Sequence analysis of early Saalian glacial lake deposits (NW Germany): evidence of local ice margin...

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Sedimentary Geology 165 (2004) 223–251

Sequence analysis of early Saalian glacial lake deposits

(NW Germany): evidence of local ice margin retreat

and associated calving processes

Jutta Winsemann*, Ulrich Asprion, Thomas Meyer

Institut fur Geologie und Palaontologie, Universitat Hannover, Callinstrabe 30, Hanover D-30167, Germany

Abstract

This paper presents a detailed analysis of high-resolution facies architecture of glaciolacustrine ice-margin deposits, which

formed at the southern margin of the Scandinavian ice shield. The ice margin depositional systems are characterised by coarse-

grained deltas and subaqueous fans, which are formed by stacked transgressive–regressive sequences, recording an overall lake

transgression, interrupted by minor short-term lake-level falls. The delta complexes at the northern margin of glacial Lake

Rinteln are thought to reflect a relatively stable position of the ice margin in front of the Weser Chains mountain ridges. The

onset of delta progradation probably represents a halt in ice advance and a related high sediment supply period. The sedimentary

facies and depositional architecture resemble those of nonglacial Gilbert-type deltas, except for the deposition of glacial debris.

The main delta progradation is recorded from the highstand systems tract, when high meltwater and sediment discharge

occurred during the melt season. A forced regression during opening of outlets led to the formation of subaerially exposed

sequence boundaries and the erosion of the highstand systems tract. Deeply incised channels (incised valleys) were filled during

the subsequent transgression.

At the eastern lake margin, subaqueous fan deposits reflect an unstable ice-front that was rapidly retreating and subject to

periodic calving, which resulted in the generation of floating icebergs, dumping ice-rafted debris and ploughing into subaqueous

fan deposits. The rapid retreat of this eastern ice margin is interpreted to result from the overall lake-level rise.

The fan-stacking pattern and bulk geometry were determined by the ice-front fluctuations, the shifting of meltwater outlets and

the short-term lake-level fluctuations. Depositional processes reflect the character of sediment supply and distance from the ice

margin. Ice-proximal upper-fan deposits are characterised by coarse-grained gravel, deposited by debris flows that document

continuous discharge. Slump and slide deposits are related to steep depositional slopes. Themid-fan contains basinward fining and

thinning deposits of quasi-steady and surge-type high- and low-density turbidity currents, indicating more ice-distal and periodic

deposition. The outer-fan deposits mainly consist of surge-type low-density turbidites and glacial debris dumped by icebergs.

The forcing parameters governing the development of depositional sequences in both delta and fan settings were lake-level

fluctuations, sediment yield rates and physiography. The depositional sequences were deposited on a time scale of seventh- and

eighth-order high-frequency cycles (101–102 years). In this short time span, lake-level fluctuations were on the order of 120 m;

therefore, accommodation space was largely controlled by lake level and subsidence can be ignored.

D 2004 Elsevier B.V. All rights reserved.

Keywords: Pleistocene; Glaciolacustrine deposits; Subaqueous ice-contact fan; Gilbert-type deltas; Sequence stratigraphy; Ice scour

0037-0738/$ - see front matter D 2004 Elsevier B.V. All rights reserved.

doi:10.1016/j.sedgeo.2003.11.010

* Corresponding author.

E-mail address: [email protected] (J. Winsemann).

J. Winsemann et al. / Sedimentary Geology 165 (2004) 223–251224

1. Introduction

Glaciolacustrine ice-margin deposits are important

and well-studied palaeogeographic and palaeoclimatic

archives and record the history and dynamics of

glacier termini in lacustrine basins (Ashley et al.,

1985; Eyles and Clark, 1988; Fyfe, 1990; Mastalerz,

1990; Postma, 1990; Brodzikowski and Van Loon,

1991; Ashley, 1995; Martini and Brookfield, 1995;

Lemons and Chan, 1999). Glaciolacustrine ice-contact

systems are rather complex depositional settings and

only a few studies include detailed analysis of high-

resolution facies architecture (e.g. Clemmensen and

Houmark-Nielsen, 1981; Martini, 1990; Mastalerz,

1990; Martini and Brookfield, 1995; Sadolin et al.,

1997).

Early Saalian glacial lake deposits are widespread

in northern Germany and have been deposited in

glacial lakes at the southern margin of the Scandina-

vian ice shield. The ice margin depositional systems

Fig. 1. Location map of the study area in NW Germany and a palaeogeogra

grained delta and subaqueous fan systems.

are characterised by coarse-grained deltas or subaque-

ous fans, which are formed by stacked transgressive–

regressive sequences. Glacial Lake Rinteln, first

named by Spethmann (1908), is located south of the

Weser Chains in northwest Germany (Fig. 1). Various

coarse-grained deltas and subaqueous fans were built

out into this lake. Deltas were fed by proglacial

streams discharging from the ice margin into the lake.

Glaciofluvial detritus carried to the lake via tunnels

near or at the base of the ice formed subaqueous fan

deposits well below the surface of the lake (Winse-

mann and Asprion, 2001; Winsemann et al., 2003).

During its initial stage, glacial Lake Rinteln stood at c.

55 m a.s.l. The lake level rose some 120 m to form a

high water surface at c. 175 m a.s.l (Fig. 2). The lake-

level rise was very fast and probably took place within

a few tens to a hundred years, inferred from varve

deposits in the basin centre. The reconstructed lake-

level curve shows minor lake-level falls, which are

interpreted by seasonal variations of water discharge

phic reconstruction of the glacial Lake Rinteln and associated coarse-

Fig. 2. Reconstructed lake level curve of glacial Lake Rinteln

(modified after Jarek, 1999; Winsemann and Asprion, 2001).

J. Winsemann et al. / Sedimentary Geology 165 (2004) 223–251 225

or drainage of the lake by the opening of outlets

(Jarek, 1999; Winsemann and Asprion, 2001).

This paper presents a detailed analysis of high-

resolution facies architecture of glaciolacustrine ice-

margin deposits of the glacial Lake Rinteln. These

subaqueous ice-margin deposits are well exposed in

sand and gravel pits and have previously been de-

scribed by several authors who generally assumed a

subaerial formation (Naumann, 1922, 1927; Grupe,

1925, 1930; Naumann and Burre, 1927; Stach, 1930;

Grupe et al., 1933; Luttig, 1954, 1960; Attig, 1965;

Wortmann, 1968; Miotke, 1971; Seraphim, 1972,

1973; Hesemann, 1975; Rausch, 1977; Merkt, 1978;

Deutloff et al., 1982; Rohm, 1985; Wortmann and

Wortmann, 1987; Rakowski, 1990; Wellmann, 1990,

1998; Groetzner, 1995; Deters, 1999). The origin and

significance of these deposits are discussed and relat-

ed to glacier termini dynamics and associated calving

events.

2. Sequence stratigraphy in glacially influenced

basins

Sequence stratigraphy is a well-established tool for

correlating sedimentary successions and has been

widely used in marine and lacustrine basins (Posa-

mentier and Vail, 1988; Van Wagoner et al., 1988,

1990; Posamentier et al., 1992; Lemons and Chan,

1999). The Emme delta and Coppenbrugge subaque-

ous fans were deposited on a time scale of seventh- and

eighth-order high-frequency cycles (101–102 years).

In this short time span, lake-level fluctuations were on

the order of 120 m; therefore, accommodation (space

made available for potential sediment accumulation;

cf. Jervey, 1988) was largely controlled by lake level

and subsidence can be ignored.

Only a few studies have been carried out using

sequence stratigraphy in glaciolacustrine basins (e.g.

Martini and Brookfield, 1995; Brookfield and Martini,

1999). In glacial lakes, glacier advance usually corre-

lates with lake-level changes and accommodation

space, water level and sediment injection points can

vary independently due to opposing and delayed

effects of glacial isostacy, glacial eustacy and glacial

advance and retreat (Martini and Brookfield, 1995;

Brookfield and Martini, 1999). Problems in applying

the sequence stratigraphic concepts on glacial systems

have been discussed in detail by Brookfield and

Martini (1999). Accommodation space and sediment

injection points in glacial basins are controlled not

only by relative water level, but also by the position of

the front of the glacier. During high water levels, the

glacier injection point may be underwater at the base

of the slope. During low water levels, the injection

point of the glacier may be on land. Since the lake level

is controlled by the position of drainage outlets, water

levels may change dramatically, abruptly and indepen-

dently of the glacial input point. The basin can empty

or fill with water rapidly, with only slight changes in

glacier position. In this case, accommodation space

has no relation to the glacier injection point.

