectory and sigmoidal

toplap geometry, and the base-level stillstand or fall reflected in the delta-brink subhorizontal or falling

trajectory and an oblique erosional toplap geometry. However, the delta-front record of base-level

rises may be erased by fluvial erosion during subsequent base-level falls, which renders the overall

record of base-level changes difficult to decipher from the delta-front deposits alone.

The present pilot study of three large Gilbert-type deltas on the southern coast of the Gulf of

Corinth indicates a genetic link between the delta-front morphodynamic responses to short-term base-

level changes and the delta-slope sedimentation processes. It is suggested that a debrite-dominated

foreset facies assemblage forms during a base-level rise, because the aggrading delta front then

tends to store sediment and undergoes frequent gravitational collapses, whereas a turbidite-

dominated facies assemblage forms during a base-level stillstand or fall, when the delta-front

accommodation is at a minimum and sediment tends to be flushed downslope by erosional

hyperpycnal flows. The delta-system autogenic variability and the allogenic impact of regional climate

come further into play, causing inevitable ‘noise’ in the delta-slope facies record. However, evidence

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from the present study seems to be sufficiently compelling to suggest that the delta foreset facies may

be used to decipher the ‘hidden’ record of base-level changes. An exception may be the narrow early-

stage bayhead deltas, whose slim front tends to be swept by hyperpycnal flood-stage flows

irrespective of the low-magnitude base-level changes and whose slope facies assemblage is thus

dominated by turbidites (hyperpycnites).

The study results bear attractive implications for a high-resolution sequence stratigraphy and point

to the importance of detailed sedimentological facies analysis. Future research on a broader range of

Gilbert-type deltas should verify whether the dynamic facies model suggested by the present study is

generic or merely case-specific.

ACKNOWLEDGEMENTS

The field study was a part of the first author's doctoral research project funded by the University of

Bergen. Nicolas Backert, Valeria Bianchi, Rob Gawthorpe and Eivind Sjursen are thanked for field

assistance and useful discussions. The authors appreciate the constructive reviews by Anna Breda,

Antonio Cattaneo, Fernando García-García, Young Sohn, David Uličný and an anonymous reviewer.

Editorial comments from J. P. Walsh helped further to improve the manuscript.

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FIGURE CAPTIONS

Fig. 1. (A) Schematic longitudinal cross-section of a Gilbert-type delta, depicting its characteristic

tripartite architecture and other common features (compiled from Bell, 2009; and Gobo et al., 2014b).

The formation of such deltas reflects a high basin/river depth ratio; no scale is given, as the delta

thickness depends on the basin accommodation and may range from a few metres to a few hundred

metres. (B) Schematic cartoon portraying the growth of a Gilbert-type delta subject to short-term base-

level changes, with a sigmoidal toplap formed during base-level rise (cases 1 and 3) and an oblique

toplap formed base-level stillstand or fall (cases 2 and 4). Note that the sigmoidal brink-zone

architectural record of base-level rise tends to be erased by fluvial incision during a subsequent base-

level fall (see case 4).

Fig. 2. (A) Geological map of the study area showing the main faults and the onshore distribution of

pre-rift bedrock and syn-rift sedimentary units. Map based on Rohais et al. (2007a) and a study by the

present authors of the Akrata area (Gobo et al., 2014a). The inset map of the Aegean region shows

the location of the Gulf of Corinth and the study area. The inset black frames refer to the detailed

maps of the study subareas shown below the figure. (B) Stratigraphy of the syn-rift sedimentary

succession on the southern side of the Gulf of Corinth; modified from Rohais et al. (2008). (C) and (D)

Detailed maps of the study subareas, indicating the location of the studied outcrop sections (inset

small frames); the numbers refer to subsequent figures.

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Fig. 3. Delta topset facies association in the Akrata delta (see locality 3 in Fig. 2C; sedimentary facies

as in Table 1). (A) Close-up view of stratified channel-fill deposits (facies Gs) attributed to longitudinal

braid bars; the mottled mudstone (facies Fl) is an erosional relic of abandoned channel slack-water

deposit. (B) Fining-upward channel-fill deposits composed of the laterally intercalated wedges of

facies Gs; transport direction roughly towards the viewer. (C) Detailed log of the topset deposits, with

palaeocurrent measurements indicating fluvial transport towards the NNE; letter symbols: CB –

channel base, CU – upward coarsening, FU – upward fining. (D) Close-up view of vertically stacked

fluvial palaeochannels; note the coarse gravelly lag (facies Gm) overlain by fining-upward stratified

channel-fill deposits (facies Gs).

