www.elsevier.com/locate/sedgeo

Sedimentary Geology 165 (2004) 67–92

Sedimentological processes in a scarp-controlled rocky shoreline to

upper continental slope environment, as revealed by unusual

sedimentary features in the Neogene Coquimbo Formation,

north-central Chile

J.P. Le Rouxa,*, Carolina Gomeza, Juliane Fennerb, Heather Middletonc

aDepartamento de Geologıa, Facultad de Ciencias Fısicas y Matematicas, Universidad de Chile, Casilla 13518, Correo 21, Santiago, ChilebBundesanstalt fur Geowissenschaften und Rohstoffe, Stilleweg 2, Hanover D-30655, Germany

cTime Frames Group, CSIRO Petroleum, Riverside Corporate Park, Delhi Road, North Ryde, NSW 2113, Australia

Received 19 November 2002; received in revised form 24 October 2003; accepted 14 November 2003

Abstract

Exceptionally good outcrops of Miocene to Pliocene deposits in the vicinity of submarine Paleozoic basement scarps at

Carrizalillo, north of La Serena, reveal a wealth of sedimentary features not commonly observed. The most proximal facies

consist of rock fall and coarse-grained debris flow deposits directly abutting the basement wall from which they originated.

Angular basement clasts are mixed with well-rounded cobbles, which probably formed as a basal gravel on a wave-cut platform

at the beginning of marine flooding, subsequently accumulated at the scarp edge and were incorporated into the debris when the

latter collapsed. The poor sorting, inverse grading, and protruding cobbles and boulders are classical debris flow features, with

good clast imbrication indicating a laminar shearing action.

A medial facies is represented by secondary channels running parallel to the major scarp about 1 km downslope of the first

locality. In the largest channel, megaflutes at the base indicate the passage of highly turbulent, nondepositing flows eroding the

soft, silty substrate. In the deepest, central part of the channel, a pebbly coquina shows horizontal and trough cross-stratification,

with most of the bivalves oriented convex side up. Meter-scale rip-up clasts of the underlying siltstone are also present, indicating

turbulent flow with a density sufficiently high to retard settling. The coquina is interpreted as a detachment deposit resulting from

a hydroplaning debris flow along the central part of the channel, where the velocity and rate of pore pressure decay were highest.

This deposit is overlain by fining upward, massive to horizontally stratified sandstone very similar in texture and composition to

the matrix of the debris flow, suggesting its formation by surface transformation and elutriation of the latter. Along the channel

margin, a basal centimeter-scale sandstone layer is virtually unaffected by the megaflute topography and clearly represents a

subsequent event. It is interpreted as a basal shear carpet driven by the overlying debris flow. Within the shear carpet, a basal

friction zone and an overlying collision zone containing a higher concentration of shell hash can be distinguished. The overlying

debris flow deposit is represented by massive coquina with scattered, angular to rounded basement clasts. It contains

disarticulated bivalves oriented with their concave side up, indicating large-scale upward fluid escape during deposition.

A smaller secondary channel shows large rip-up rafts of the underlying substrate. Some rafts appear to have been plucked

from the substrate by a process of sand injection from an overriding high-density sandy debris flow, which probably originated

during a tsunami. Such clasts can climb upward into a laminar flow by down-current tilting and tumbling.

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

doi:10.1016/j.sedgeo.2003.11.006

* Corresponding author.

E-mail address: jroux@cec.uchile.cl (J.P. Le Roux).

J.P. Le Roux et al. / Sedimentary Geology 165 (2004) 67–9268

The most distal facies occurs below a second scarp oriented more or less parallel to the present coastline, where finer-grained

turbidites onlap and backlap onto the stoss and lee sides of an obstacle formed by eroded boulder conglomerates. The onlap

deposits resemble inclined sandy macroforms recently described in submarine canyon settings. They are interbedded with

diatom-containing, volcanic ash beds with cross-stratification dipping eastwards and containing deepwater microflora typical of

continental upwelling zones.

D 2004 Elsevier B.V. All rights reserved.