In ‘‘lake-level systems’’ (cf. Brookfield andMartini,

1999), deposition is controlled by water level and

sediment supply from streams in the same way as lake

level controls deposition in conventional sequence

stratigraphy since the sediment input point is a shore-

line. Thus, within one lake basin, sequence stratigraphy

needs little modification, although water-level changes

may be more abrupt and of greater magnitude. In the

‘‘glacier input-point system’’ (cf. Brookfield and Mar-

tini, 1999), deposition is controlled by the sediment

input point at or near the end of the glacier. If the glacier

terminates on land, then the sequences are basically

controlled by lake-level systems. But if the glacier

terminates underwater, then the position of the sedi-

ment input point can fluctuate independently of lake

level. In the Exxon model, a relatively rising water

level corresponds to maximum flooding surfaces cul-

minating in highstand systems tracts, and relatively

dropping water levels correspond to unconformities

culminating in lowstand systems tracts. But rising lake

levels often correspond with glacial advances into a


basin, as the glacier isostatically depresses the land and

blocks drainage outlets. This causes a shift in the locus

of proglacial sedimentation to deeper waters and cor-

responds to the lowstand fans of the Exxon model.

Thus, lowstand systems tracts for the glacier input point

system correlate with highstand systems tracts for

deltas of the water-level system and there, input points

fluctuate independently. It is thus essential to keep

systems controlled by relative water levels and glacier

input points separate.

Where glaciers terminate on land, standard Exxon

sequence stratigraphy can easily be applied to adja-

cent lakes. In such cases, the sediment injection point

in the basin is directly related to water level. The basal

sequence boundary of the lake-level system forms as

lake level drops during glacier retreat, when outlets

are opened. Unblocking these outlets may cause

catastrophic drainage and almost instantaneous and

very large irregular drops in lake level. Thus, each

lake-level sequence will probably start and end with

very marked erosional surfaces and incised valleys of

type 1 sequence boundaries (Posamentier and Vail,

1988). The lowstand systems tract of the lake-level

system develops during and after the drop in lake

level, and in some places, deep incised valleys may be

cut through the delta plain and lowstand fans may

develop at the end of these valleys. The transgressive

systems tract marks a rise in relative water level due to

glacier advance causing blocking of drainage outlets.

Condensed fine-grained sediments may abruptly over-

lie the lowstand wedges and fill any valleys incised

during lowstand. Diamict slumps, stream or shoreline

sands may mark the transgression in incised valleys

passing upwards into increasingly distal rhythmites.

The highstand systems tract corresponds with the

maximum height of the lake. During this equilibrium

phase in glacier development, sediment supplied from

glaciers may build out deltas in places, separated by

areas dominated by continued condensed sedimenta-

tion with ice-rafted debris.

Where glaciers terminate underwater, sediment in-

put points can fluctuate independently of lake level.

Sediment may be supplied in by subglacial, englacial

or supraglacial meltwater flows at the ice front and by

rainout from a floating glacier. According to Brook-

field and Martini (1999), the sequence boundaries and

systems tracts for this glacier input-point systems

equate entirely with the lake-level highstand. Keeping

the lake level at highstand, input points of the glacier

system vary according to the position and nature of the

ice front. As a glacier retreats out of the basin, a series

of retreating fining-upwards subaqueous fans forms.

The resulting fining-upwards section would normally

be interpreted as a transgressive systems tract caused

by rising relative water levels. The sequence boundary

of this input-point sequence is the ice-scoured uncon-

formity, and its top (which looks like the maximum

flooding surface) is simply the sediment-starved area

furthest from the ice-front. When the glacier readvan-

ces into the basin, it may deposit a coarsening-upwards

subaqueous outwash section capped by an erosion

surface. Sequence boundaries of lake-level systems

therefore may correlate with glacial retreat deposits in

glacier input-point systems, and sequence boundaries

of the glacier input system may correlate with trans-

gressive systems tracts of the lake-level system. Fur-

thermore, local fluctuating ice fronts through ice-front

calving may produce successive subaqueous fans,

overlain by thin mud drapes, which could be misinter-

preted as a glacier retreat by a rising lake level. Only

where the lake level drops below the glacier injection

point, the formed erosional surfaces (sequence bound-

ary) can be correlated with sequence boundaries of the

lake-level system.

3. Emme delta

The Emme delta complex is about 2 km long, 1.5

km wide and 70 m thick and overlies glaciolacus-

trine mud or Jurassic basement rocks, forming a

steep dipping ramp surface. The delta deposits are

exposed in various gravel pits (Fig. 3), at an altitude

of 95–165 m, which allow a detailed reconstruction

of the facies architecture. The tripartite structure of

the exposed sedimentary successions with well-de-

veloped bottomsets, foresets and topsets indicates

Gilbert-type deltas (Fig. 4). A description and inter-

pretation of lithofacies is given in Table 1. The

terminology for gravel characteristics is after Walker

(1975). The fabric notation uses symbols a and b for

the clast long axes, with indices (t) and (p) denoting

axis orientation transverse or parallel to flow direc-

tion, and index (i) denoting axis imbrication. The

notation of turbidites, Tabcd, refers to the Bouma

divisions (cf. Bouma, 1962).

Fig. 3. Location map of the open-pit outcrops of the Emme delta

complex; the topographic contour values are in metres.


3.1. Facies associations

3.1.1. Bottomsets

The bottomsets consist of 1–2.5 m thick planar

parallel-stratified granule to medium-grained sand

(F11) that pass updip into steeply inclined gravelly

foreset beds (Fig. 4C). The planar parallel-stratified

sand is interpreted to result from surge-type turbulent

low-density gravity flows (Tb), triggered by the

release of limited sediment volumes by discrete fail-

ures of the delta’s upper slope and/or brink zone. The

finer-grained, low-density parts of these currents have

apparently bypassed the delta slope and deposited

their load in the delta toe and prodelta zone. The

occurrence of scattered pebbles at the bed boundaries

is interpreted to result from coeval debris fall process-

es (Nemec et al., 1999). In the toe part of the delta

foreset, the planar parallel-stratified sand forms

mound-shaped successions, whose convex-upward

tops suggest a series of depositional lobes, similar to

those described by Postma and Cruickshank (1988)

and Nemec et al. (1999), which apparently coalesced

with one another into an aggradational ramp (Fig.

4C). The multiple delta toe and lobes suggest that the

delta was advancing by the alternating episodes of

slope-base aggradation and progradation. The toe of

the delta slope aggraded in response to the slope

steepening, causing intense sediment sloughing by

means of chutes and occasional large-scale failures

(Nemec et al., 1999).

3.1.2. Foresets

The foresets mainly consist of 5–15 m thick steeply

dipping medium- to thick-bedded matrix- and clast-

supported gravel (F7, F8), often showing an a(p) a(i)

steep-clast fabric, dipping in an upslope direction (Fig.

4F). The absence of current-produced structures and

the occurrence of steep-clast fabric suggest a steep

slope with gravity-driven sediment transport (Postma

and Cruickshank, 1988; Martini, 1990; Massari and

Parea, 1990; Nemec, 1990). The vast majority of beds

have a coarse-grained sandy matrix nearly avoid of

mud, and these debris flows were probably cohension-

less (Nemec and Steel, 1984), controlled mainly by the

sediment’s frictional strength, which would explain

their low mobility and preferential deposition on the

delta’s upper to middle slope (Nemec et al., 1999). In

dip-parallel sections, the debris-flow beds are exten-

sive and fairly tabular. Floating outsized clasts are

common and beds are often inversely graded, indicat-

ing a loss of the largest clasts from the lower, faster-

shearing and rheologically weakest part of the debris

flow (Naylor, 1980). Many beds show upslope-dip-

ping internal shears, listric or sigmoidal in shape,

marked by pebble stringers or sandy bands nearly

avoid of gravel (Fig. 4F). These features are thought

to be syndepositional thrusts. The occurrence of an

a(p) or a(i) clast fabric is attributed to laminar shear

(Nemec, 1990). Subordinate, thin- to medium-bedded

pebbly sand beds occur, which are massive or inverse-

ly graded (F9) or show normal grading or planar

parallel stratification (F10), interpreted to have been

deposited by sandy debris flows or surge-type low-

density turbidity flows, respectively (cf. Shanmugam,

2000; Nemec et al., 1999).