Fig. 4. Delta-front facies association in a sigmoidal toplap segment of the Akrata delta (see locality 4B

in Fig. 2C; facies labels as in Table 1). (A) Longitudinal outcrop section roughly parallel to the direction

of delta progradation, showing vertically-stacked mouth-bar deposits. (B) Panoramic view and (C)

overlay line-drawing of transverse outcrop section; note the location of detailed logs, the inset frames

indicating photographic details, the general upward coarsening of the delta-front succession and the

differing depositional architecture of the lower and upper parts. (D) Close-up detail of mouth-bar

deposit, showing alternation of facies Giw and Ssw. (E) Close-up detail of a debrisflow deposit of

facies Gms, with coarse-tail inverse grading and listric shear bands, overlain by a stratified deposit of

facies Gsp. (F) Conglomeratic debrisflow lenses (facies Gms) enveloped by cross-laminated silty

sandstones (facies Ssp). (G) Slope gully scour at the transition to delta foreset, filled with debrisflow

deposits; the hammer is 30 cm long. (H) Small and large scours filled with facies Gsp and Gms. (I)

Detailed sedimentological logs of the delta-front deposits; depositional inclination disregarded.

Fig. 5. Outcrop details of delta foreset deposits (facies labels as in Table 1). (A) Debrisflow

conglomerates of facies Gms intercalated with the debrisfall conglomerates of facies Go; note the

coarse-tail inverse grading and listric shear bands in the thicker debrisflow beds. (B) Alternating

conglomeratic facies Gms, Gsa and Go. The hammer in (A) and (B) is 30 cm long. (C) Stratified and

normally graded sandy turbidites (facies Smg) overlain by a debrisflow deposit (facies Gms); the

measuring stick shown is 32 cm. (D) Fine-grained sandstone bed of facies Sru with ripple cross-

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lamination indicating an upslope transport direction, attributed to tidal flood current; the coin for scale

is ca 2 cm. (E) Turbidite Gsa overlain by debrisflow deposit Gms and truncated by a slope chute filled

with backset deposit Sxu; the delta foreset dip is away from the viewer, obliquely to the left; the

notebook for scale is 21 cm long. (F) Debris-flow deposits of facies Gms intercalated with the

openwork debrisfall deposits of facies Go and a tidal sandstone layer of facies Ssr; the measuring

stick shown is 23 cm long.

Fig. 6. Histograms showing the relative thickness proportion of the three main genetic facies

categories of delta foreset deposits in the five studied outcrop sections of Gilbert-type deltas (see

location in Fig. 2).

Fig. 7. Panoramic views of the five outcrop sections studied (left-hand column) and their bedding

architecture highlighted by line-drawing, with the logging routes and correlative marker surfaces (right-

hand column). Note the alternating oblique and sigmoidal geometry of the delta topset/foreset contact,

and the ascending (red line) and descending (blue line) delta-brink trajectories with the corresponding

angle of inclination. The white lines indicate the logging routes, with the key marker surfaces labelled

with letter symbols.

Fig. 8. Example portions of the sedimentological logs of delta foreset deposits (depositional dip

disregarded), showing the debrite-dominated assemblages (DFA; highlighted in red) and the turbidite-

dominated assemblages (TFA; highlighted in blue). Detailed facies composition of the assemblages is

indicated by the chequer-plots at right-hand margin of the logs. The log locality labels refer to the

outcrops in Figure 7. The letters at the log right-hand margin refer to the marker bedding surfaces

indicated in Figures 7 and 9.

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Fig. 9. Facies chequer-plots (as in Fig. 8) of the delta foreset logs indicated in Figure 7. The debrite-

dominated assemblages (DFA) facies assemblages are highlighted in red and the turbidite-dominated

assemblages (TFA) assemblages are shown in blue. The letters at the log left-hand margin refer to the

marker bedding planes indicated in Figure 7.

Fig. 10. Histograms summarizing the bulk relative thickness proportion of the three main genetic

facies categories in the oblique and sigmoidal foreset bed packages in the studied delta sections

(Fig. 7).

Fig. 11. Schematic cartoon showing the suggested link between the delta foreset facies and base-

level changes. (A) A sigmoidal delta-front geometry and debrite-dominated foreset facies assemblage

(DFA) form during a base-level rise. (B) An oblique delta-front geometry and turbidite-dominated

foreset facies assemblage (TFA) form during the subsequent base-level stillstand or fall. (C) A new

relative sea-level rise leads to the formation of sigmoidal delta-front geometry and debrite-dominated

foreset facies assemblage (DFA). (D) A subsequent fall of base level causes fluvial incision, with the

deposition of another TFA assemblage of foreset facies and formation of oblique toplap geometry.

Only the recognition of DFA and TFA facies assemblages in the delta foreset may now help to

recognize the previous short-term base-level changes experienced by the prograding delta.

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Table 1. An overview of the main sedimentary facies of Gilbert-type deltas distinguished in the present study.

Facies letter codes are modified from Miall (1985); G – gravel, S – sand, F – fine-grained deposit (silt or mud).