Keywords: Coastal sedimentation; Submarine canyon; Continental shelf; Continental slope; Debris flow; Turbidite

1. Introduction 100-m topographic and bathymetric contours and

One of the least studied and most complex sedi-

mentary environments is that of rocky shorelines

(Felton, 2002). This environment is bounded by the

Fig. 1. Map indicating location of study area (small square north of La

described in text. The direction of transport as derived from channel trend

(n= 44).

occurs preferentially in tectonically active settings

(e.g., on the narrow shelves of active continental

margins). Typical geomorphic features are terraces

and cliffs transected by canyons, where erosional

Serena on inset of South America), as well as places and outcrops

s, cross-lamination, and flute casts is indicated in the rose diagram

J.P. Le Roux et al. / Sedimentary Geology 165 (2004) 67–92 69

and depositional agents such as waves, ocean currents,

gravity flows, and tsunamis are dominant (Felton,

2002). In this paper, we describe Miocene–Pliocene

deposits of the Coquimbo Formation in a remote

section of the Chilean coast some 80 km north of

La Serena (Fig. 1), which are attributed partly to this

environment. Exceptionally good exposures along

Quebrada Chanaral to the north of the hamlet Carri-

zalillo exhibit a wealth of sedimentary features that we

interpret in the light of various transport and deposi-

tional processes. This provides some new insights into

the sedimentology of rocky shoreline and associated

environments.

2. Geologic and tectonic setting

The study area lies within a region characterised by

four topographic domains, from west to east, com-

prising the Coastal Plain, Coastal Range, Central

Depression, and Andean Range. Basement rocks, onto

which the Coquimbo Formation was deposited direct-

ly, are represented by the Paleozoic Coastal Metamor-

phites, a metasedimentary succession underlying most

of the Coastal Plain.

The region falls within the Chilean flat-slab zone

between about 26jS and 33jS (Pardo et al., 2002).

Le Roux et al. (in press) studied the tectonic

history of the area during the Miocene–Pliocene

using paleobathymetry, backstripping, and dating of

the stratigraphic units based on Sr isotopes as well

as microfossils and macrofossils. They concluded

that the area underwent strong subsidence during

the late Burdigalian–early Langhian, which caused

marine transgression to continental shelf depths

over the coastal basement platform. The subsequent

approach and subduction of the Juan Fernandez

Ridge between about 13 and 8 Ma produced

tectonic uplift, but not sufficient to elevate the shelf

above the contemporaneous sea level. Renewed

submergence of the shelf between 8 and 3 Ma in

the wake of the southeastward-migrating Juan Fer-

nandez Ridge was followed by rapid uplift, with

the shelf finally emerging above sea level during

the Quaternary.

Water depths on the shelf after the initial flooding

fluctuated between about 20 and 200 m, as indicated

by the sedimentological facies (including bored phos-

phatic hardground) and foraminifera such as Bulimina

elongata, Buliminella elegantissima, Globigerina bul-

loides, Globigerinella calida, and Globigerinita glu-

tinata. On the seaward side of a basement scarp

running parallel to the present coastline, upper conti-

nental slope sedimentation is reflected by foraminifera

such as Brizalina aenariensis, sinistrally coiled Neo-

globoquadrina pachyderma, Globorotalia inflata, and

Orbulina universa, as well as a diatom assemblage

typical of marine upwelling on the continental slope

or outer shelf, including the Thalassiosira spp. Tha-

lassiosira eccentrica, Thalassiosira oestrupii, and

Thalassiosira leptopus; Stephanopyxis spp.; as well

as Azpeitia curvatulus, Actinocyclus curvatulus, Acti-

noptychus senarius, Actinoptychus splendens, Paralia

sulcata, Thalassionema nitzschiodes, and Nitzschia

fossilis (Le Roux et al., in press). Chaetoceros resting

spores, one of the most evident signs of upwelling, are

also common.

The best exposures are encountered along three dry

river valleys or ‘‘quebradas’’ that traverse the coastal

plain in a westerly to southwesterly direction. Along

the northernmost and largest of these, Quebrada

Chanaral (Fig. 1), a variety of Miocene–Pliocene

continental shelf and associated environments can be

identified. In its upper reaches, where the quebrada

approaches the rising Coastal Range, beach conglom-

erates with wave-polished pebbles partly covered with

barnacles are associated with poorly sorted fluvial

conglomerates showing upstream-dipping imbrica-

tion. Downstream along the quebrada, these rocks

pass into typical lower shoreface deposits with inter-

bedded, bioturbated (Thalassinoides) sandstones and

shales. Rare hummocky and swaley cross-lamination

attest to deposition by storms below the fair weather

wave base.