Sediment was supplied to the delta front by braided

glaciofluvial streams. Material that bypassed the braid

plain avalanched downslope as debris flows and

stopped by freezing when the slope diminished. The

occurrence of debris flows corresponds with the

combination of high-bedload events of fluvial dis-

charge to the delta front (Nemec, 1990; Prior and

Bornhold, 1990) or slope instability events triggered

Fig. 4. Architectural elements of the Emme delta complex. Lithofacies types (F1–11) refer to Table 1. (A) Delta foreset, unconformably overlain

by a channelised topset, Fell open pit. Note gravel lag at the base of the channel. (B) Close-up view of (A), showing large-scale trough cross-

bedding of topset channel fill (F1). (C) Coarse-grained sandy bottomset (F11), overlain by steeply dipping gravelly foreset beds (F8). The

foreset beds are dipping southwestwards and the delta toe rises in that direction, indicating aggradation of the bottomset zone during delta-front

progradation. The mound-shaped geometries of bottomset successions (arrows) suggest a series of coalescing depositional lobes; Prange open

pit. (D) Sheetlike topset facies (F5), consisting of sand, silt and mud alternations with ripple cross-lamination, Prange open pit. (E) Steeply

dipping matrix-supported gravel and pebbly sand, Fell open pit (F7). (F) Steep-clast fabric within foreset bed, indicating laminar shear during or

immediately after the flow’s stop (F8), Fell open pit.


Table 1

Lithofacies classification and architectural elements, Emme delta

Lithofacies and sedimentary structures Geometry Interpretation

Topsets

Channel deposits

F1 Pebbly sand with large-scale

trough-cross-bedding. Bed

thickness up to 2 m.

Lenticular

(5–10 m thick and

up to 70 m wide)

Bedload deposition of braided streams

F2 Medium-bedded medium- to

coarse-grained sand with

planar parallel stratification

Lenticular, up to

1.5 m thick


F3 Clast-supported gravel with

large-scale trough cross-bedding

Lenticular, up to

0.5 m thick



blocks up to 1.3 m across

Lenticular (0.5–1.5 m

thick and 70 m wide)

Gravel lag at channel floor

F5 Alternations of thin- to

medium-bedded fine- to medium-

grained sand, silt and mud with

plan parallel lamination, ripple

cross-lamination, wave-ripple

cross-lamination and dropstones

Lenticular (up to 3 m

thick and 60 m wide)

Low-energy unidirectional and oscillatory flows

with coeval dumping of ice-rafted debris.

Glaciolacustrine sedimentation in incised valley.

Interchannel deposits

F6 Thin- to medium-bedded (1–20 cm)

alternations of fine- to medium-grained

sand, silt, mud and clay with plan

parallel stratification, ripple cross-

lamination and plan-parallel lamination

Sheetlike (0.5–1 m thick,

up to 30 m wide)

Unconfined low- to high-energy

flows on delta plain

Foresets

F7 Matrix-supported gravel with normal

or inverse grading. The matrix consists

of middle- to coarse-grained sand.

Outsized clasts up to 50 cm across.

Bed thickness between 10 and 40 cm;

sharp bed contacts. Long axes of large

clasts are often oriented parallel to dip

and may show an a(p) a(i) fabric.

Wedge (5–15 m, up to

150 m long). High-angle

bedding (10–30j)

Deposition from debris flows; the steep-clast

fabric indicates laminar shear during or

immediately after the flow’s stop (Nemec, 1990)


normal or inverse grading. Sharp bed

contacts. Clasts up to 50 cm in diameter.

The matrix consists of middle- to

coarse-grained sand. Bed thickness

between 10 and 60 cm. Clasts often show

a steeply imbricate fabric with the a-axes

dipping in an upslope direction. Long

axes are mainly oriented parallel to dip.

Wedge (5–15 m thick,

up to 50 m long).

High-angle bedding

(10–30j)

Deposition from debris flows; the steep-clast

fabric indicates laminar shear during or

immediately after the flow’s stop (Nemec, 1990)

(continued on next page)


Lithofacies and sedimentary structures Geometry Interpretation

F9 Massive or inversely graded pebbly sand.

Bed thickness between 2 and 20 cm.

Long axes of large clasts are often

oriented parallel to dip and may show

an a(p) a(i) fabric.

Wedge (5 m thick,

up to 50 m long).

High-angle bedding

(10–20j)

Deposition from sandy debris flows

(Shanmugam, 2000)

F10 Normally graded or planar

parallel-stratified

pebbly sand. Bed thickness

between 2 and 20 cm.

Wedge (5 m thick, up

to 50 m long).

High-angle

bedding (10–20j)

Deposition from surge-type turbidity flows

(Shanmugam, 2000; Lowe, 1982)

Bottomsets

F11 Planar parallel-stratified fine- to

coarse-grained sand with scattered

pebbles. Bed thickness

between 2 and 10 cm.

Mound (0.5 m thick,

up to 5 m wide).

Horizontal to low-angle

bedding (0–3j).

Deposition from surge-type low-density

turbidity flows and coeval debris fall

(cf. Nemec et al., 1999). Mounds are interpreted

as individual depositional lobes.

Table 1 (continued)


by the shear of overriding flows or by gravity (Nemec

et al., 1999; Plink-Bjorklund and Ronnert, 1999). The

finer-grained sandy material moved to the basin floor

where it was deposited from turbidity flows.

3.1.3. Topsets

The gravelly foresets are overlain by 0.5–2 m

thick subhorizontal sheet-like sand, silt, mud alter-

nations and/or channelised sand and gravel (Fig. 4A,

B and D). The sheet-like strata show ripple cross-

lamination and planar parallel lamination/stratifica-

tion (F6), interpreted to result from unconfined low-

and high-energy flows on the delta plain. Channel-

fill deposits (F1–F5) are 5–15 m thick and up to 70

m wide and mainly consist of large-scale trough

cross-bedded and planar parallel-stratified pebbly

sand, indicative of 3-D dunes and upper flow-regime

deposits of braided glaciofluvial channels. Subordi-

nate thin- to medium-bedded sand, silt, mud alter-

nations with ripple cross-lamination, planar parallel

lamination and wave-ripple cross-lamination occur,

interpreted to result from low-energy unidirectional

and oscillatory flows. Within these fine-grained

channel deposits, occasionally, dropstones can be

observed.

3.2. Open-pit Fell section

The basal succession of the Emme delta complex is

exposed in the open-pit Fell (Fig. 3) at an altitude of c.

108–124 m (Fig. 5A and B). The lower foreset is up to

15 m thick and can be traced laterally for about 150 m.

Individual foreset beds are 10–60 cm thick and

remarkably constant in thickness along dip. The beds

consist of matrix- or clast-supported sandy gravel,

which are massive or show inverse or normal grading

(F7, F8). Outsized clasts often show an upslope

dipping steep-clast fabric, interpreted to result from

laminar shear (Nemec, 1990). Bed contacts are usually

sharp. The foreset beds steeply dip (18–30j) into

westerly directions. This foreset is unconformably

overlain by a c. 7-m-thick foreset, dipping (10j–20j) into southerly directions. Foreset beds are 2–20

cm thick and mainly consist of massive pebbly sand

(F9) with outsized clasts, oriented parallel to dip.

Occasionally, normally or inversely graded pebbly

sand beds can be observed. Subordinate thin- to

medium-bedded intercalations of planar parallel-strat-

ified medium- to coarse-grained sand can be observed

(F10), interpreted to have been deposited from low-

density turbidity flows (Tb).

Both foresets are unconformably overlain by a

channel-fill, 5 m thick and c. 70 m wide. The basal

channel-fill consists of a 1-m-thick clast-supported

gravel lag (F4), which passes upward into pebbly

sand with large-scale trough cross-bedding (F1). The

succession is interpreted to be a glaciofluvial channel-

fill. The erosive gravel lag is apparently a channel-

floor deposit, whereas the overlying cosets of trough

cross-strata represent 3-D dunes. The channel-fill

deposits are overlain by a c. 5-m-thick gravelly foreset

(F7, F8) dipping (10j–25j) towards the southwest.

Fig. 5. (A) Sketch of the open-pit outcrop Fell with location of the measured section. (B) Measured section of the open-pit outcrop Fell. For

legend, see Fig. 6.


Fig. 6. Sketch of the open-pit outcrop Prange with location of measured sections.