Facies associations

Facies Main descriptive characteristics Interpretation

Delta topset deposits

(subhorizontal)

Gm Massive, clast-supported cobble to boulder conglomerates up to 70 cm thick, paving erosional surfaces. Well-rounded clasts with a ‘rolling’ a(t)b(i) fabric

Tractional deposition of bedload gravel as pavement in braided-stream channels (Nemec & Postma, 1993; Miall, 1996)

Gs Matrix to clast-supported, parallel-stratified pebbly sandstones or sandy conglomerates with horizontal to gently inclined (≤15°) strata. Coarsening or fining-upward strata sets stacked in a compensational manner into cosets up to 180 cm thick

Tractional deposition of sand and gravel as longitudinal bars ('sheet bars') in braided-stream channels (Boothroyd, 1972; Boothroyd & Ashley, 1975; Nemec & Postma, 1993)

Fl Mottled mudstones, up to 20 cm thick, with an admixture of fine sand and local faint lamination. Occur sporadically as erosional relics on top of facies Gs

Slack-water deposition in abandoned stream channels (Miall, 1996; Bridge, 2003)

Delta-front deposits

(gently seaward inclined)

Giw Clast-supported conglomerates forming sheet-like beds ≤30 cm thick, with common seaward-dipping imbrication of disc and blade clasts.

Swash-dominated deposition by storm waves on delta beachface (Bluck, 2010)

Ssw Sheet-like beds of planar parallel-stratified gravelly sandstone, 20 to 50 cm thick

Deposition by fair-weather waves with high orbital velocities on delta beachface (Bluck, 2010)

Gsp Planar parallel-stratified pebbly sandstones or sandy conglomerates with frequent grain-size fluctuations on a thickness scale of 6 to 80 cm

Tractional deposition on mouth-bar slope by frictional high-stage river effluent (Wright, 1977)

Ssp Silty sandstone beds 2 to 60 cm thick, with planar parallel stratification and isolated gravel clasts, passing

Tractional deposition on mouth-bar slope by frictional low-stage river

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down-dip into seaward ripple cross-lamination; stratification locally obscured by bioturbation

effluent (Wright, 1977)

Gms Sand to clast-supported massive conglomerates, non-graded or coarse-tail inversely graded, forming solitary or amalgamated beds 6 to 120 cm thick with non-erosional bases and occasional listric shear bands

Deposition by low-mobility cohesionless debrisflows subject to a low to moderate-rate shear strain (frictional shear regime, sensu Drake, 1990)

Delta foreset deposits

(steeply seaward inclined)

Gms Sand to clast-supported massive conglomerates, non-graded or coarse-tail inversely graded, forming solitary or amalgamated beds 10 to 850 cm thick with non-erosional bases and occasional listric shear bands

Cohesionless debrisflows subject to a low to moderate-rate strain (frictional shear regime, sensu Drake, 1990)

Sm Massive, non-graded pebbly sandstone beds, 5 to 70 cm thick, commonly weakly planar parallel-stratified at the top

Sandy debrisflow accompanied or followed by low-density turbidity current; or a co-genetic debrisflow spawned by basal collapse of high-density turbidity current (Postma et al., 1988; Mulder & Alexander, 2001)

Go Isolated large subspherical clasts or openwork conglomerate lenses, 5 to 70 cm thick, with cobbly downslope 'heads' and upslope-fining pebbly 'tails'

Deposition by debrisfall (sensu Holmes, 1965; Nemec, 1990)

Gsa Planar parallel-stratified pebbly sandstones or sandy conglomerates with frequent grain-size fluctuations on a thickness scale of 3 to 70 cm, forming amalgamated strata sets up to 3 m in thickness

Tractional deposition by low-density turbidity current (hyperpycnal flow, sensu Lowe, 1982)

Smg Massive, normally-graded pebbly sandstone beds, 18 to 70 cm thick, commonly passing upwards into faintly planar parallel-stratified sandstone

Deposition by high-density turbidity current (sensu Lowe, 1982)

Ssr Silty sandstone beds 3 to 12 cm thick, with planar parallel stratification and isolated gravel clasts, passing down-dip and/or upwards into seaward-directed ripple cross-lamination; stratification locally obscured by bioturbation

Tractional deposition by low-density turbidity current (hyperpycnal flow, sensu Lowe, 1982)

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Sxu Solitary backsets of sandy/gravelly cross-strata dipping upslope at ca 30° relative to the foreset bedding, filling trough-shaped scours 13 to 500 cm deep

Slope chute-fills formed by low-density turbidity current subject to hydraulic jump (Nemec, 1990)

Sru Fine-grained sandstone layers, up to 6 cm thick, showing upslope-directed ripple cross-lamination

Deposition by weak tidal-flood currents (Corner, 2006)

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Gilbert-type deltas recording short-term base-level changes: Delta-brink morphodynamics and related...

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Transcript of Gilbert-type deltas recording short-term base-level changes: Delta-brink morphodynamics and related...