3.3. Open-pit Prange section

The upper part of the delta complex is exposed in

the open-pit Prange (Fig. 3) at an altitude of 138–158

m. At the base of the open pit, 2.5-m-thick sandy

bottomsets are exposed, which pass updip into coarse-

grained gravelly foreset beds (Fig. 6). The foreset

beds are steeply dipping (25j–30j) southwestwards


and the delta toe rises in that direction (Fig. 4C),

indicating aggradation of the bottomset zone during

delta-front progradation. The mound-shaped geome-

Fig. 7. (A) Sketch of the open-pit outcrop Im Teufelsbad with location of th

Teufelsbad. For legend see Fig. 6.

tries of bottomset successions suggest a series of

coalescing depositional lobes, forming an aggrada-

tional ramp. Foreset beds consist of matrix- or clast-

e measured section. (B) Measured section of the open-pit outcrop Im


supported gravel (F7, F8, Table 1), frequently show-

ing inverse or normal grading and an upslope dipping

a(p) a(i) steep-clast fabric. The gravelly foreset is up

to 8 m thick and deeply incised by a c. 10-m deep and

60-m-wide channel, interpreted to represent an incised

valley. The basal part of this incised valley-fill con-

sists of c. 5-m-thick mud, silt and sand alternations

with wave-ripple cross-lamination, planar-ripple

cross-lamination and dropstones (F5). These fine-

grained glaciolacustrine valley-fill deposits are ero-

sively overlain by large-scale trough cross-stratified

and planar parallel-stratified sand and pebbly sand

(F1, F2), interpreted as glaciofluvial bedload deposits

with numerous phases of erosion and reactivation.

Laterally, these channelised sand and pebbly sand

pass into sheet-like topset deposits (Figs. 4D and 6).

3.4. Open-pit Im Teufelsbad section

The uppermost part of the Emme delta complex is

exposed in the open-pit Im Teufelsbad at an altitude of

c. 150–160 m (Fig. 3). The open pit has already been

refilled and exposures are poor. The sedimentary

succession starts with a 3-m-thick gravelly foreset,

which is bounded by an upper erosional surface (Fig.

7A and B). Foreset beds are 10–50 cm thick and

consist of matrix- or clast-supported gravel, steeply

dipping (15–25j) into southerly directions. Normal or

inverse grading is commonly developed and larger

Fig. 8. Stackening pattern and sequence stratigraph

clasts often show an a(p) a(i) steep-clast fabric,

dipping in an upslope direction (F7, F8). The upper

erosional surface is overlain by a 0.5–1.5 m thick

clast-supported gravel-lag, which passes upwards into

1.5-m-thick coarse- to medium-grained sand with

planar parallel stratification (Fig. 7B). The erosive

gravel lag is interpreted to represent a channel-floor

deposit, whereas the overlying planar parallel-strati-

fied sand is interpreted as a glaciofluvial bedload

deposit. These deposits are overlain by 6-m-thick

matrix-supported disorganized sandy gravel of a

young subaerial debris flow. Locally, at the base of

this unit, a 0.2–1 m thick boulder lag can be ob-

served, which is traceable for about 10 m (Fig. 7A).

3.5. Large-scale stackening pattern and depositional

sequences

The Emme delta complex consists of six vertically

stacked Gilbert-type deltas, which are interpreted to

represent eighth-order high-frequency depositional

sequences (Fig. 8). Sequence boundaries are indicated

by erosional unconformities, produced by stream

entrenchment during lake-level falls. The incised

channels and valleys were filled with lake and glacio-

fluvial deposits during the subsequent transgression.

The fluvial incision resulted in steep unstable slopes

along the channel-margins, and a series of slides and

slumps has occurred along these slopes and could be

ic interpretation of the Emme delta complex.


observed during excavation. A minimum number for

the lake-level fall is obtained from the erosion depth

of the glaciofluvial channels, which must be near or

above lake level landward of the shoreline (Postma,

1995). The main delta progradation occurred during

lake-level highstands, when high meltwater and sed-

iment discharge occurred during the summer.

The large-scale stacking pattern of the Emme delta

complex indicates an overall lake-level rise, interrup-

ted by minor short-term lake-level falls and closely

resembles the Rf type of Postma (1995). A similar Rf

lake-level control on delta architecture and facies has

been described by Fernandez et al. (1993). Opening of

outlets probably led to short-term lake-level falls, a

forced regression and the formation of subaerially

exposed sequence boundaries and the erosion of the

highstand systems tract (cf. Posamentier and Vail,

1988; Van Wagoner et al., 1988, 1990; Posamentier

Fig. 9. Location map of the open-pit outcrops of the Coppenbrugg

et al., 1992). The delta was finally abandoned due to

the northward retreat of the ice front.

4. Coppenbrugge subaqueous fan

The Coppenbrugge fan complex is situated on the

eastern margin of glacial Lake Rinteln and is about 10

km long and up to 10 km wide (Fig. 9). The fan

sediments form approximately NW–SE striking sed-

iment ridges and are exposed in various gravel pits

(Fig. 9), at an altitude of 90–170 m. They overlie

glaciolacustrine mud and a diamicton, interpreted to

represent a basal till (Deters, 1999) and lack any

subaerial glaciofluvial or distributary delta-plain com-

ponents. Clasts consist mainly of resedimented fluvial

material (95%), previously deposited by the River

Leine and River Weser or originated from the adjacent

e fan complex; the topographic contour values are in metres.

Table 2

Classification of facies associations, Coppenbrugge subaqueous fan

Facies association Geometry Interpretation of depositional processes

FA1 Massive clast-supported gravel (60–95%) with

bed thickness between 10 and 40 cm. The

matrix consists of fine- to coarse-grained

sand and bed contacts are mainly sharp;

occasionally, erosive bed contacts occur.

Larger clasts can be oriented parallel to dip

and show a steeply imbricate clast fabric with

the a-axes dipping in upslope direction.

Intercalations of thin- to medium-bedded

inversely or normally graded gravel, massive

or normally graded pebbly sand and massive

or plan parallel-stratified sand amount to 5–30%.


up to 50 m long).

High-angle bedding

(16–34j). Lenticularin chutes and channels.

The massive clast-supported gravel and pebbly

sand with sharp bed contacts indicate deposition

from noncohesive debris flows (Shanmugam, 2000);

the steep-clast fabric indicates laminar shear during

or immediately after the flow’s stop (Nemec, 1990).

The normally graded gravel, pebbly sand and

stratified sand are interpreted to result from

surge-type high- and low-density turbidity flows

(Bouma, 1962; Lowe, 1982; Nemec et al.,

1999; Plink-Bjorklund and Ronnert, 1999).

Usually found in the proximal upper fan and both

feeder and distributary channels.

FA2 Thin- to thick-bedded clast-supported massive

or inversely graded gravel (40%), alternating with

thin- to thick-bedded massive, normally or

inversely graded pebbly sand (33%), massive,

normally graded, planar parallel-stratified, ripple

cross-laminated sand (20%) or large-scale trough

cross-bedded coarse-grained sand (7%). Subordinate

thin beds of massive silt or silty sand (1%) occur.


up to 50 m long).

Low- to high-angle

bedding (5–25j).Lenticular in

chutes and channels.

The massive or inversely graded gravel and pebbly

sand with sharp bed contacts indicate deposition from

noncohesive debris flows (Nemec, 1990; Shanmugam,

2000). The normally graded gravel, pebbly sand and

stratified sand are interpreted to result from surge-type

high-density turbidity flows (Lowe, 1982; Nemec et al.,

1999; Plink-Bjorklund and Ronnert, 1999). The

intercalated thin- to thick-bedded sand, silt and

mud alternation with normal grading, planar

parallel-stratification, climbing-ripple cross-lamination

and planar parallel lamination are interpreted to result

from surge-type low-density turbidity flows (Ta–d).

Large-scale trough cross-bedding indicates the downslope

migration of dunes driven by quasi-steady low-density

turbidity flows (Nemec et al., 1999; Lowe, 1982; Mulder

and Alexander, 2001). Usually found in the distal upper

fan and both feeder and distributary channels.

FA3 Alternations of thin- to medium-bedded massive,

normally, inversely graded or planar parallel-

stratified pebbly sand (10–25%), thin- to

medium-bedded large-scale trough cross-bedded

medium- to coarse-grained sand (0–15%), very

thin- to thick-bedded massive, normally graded,

planar parallel-stratified and ripple cross-laminated

fine- to coarse-grained sand (20–80%) and

massive or normally graded gravel with erosive bed

contacts (10%). Subordinate thin beds of massive or

horizontally laminated silt and mud occur (1–3%).

Wedge (2.5–6 m thick,

up to 200 m long).

High- to low-angle

bedding (3–19j).Lenticular in channels.

The graded-stratified gravel and pebbly sand are

interpreted to have been deposited from surge-type

high-density turbidity flows (Lowe, 1982; Nemec et al.,

1999; Plink-Bjorklund and Ronnert, 1999). The

intercalated thin- to thick-bedded sand, silt

and mud alternation with grading, planar parallel

stratification, climbing-ripple cross-lamination and

planar parallel lamination are interpreted to result


Usually found in the proximal mid fan and both


FA4 Medium- to very thick-bedded fine- to coarse-

grained sand with scattered pebbles. Beds are

massive (10–50%), or show planar parallel

stratification (30%), ripple-trough cross-lamination

(10–25%), climbing-ripple cross-lamination

(0–40%) or large-scale trough cross-bedding

(5%). Very thin-bedded silt and mud beds

occasionally can be observed at the top of

climbing-ripple cross-laminated beds (1–2%).

These beds often show a fining-upward where


up to 80 m long).

Low-angle bedding

(3–9j). Lenticularin channels.

The thick fining-upward beds with planar parallel

stratification, large-scale trough cross-stratification

and ripple cross-lamination are interpreted to have

been deposited from pulsating quasi-steady low-density

turbidity flows (Mulder and Alexander, 2001). The

scattered pebbles are interpreted as coeval debris

fall from the steep upper fan slope (cf. Nemec et al.,

1999). Intercalated thin- to medium-bedded sand, silt

and mud alternation with planar parallel

stratification, climbing-ripple cross-lamination and


Facies association Geometry Interpretation of depositional processes

a lamination with eroded ripple stoss sides

commonly passes upwards into lamination with

preserved stoss sides and into draping

lamination. Bed contacts are usually sharp.

planar parallel lamination is interpreted to result


Usually found in the distal mid fan and both


FA5 Alternations of thin- to medium-bedded

fine-grained sand, silt and mud. The

fine-grained sand beds are massive (25%),

planar parallel-laminated (50%)

or show climbing-ripple cross-lamination

(15%). Silt and mud beds are massive

or show planar parallel-lamination (10%).

Bed contacts are usually sharp.

Dropstones can frequently be observed

and are often concentrated in mud layers.

Blanket, 0.5–2 m

thick. Lenticular

in channels.

The thin- to medium bedded sand, silt and mud

alternation with plan parallel stratification,

climbing-ripple cross-lamination and plan parallel

lamination is interpreted to result from surge-type

low-density turbidity flows (Ta–d). The frequent

occurrence of dropstones indicates dumping of ice-rafted

debris from icebergs (Lønne, 1995). Found in the outer

fan and feeder channels.

Table 2 (continued)


Mesozoic bedrock (Rausch, 1977; Rohde, 1994;

Deters, 1999). Frequent occurrences of ice-rafted de-

bris and compressive deformational structures point to

an ice-contact fan setting (Deters, 1999; Winsemann et

al., 2003). A description and interpretation of facies

associations is given in Table 2. The terminology for

gravel characteristics is after Walker (1975) and the

notation of turbidites, Tabcd, refers to the Bouma

divisions (cf. Bouma, 1962).

Fig. 10. Facies associations of the Coppe

4.1. Facies associations

4.1.1. Upper-fan facies associations (FA1 and FA2)

The proximal upper-fan deposits consist of 3–7 m

thick steeply dipping (16–34j) medium- to thick-

bedded clast-supported gravel (Fig. 10). The matrix is

sandy and bed contacts are mainly sharp. Larger clasts

can be oriented parallel to dip and a steeply imbricate

clast fabric with the a-axes dipping in upslope direction

nbrugge subaqueous fan complex.

Fig. 11. Facies associations of the Coppenbrugge subaqueous fan complex. (A) Steeply dipping gravel of the proximal upper fan (FA1),

unconformably overlain by outer- to mid fan-facies (FA5, FA4), Heerburg open pit. (B) Channelised thick- to medium-bedded sand and gravel

of the distal upper fan (FA2), Heerburg open pit. (C) Ripple cross-laminated sand of the distal mid fan (FA4), Steinbrink open pit. (D) Thin- to

medium-bedded sand, silt and mud of the outer-fan facies association (FA5), channel fill, Heerburg open pit. (E) Alternation of normally graded

gravel, massive and planar parallel-stratified pebbly sand, proximal mid-fan (FA3), Steinbrink open pit. (F) Climbing-ripple cross-lamination

and large-scale trough cross-bedding in sandy distal mid-fan deposits (FA4), Heerburg open pit.


is occasionally observed. Massive beds and sharp bed

contacts indicate deposition from noncohesive debris

flows (Nemec and Steel, 1984; Nemec, 1990; Shanmu-

gam, 2000). The scarcity of normal grading at the bed

tops suggests relatively slow, low-mobility debris

flows, with negligible shear and no turbulent churning


at the upper interface (Hampton, 1972; Nemec et al.,

1999). Intercalations of thin- to medium-bedded in-

versely or normally graded gravel, pebbly sand or

planar parallel-stratified sand amount to 5–30% and

increase towards the distal upper fan zone (FA2; Figs.

10, 11A and B). These deposits are interpreted to result

from surge-type high- and low-density turbidity flows

(Lowe, 1982; Mulder and Alexander, 2001). Occasion-

ally, large-scale trough cross-bedding can be observed,

indicating the downslope migration of dunes driven by

quasi-steady low-density turbidity flows (Cheel and

Rust, 1982; Nemec et al., 1999). Local chaotic bedding

indicates slumps caused by fan slope collapse probably

due to the alternating periods of fan-head aggradation

and high meltwater discharges (Postma, 1984, 1990;

Nemec, 1990; Ashley, 1995; Lønne, 1997; Plink-Bjor-

klund and Ronnert, 1999). Shallow chutes, filled with

massive or normally graded gravel or pebbly sand, can

frequently be observed and are 3–5 m wide and up to

0.3m deep, probably reflecting cut and fill processes on

the fan surface. Typical for the upper-fan environment

is the occurrence of large U-shaped feeder channels, up

to 25 m wide and 6 m deep. These feeder channels are

filled with coarse-grained gravelly deposits (FA1, FA2)

as well as finer-grained sandy and muddy deposits

(FA3–FA5, Figs. 10 and 11D).

4.1.2. Mid-fan facies associations (FA3 and FA4)

The proximal mid fan deposits (Figs. 10 and 11E)

consist of 2–6 m thick alternations of subhorizontal to

moderately dipping (3–19j) thin- to medium-bedded

massive, normally, inversely graded or planar parallel-

stratified pebbly sand, thin- to medium-bedded large-

scale trough cross-bedded sand and very thin- to

thick-bedded massive, normally graded, planar paral-

lel-stratified and ripple cross-laminated sand. Interca-

lations of massive or normally graded gravel with

erosive bed contacts can occasionally be observed.

Subordinate thin beds of massive or horizontally

laminated silt and mud occur. The proximal mid fan

slope is characterised by distributary channels, up to

20 m wide and 2.5 m deep, filled with graded or

stratified gravel, sand and silt. The graded-stratified

gravel and pebbly sand beds are interpreted to have

been deposited from surge-type high-density flows

(Lowe, 1982; Mulder and Alexander, 2001), whereas

the intercalated thin- to thick-bedded sand, silt and

mud alternation with grading, planar parallel stratifi-

cation, climbing-ripple cross-lamination and planar

parallel lamination are interpreted to result from

surge-type low-density turbidity flows (Ta–d). Fur-

ther downslope, the deposits grade into low-angle

dipping (3j–9j) fine- to coarse-grained sand beds

with scattered pebble, showing planar parallel strati-

fication and/or ripple cross-lamination. Typical are

thick beds with climbing-ripple cross-lamination often

showing a fining-upward where a lamination with

eroded ripple stoss sides commonly passes upwards

into lamination with preserved stoss sides and into

draping lamination (Figs. 10, 11C and F). The thick

beds with planar parallel stratification, trough cross-

stratification and ripple cross-lamination are inter-

preted to have been deposited from quasi-steady

low-density turbidity flows (Nemec et al., 1999;

Mulder and Alexander, 2001). High-suspension fall-

out rates are indicated by climbing-ripple cross-lam-

ination and graded or massive sand and silt beds. The

ripple cross-lamination shows evidence of a fluctuat-

ing and periodically waning flow, indicated by an

increase in the vertical aggradation rate of migrating

current ripples. The draping lamination indicates rip-

ples whose migration nearly ceases and vertical ac-

cretion prevailed (Ashley, 1995). The multiple

sequences of fining-upward beds reflect pulsating

quasi-steady underflows (cf. Mulder and Alexander,

2001), although the meltwater outflow was probably

semicontinuous. The pulsatory sediment discharges

may therefore be attributed to an autocyclic process of

fan-head aggradation and erosion or upper fan-slope

collapses (cf. Nemec et al., 1999). The scattered

pebbles are interpreted as coeval debris fall from the

steep upper fan slope.

4.1.3. Outer-fan facies association (FA5)

The outer-fan deposits are up to 4 m thick and

consist of thin- to medium-bedded fine-grained sand,

silt and mud (Figs. 10 and 11D), reflecting the decrease

of flow power with distance from subglacial tunnel

outlets. The fine-grained sand beds are massive (25%),

planar parallel-laminated (50%) or show climbing-

ripple cross-lamination (15%). Silt, mud and clay beds

are massive or show planar parallel lamination (10%).

Bed contacts are usually sharp. Dropstones can fre-

quently be observed and are often concentrated in mud

or clay layers. The thin- to medium-bedded sand, silt

and mud alternations are interpreted to result from


surge-type low-density turbidity flows (Ta–d). Mas-

sive beds and climbing-ripple cross-lamination indicate

high suspension fallout rates. The frequent occurrence

of dropstones indicates dumping of ice-rafted debris

from icebergs. Outer-fan deposits are mainly recorded

from channels or form blankets on more proximal fan

deposits.

4.2. The open-pit Otto section

The basal part of the Coppenbrugge fan complex is

exposed in the open-pit Otto (Fig. 9). Large parts of the

open pit have already been refilled and only a few

outcrop walls are still available for observation. The

fan deposits are exposed at an altitude of c. 84–100 m

and overlie glaciolacustrine mud and Mesozoic base-

ment rocks (Naumann, 1922; Deters, 1999). The

sedimentary succession is strongly deformed and part-

ly overlain by a basal till. The deformation is charac-

terised by large-scale folds and thrusts, steeply dipping

into easterly directions. The exposed facies includes

upper-fan coarse-grained gravels and mid-fan sands,

steeply dipping (16–30j) into southerly to southwest-

erly directions and interpreted to represent individual

foresets of small prograding fan lobes. Exposed fore-

sets are about 2–3 m thick and consist of thin- to

medium-bedded (5–25 cm) massive, matrix- or clast-

supported gravel. The sandy mid-fan deposits consist

of thin- to medium-bedded coarse- to medium-grained

sand or pebbly sand. Pebbly sand beds are massive or

show large-scale trough cross-bedding, whereas sandy

beds show planar-ripple cross-lamination or climbing-

ripple cross-lamination. Bed contacts are usually

sharp. Intercalated channels (chutes) are up to 1 m

deep and several meters wide and filled with normally

graded gravel and/or normally graded pebbly sand.

4.3. The open-pit HBT section

The sedimentary succession of the open-pit HBT,

exposed at an altitude of c. 147–165 m, overlies a

basal diamicton (Deters, 1999). As the open-pit Otto,

the HBT open pit has already been partly refilled and

only a few outcrop walls are available for observation.

The basal sedimentary succession consists of coarse-

grained channelised gravelly foresets, dipping (14–

20j) into northeasterly and southeasterly directions.

Individual foreset beds are 10–40 cm thick and consist

of massive clast-supported gravels with a coarse-

grained sandy matrix. Subordinate inversely and nor-

mally graded gravel beds occur. The coarse-grained

gravelly foresets are unconformably overlain by sandy

low-angle foresets, dipping (3–11j) into northwester-

ly directions. Beds are 5–70 cm thick and consist of

sand and pebbly sand, which are massive or show

climbing-ripple cross-lamination. Climbing-ripple

sequences often show fining-upward beds with eroded

stoss sides at the base and preserved stoss sides at the

upper parts of the beds. The major erosional surface is

only partly exposed. In one outcrop wall, c. 1 m above

this unconformity (c. 152-m altitude), wave-rippled

sand could be observed, indicating very shallow water

depths.

4.4. The open-pit Heerburg section

The sedimentary succession of the open-pit Heer-

burg is exposed at an altitude of c. 143–165 m and

overlies a basal diamicton, interpreted to represent a

basal till (Deters, 1999). The basal succession consists

of coarse-grained foresets, bounded by an erosional

upper surface. The foreset beds are 10–40 cm thick

and mainly consist of massive clast-supported gravel

with a coarse sandy matrix. The foreset beds are

steeply dipping (20–30j) into northwesterly and

southwesterly directions. The truncated gravelly fore-

sets are overlain by sandy mid-fan deposits (Figs. 11

and 12) whose deposition commenced with the infill-

ing of large channels incised into the underlying fan

body. These palaeochannels are up to 6 m deep and 25

m wide and exposed at an altitude of c. 148–154 m.

The basal channel-fill consists of thin- to medium-

bedded sand, which are massive, normally graded

and/or show planar parallel stratification and ripple

cross-lamination. Bed contacts are usually sharp.

Alternations of very thin-bedded clay, mud and silt

with thin- to medium-bedded sand occur upward in

the channel-fill succession. The fine-grained sand

beds are massive or normally graded. The silt, mud

and clay layers are massive and show sharp undula-

tory contacts. Dropstones are common and scattered

or accumulated in mud or clay layers. The sandy

channel-fill deposits are overlain by alternating thin-

to medium-bedded sand and gravel. The sand beds are

massive or normally graded and often pass upwards

into planar parallel-stratified and/or ripple cross-lam-

Fig. 12. Correlation panel of measured logs showing the facies stratigraphy in the Heerburg open-pit section.


inated sand. Gravel beds are matrix- or clast-sup-

ported and most often normally graded. Bed contacts

are sharp.

The erosional unconformity in the interchannel

areas is overlain by alternating beds of sand, silt

mud and clay, passing upwards into thick sand beds

with climbing-ripple cross-lamination, commonly

with an upward increasing stoss-side preservation. In

these sandy interchannel deposits, a localized erosion-

al structure, up to 1.5 m deep and 0.8–1.5 m wide, has

been recognized and attributed to iceberg ploughing

(Winsemann et al., 2003). Repeated observations of

the open-pit exposure during excavation revealed that

the scour structure is laterally continuous, extending

approximately 20 m in an east–west direction. The

scour has a concave-upward lower surface that trun-

cates the underlying deposits. The upper bounding

surface is formed by a subhorizontal erosional surface.

The steep-sided scour occurs in deformed sedi-

ments, and on both sides, downbending of marginal

strata into the central trough can be observed (Fig.

13). In the frontal (eastern) part of the scour, a wedge-

shaped deformation zone could be observed (Fig.

13C), which is interpreted to represent a relic of the

frontal ridge (berm). A detailed description and dis-

cussion of the ice scour formation is given in Winse-

mann et al. (2003). This scour feature is overlain

unconformably by c. 15 m of mid-fan deposits,

Fig. 13. Photographs of the iceberg scour recognized in the Heerburg open pit. The iceberg scour is laterally continuous, extending approx. 20 m

in an east–west direction. Flow direction of the iceberg was from west to east. (A) Frontal zone of the iceberg scour. The scour occurs in

deformed mid-fan sediments. (B) Close-up view of (A), showing the central trough and deformed marginal sediments. The zone of the central

trough is characterised by a pronounced downbending of overlying strata. These downbent strata are cut by vertical dewatering structures. (C)

Close-up view of (A), showing a wedge-shaped deformation zone at the right-hand side of the central trough, interpreted to represent a relic of

the frontal ridge (berm) of the iceberg scour. (D) Rear zone of the iceberg scour, approx. 15 m to the west of (A). The scour infill consists of

downbent marginal strata and turbated sandy mid-fan deposits.


consisting in the lower part of thinly to thickly bedded

cross-stratified sand and ripple cross-laminated silt.

Extensional faulting was active above the central

trough during deposition of these ‘‘post-iceberg-scour

strata,’’ as indicated by growth faults.

Normally or inversely graded gravel lenses occur

in the upper part, alternating with beds of pebbly

sand and sand, which are massive or normally gra-

ded or show planar parallel stratification and ripple

cross-lamination. Bed contacts are mainly sharp,

planar or undulatory, and dropstone horizons are

common. Channels in this part of the section are

10 m wide and up to 1.5 m deep, filled with thin

beds of massive or normally graded sand and pebbly

sand that pass upwards into fine-grained sand and

silt with horizontal or undulatory lamination. Bed

contacts are sharp, often erosive, and dewatering

structures are common.

4.5. The open-pit Steinbrink section

Deposits of the Steinbrink open pit are exposed at

an altitude of c. 150–162.5 m (Fig. 14) and overlie

glaciolacustrine mud and a diamicton, interpreted to

represent a basal till (Deters, 1999). Well data show

that in the western part of the open pit, sand predom-

inates the sedimentary succession, whereas the eastern

part is dominated by coarse gravel facies. This bipar-

tite division corresponds with opposite flow directions

measured in sandy and gravelly foresets.

Fig. 14. Correlation panel of measured logs showing the facies stratigraphy in the Steinbrink open-pit section.

J.Winsem

annet

al./Sedimentary

Geology165(2004)223–251

243


The exposed part of the eastward dipping sandy

mid-fan foreset can be traced laterally for about 40 m

and is up to 7 m thick. Foreset beds are 10–50 cm

thick and consist of coarse- to medium-grained sand

and pebbly sand. Beds are often fining-upwards and

show planar parallel stratification, large-scale cross-

stratification or ripple-trough cross-lamination and

planar-ripple cross-lamination, passing upwards into

climbing-ripple cross-lamination. Within the foreset

beds, small-scale channels, 1–1.5 m deep and up to

10 m wide, occur, which are filled with normally

graded gravel, pebbly sand and sand. Dip angles of

the sandy foreset are between 5j and 28j.The up to 7-m-thick upper-fan gravelly foreset

beds are steeply dipping (14–33j) into westerly

directions and consist of clast- and matrix-supported

gravel. Beds are 10–30 cm thick and contacts are

sharp. Both foresets are truncated by a major uncon-

formity, exposed at an altitude of c. 151–156 m and

partly evolved as a large U-shaped channel, up to 5 m

deep and about 25 m wide (Fig. 14). The basal

channel-fill consists of c. 2-m thin- to medium-bedded

sand, silt and sandy mud with inverse grading or

planar parallel lamination. These fine-grained chan-

nel-fill sediments are overlain by massive matrix- and

clast-supported gravels and massive pebbly sand.

The channel-fill and the truncated sandy and grav-

elly foresets are overlain by a c. 5-m-thick gravelly

foreset, steeply dipping (30j) into westerly directions.

This foreset again is incised by a channel, up to 2.5 m

deep and 25 m wide, and filled with thin- to medium-

bedded sands, showing climbing ripple cross-lamina-

tion, planar ripple cross-lamination, trough cross rip-

ple-lamination and thin- to medium-bedded massive

or planar parallel-stratified coarse sand and pebbly

sand. Bed often shows fining-upward trends and bed

contacts are sharp, erosive or gradual.

To determine the stratigraphic relation between the

oppositely dipping sandy and gravelly foresets, sev-

Fig. 15. (A) Georadar line from the open-pit Steinbrink at 150-m altitude.

the sandy foresets exposed in the western part of the open pit (Fan A). This

as a channel, which has been laterally filled by west- and eastward dippin

higher signal attenuation (1) limits further penetration. This sudden termin

attenuating finer-grained silt and mud, which underlie the fan deposits.

approximately 20 m northwards of georadar line A. The radar line shows

dipping reflectors (3) in the east (Fan B), which are correlated with the sa

Both foresets are incised by a channel (4), which has been laterally filled

channel towards the north. Penetration depth is c. 7 m (130 ns).

eral georadar lines were measured in the open pit at

different levels (150- and 157-m altitude). These

georadar lines have been measured in February and

October 2002, respectively. Georadar line A (Fig.

15A) was measured at 150-m altitude in February

2002. The GPR device used was a GSSI SIR-10B

with different antennae. The best penetration was

obtained using an 80-MHz bistatic antenna. The

dipoles have been separated by 1.8 m, and a trace

distance of 0.3 m was applied. A horizontal stacking

of 64 traces was applied in the field, and thus no

further processing was needed for enhancing the data

quality. The weather conditions have been dry; how-

ever, it was the first day after a period of heavy

rainfall.

The radar line shows eastward dipping reflectors

(2), ending at a depth of 200 ns (Fig. 15A). These

reflectors are correlated with the sandy foreset, ex-

posed in the western part of the open pit (Fan A). The

sandy foreset is bounded by an upper erosional

surface (4), partly evolved as a channel, which has

been laterally filled by west- and eastward dipping

strata (5). At the base of the georadar line, a layer

with a higher signal attenuation (1) limits further

penetration. This sudden termination of penetration

depth at c. 210 ns (c. 7 m) is attributed to finer-

grained silt and mud or diamicton beds, which un-

derlie the fan deposits.

Georadar line B (Fig. 15B) wasmeasured in October

2002 at 157-m altitude, approximately 20 m north-

wards of georadar line A. The best penetration was

obtained using a 100-MHz, bistatic antenna in a con-

tinuous mode. The radar data have been bandpass-

filtered with a butterworth filter to increase the data

quality. Penetration depth was limited by a high

groundwater level and there is no information below

130 ns (c. 7 m).

The radar line shows eastward-dipping reflectors

(1) in the west (Fan A) and westward-dipping reflec-

The radar line shows eastward dipping reflectors (2), correlated with

foreset is bounded by an upper erosional surface (4), partly evolved

g strata (5, Fan B). At the base of the georadar line, a layer with a

ation of penetration depth at c. 210 ns (c. 7 m) is attributed to higher

(B) Georadar line from the open-pit Steinbrink at 157-m altitude,

eastward dipping reflectors (2) in the west (Fan A) and westward

ndy and gravelly foresets exposed at the lower level of the open pit.

(5) by westward and eastward dipping strata. Note shallowing of


tors (3) in the east (Fan B), which are correlated with

the sandy and gravelly foresets exposed at the lower

level of the open pit. Both foresets are incised by a

channel (4), which has been laterally filled (5) by

westward- and eastward-dipping strata and is shallow-

ing towards the north.


4.6. Large-scale stackening pattern and depositional

sequences

The Coppenbrugge subaqueous fan complex con-

sists of laterally and vertically stacked small-scale

fan bodies, which are exposed in topographic heights

between c. 90–170 m. Palaeoflow directions are

highly variable and individual fan bodies are up to

10 m thick and several tens to hundreds of meters

across, lacking any subaerial, glaciofluvial or distrib-

utary delta-plain components. The small-scale fans

typically consist of coarse gravel in the proximal

core part, showing large-scale foreset bedding (cli-

Fig. 16. Stackening pattern and sequence stratigraphic interpr

nothems). The coarse gravel grades distally into

better-sorted, finer-grained sandy facies. Various

types of grading, cross-bedding and cross-lamination

are present, recording the downflow of both surge-

type and quasi-steady turbidity flows. Proximal to

distal fining reflects the drop-off of flow velocities

with distance from the tunnel mouth. Towards the

fan margin, high-suspension sedimentation rates pro-

duce climbing ripple-drift and graded or massive

sands and silts. Ball and pillow and flame structures

are common, owing to dewatering.

Within the subaqueous fan complex, three geneti-

cally distinct clastic units can be defined from facies-

etation of the Coppenbrugge subaqueous fan complex.


architectural analysis, which are separated by major

unconformities.

(1) A lower unit, which records proximal deposition on

a subaqueous fan during glacier advance. This unit,

exposed in the Otto open pit at altitudes between 84

and 100 m, is characterised by steeply dipping

coarse-grained foresets, dipping away from the

former ice margin. The fan deposits overlie

glaciolacustrine mud and are strongly deformed

by synsedimentary glacial tectonics. These deposits

mark the maximum ice-front position at the eastern

margin of glacial Lake Rinteln. Flow directions are

mainly from the northeast. The basal sequence

boundary is indicated by the abrupt onset of coarse-

grained fan deposition on fine-grained glaciolacus-

trine mud. The upper sequence boundary is

indicated by a basal till, which unconformably

overlies the fan deposits (Fig. 16).

(2) A medial unit, which records proximal deposition

on a subaqueous fan during glacier stillstand, after

the glacier and its injection points had retreated

towards a more upslope position. This medial unit

overlies a basal till and is exposed in the open-pits

HBT, Heerburg and Steinbrink at altitudes be-

tween c. 143 and 154 m. This medial unit is

characterised by steep coarse-grained foresets,

dipping away from the former ice margin and

showing only minor synsedimentary glaciotecton-

ic deformation. Flow directions are mainly

recorded from the west and southwest, and the

steeply dipping foresets are bounded by an upper

unconformity (Fig. 16), which is associated with

the formation of deep U-shaped channels, inter-

preted to represent a type-1 sequence boundary,

which resulted from a major lake-level fall.

Evidence for a lake-level fall is given by the

occurrence of wave ripple-cross-lamination, ob-

served in the open-pit HBT on top of the

unconformity at an altitude of c. 152 m. The

basal sequence boundary is indicated by the basal

till.

(3) An upper unit, which unconformably overlies the

coarse-grained foresets of the medial lower unit.

These fine-grained deposits consist of outer- to

mid-fan sediments, dipping towards the former ice

margin and thus representing an ice margin retreat

(Lønne, 1995, 2001). As the glacier retreated, the

deposition gradually switched from the fan’s

frontal slope to its ice-proximal backslope. Flow

directions, obtained from the climbing-ripple

sequences, show clear transport direction towards

the west, climbing up the older foreset slope

(open-pit Heerburg section). The rapid retreat of

the glacier resulted in calving and an abrupt cutoff

of the sediment flux to the fan. Evidence of

iceberg grounding and reworking is given by a

prominent ice scour mark, cut into mid-fan

deposits of the Heerburg section. The upper unit

shows an overall upward coarsening, indicating

the progradation of the new fan system.

5. Discussion

The Emme delta and Coppenbrugge subaqueous

fan are formed by stacked transgressive–regressive

sequences, indicating lake-level fluctuations. The

Emme delta complex, formed at the northern margin

of glacial Lake Rinteln, is thought to reflect a rela-

tively stable position of the ice margin in front of the

Weser Chains mountain ridges. The onset of delta

progradation probably represents a halt in ice advance

and a related high sediment supply period (cf. Lønne,

1995, 2001; Plink-Bjorklund and Ronnert, 1999). The

sedimentary facies and depositional architecture re-

semble those of nonglacial Gilbert-type deltas, except

for the deposition of glacial debris. Gravelly foreset

beds suggest a steep slope with gravity-driven flows.

Material that bypassed the braid plain avalanched

downslope as debris flow and stopped by freezing

when the slope diminished. The finer-grained sandy

material moved to the basin floor where it was

deposited from surge-type turbidity flows. The delta

stacking patterns indicate a long-term transgression,

interrupted by minor short-term lake-level falls and

closely resemble the Rf type (an overall rise with

superimposed falls) of Postma (1995). A similar Rf

lake-level control on delta architecture and facies has

been described by Fernandez et al. (1993). The main

delta progradation is recorded from the highstand

systems tract, when high meltwater and sediment

discharge occurred during the melt season. A forced

regression during opening of outlets led to the forma-

tion of subaerially exposed sequence boundaries and

the erosion of the highstand systems tract. The deeply


incised channels were filled with lake and glacioflu-

vial deposits during the subsequent transgression.

In contrast, the Coppenbrugge subaqueous fan was

associated with a retreating ice margin, indicated by the

shift of depocentres to more upslope positions. The

maximum ice advance is recorded from the open-pit

Otto, where the subaqueous fan deposits directly over-

lie glaciolacustrine mud. The sediments are strongly

deformed, thus indicating an advancing ice margin

(Lønne, 1995; Lønne and Syvitski, 1997). As the

glacier advanced into the lake basin, a conformable

coarsening-upwards sequence from lake sediments to

subaqueous outwash occurred, culminating in an ero-

sional sequence boundary and the deposition of a basal

till on top of the fan sediments. A subsequent glacier

retreat is recorded from the abrupt backstepping of fan

bodies, exposed in the open-pits HBT, Heerburg and

Steinbring at topographic heights between 143 and 170

m. These fan deposits overlie the basal till and form the

stratigraphic upper part of the Coppenbrugge fan com-

plex. The deposits only show minor synsedimentary

deformation structures and therefore probably mark

stillstand positions of the ice margin (cf. Lønne, 1995).

A prominent unconformity, exposed in topographic

heights between c. 145 and 156 m, is observable in

the open-pit outcrops HBT, Heerburg and Steinbrink.

This unconformity is associated with a major change in

palaeoflow directions. Below the unconformity,

palaeoflow directions are mainly from westerly and

southwesterly directions, whereas above the unconfor-

mity, palaeoflow directions are mainly from the east.

The formation of this unconformity is attributed to a

major lake-level fall, which probably led to a partial

subaerial exposure of the fan surface. Evidence for a

major drop in lake level is given by the occurrence of

wave-ripple cross-lamination on top of the unconfor-

mity, found in the HBT open pit at an altitude of c. 152

m. Therefore, it seems possible that this prominent

unconformity corresponds to the incised valley recog-

nized in the Emme delta at altitudes of c. 140–150 m

(Fig. 6). A subsequent lake-level rise is deduced from

the deposition of outer- to mid-fan deposits above the

unconformity, observable in the open-pits Heerburg

and HBT. In this section, an iceberg scour is recog-

nized, which is comparable in width and depth to

analogous features reported from modern and Pleisto-

cene ice scours and would require a water depth of at

least c. 10–20m above the fan surface (cf.Woodworth-

Lynas, 1996; Eden and Eyles, 2001). The trend of the

scour runs roughly west–east and a flow direction from

west to east is inferred from the formation of a frontal

ridge (berm). This flow direction is opposite to themain

flow directions measured in the upper sequence in the

open-pit outcrop Heerburg.

The deposition of the Coppenbrugge subaqueous

fan does not correlate with the maximum lake-level

highstand of glacial Lake Rinteln. Most probably, a

coeval ice advance from the western basin margin

successively closed lake outlets and led to the ob-

served overall lake-level rise. Possible lake outlets at

the southwestern margin of glacial Lake Rinteln lie at

altitudes between 49 and 207 m and increase in

altitude towards the east (Thome, 2001) and correlate

remarkably well with the topographic heights of

foreset–topset transitions of the Emme delta.

The ice front at the eastern lake margin, affected by

this lake-level rise, was thus probably very unstable,

subject to periodic calving, short-term oscillations and

possibly surges. These conditions can be expected to

have generated floating icebergs and caused switching

of meltwater outlets and variation in sediment dis-

charge (Boulton, 1986, 1990; Powell, 1990; Lønne,

1995; Plink-Bjorklund and Ronnert, 1999).

6. Conclusion

The forcing parameters governing the development

of depositional sequences were lake-level fluctuations,

sediment yield rates and physiography. Compared to

the changing lake level, which was on the order of 120

m, subsidence of the basin was insignificant.

Depositional processes reflect the character of sed-

iment supply and distance from the ice margin. The

sedimentary facies and depositional architecture of ice-

contact deltas resemble those of nonglacial Gilbert-

type deltas, except for the deposition of glacial debris.

Gravelly foreset beds suggest a steep slope with grav-

ity-driven flows. Material that bypassed the braid plain

avalanched downslope as debris flow and stopped by

freezing when the slope diminished. The finer-grained

sandy material moved to the basin floor where it was

deposited from surge-type turbidity flows. The main

delta progradation is recorded from the highstand

systems tract, when high meltwater and sediment

discharge occurred during the melt season. A forced


regression during opening of outlets led to the forma-

tion of subaerially exposed sequence boundaries and

the erosion of the highstand systems tract. The deeply

incised channels were filled with lake and glaciofluvial

deposits during the subsequent transgression.

A subaqueous fan complex formed where detritus

were carried to the lake via tunnels near or at the base

of an ice cliff. The Coppenbrugge fan complex was

constructed by the coalescence of several fan bodies,

which are recognizable as backstepping, southwest-

trending parallel morphologic sediment ridges. The

fan-stacking pattern and bulk geometry were deter-

mined by the ice-front fluctuations, the shifting of

meltwater outlets and the short-term lake-level fluc-

tuations. Individual fan bodies are up to 10 m thick

and several tens to hundreds of meters across, lacking

any subaerial, glaciofluvial or distributary delta-plain

components. Upper-fan deposits are characterised by

coarse-grained ice-proximal gravel, deposited by de-

bris flows that document continuous discharge. Slump

and slide deposits are related to steep depositional

slopes. The mid-fan contains basinward fining and

thinning deposits of quasi-steady and surge-type high-

and low-density turbidity currents, indicating more

ice-distal and periodic deposition. The outer-fan

deposits mainly consist of surge-type low-density

turbidites and glacial debris dumped by icebergs.

The formation of a prominent unconformity, associ-

ated with deeply incised channels, is attributed to a

major lake-level fall.

Acknowledgements

We would like to thank H.-B. Deters, P. Groetzner,

K. Skupin, P. Rohde, E. Speetzen, W. Thiem, K.

Thome, C. Vandre and P. Victor for discussions. A.

van Loon and two anonymous reviewers are thanked

for their critical comments, which helped to improve

the manuscript.

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Sequence analysis of early Saalian glacial lake deposits (NW Germany): evidence of local ice margin...

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