3. Basement topography

In the study area north of Carrizalillo, the Coastal

Plain is 10–14 km wide and slopes at about 1j from

sea level to the foot of the Coastal Range, which rises

from 200 m to elevations exceeding 1000 m in places.

Le Roux et al. (in press) suggested that the Coastal

Plain mirrors the palaeotopography during the Mio-

cene, when a coastal platform was carved into the

Coastal Metamorphites. A somewhat irregular plat-

J.P. Le Roux et al. / Sedimentary Geology 165 (2004) 67–9270

form topography is shown by the presence of local

basement hillocks rising to elevations of about 200 m

above present sea level, with the surrounding plain

lying between 125 and 175 m over most of the study

area (Fig. 2). That the basement platform was at sea

level during the early Langhian is suggested by the

fact that it is overlain by large, wave-smoothed

basement boulders partly covered with Balanus sp.

on the southeastern side of Quebrada Chanaral just

southwest of Chanaral de Aceituna (Fig. 1). A unit

directly overlying the boulder bed was dated by87Sr/86Sr at 14.6 Ma (Le Roux et al., in press). Ostrea

sp., which like Balanus sp. are particular to the littoral

(Stenzel, 1944) or sublittoral zone, are also found

clinging to the rocky surface of small gullies draining

into Quebrada Chanaral.

Basement rocks crop out mostly to the area

northwest of Quebrada Chanaral, which may be

due to a major fault displacing them downward to

the southeast. Evidence indicating such a fault is

seen, for example, at locality 1 (Fig. 1), where the

basement forms a vertical to slightly overhanging

cliff striking 204j, against which the Coquimbo beds

abut sharply from the southeast. There is a 2-m-wide

sheared zone in the basement along its contact with

the Coquimbo Formation, but towards the northeast

and base of the cliff, this zone is separated from the

Coquimbo Formation by a few meters of unbrecci-

Fig. 2. Diagrammatic representation of the basement topography

ated metamorphites. As the Coquimbo Formation

itself shows no sign of shearing, it probably post-

dates the fault. This is supported by an outcrop in a

small secondary quebrada northeast of Chanaral de

Aceituna, where the Coquimbo Formation similarly

abuts against a vertical basement wall, but also

overlies a horizontal ledge within the latter without

any sign of displacement.

This SSW-striking, probably fault-controlled

scarp formed by the basement appears to have had

a major influence on sedimentation patterns during

the Miocene–Pliocene, as suggested by the deposi-

tional facies encountered in its vicinity and de-

scribed in this paper. Downstream along Quebrada

Chanaral, however, the scarp gradually changes into

a southeastward-sloping surface (Fig. 2) onlapped

towards the northwest by younger stratigraphic units

of the Coquimbo Formation. It possibly formed by

erosion of the scarp prior to the Miocene–Pliocene

sedimentation.

About 1 km from the present coastline on both

sides of Quebrada Chanaral, another 320j trending

basement scarp dipping about 50jSW (Fig. 2) sep-

arates mainly shelf deposits on its landward side

from upper continental slope deposits on its seaward

side. This scarp possibly represents an ancient shore-

line cliff that was converted into the Miocene con-

tinental shelf break after the early Langhian

(not to scale) showing the location of the submarine scarps.

J.P. Le Roux et al. / Sedimentary Geology 165 (2004) 67–92 71

transgression, with the Coquimbo beds dipping at

about 3j seaward below it and draping landward

over the platform. At the platform edge/shelf break,

the basement slopes gently towards Quebrada Cha-

naral from both sides (Fig. 2), indicating that a

valley already existed here before deposition of the

Coquimbo Formation.

The SSW-trending scarp and valley described

above have about the same strike as the present

Quebrada Chanaral, possibly indicating topographic

inheritance (Le Roux, 1994). This is confirmed by

the trends of secondary channels as well as sedimen-

tary structures such as flutes and trough cross-beds

in the vicinity of these features. A rose diagram of

44 measured directional features (Fig. 1) shows a

mean transport direction of 194j, as compared to the

218j mean trend of Quebrada Chanaral. The ba-

thymetry of the present continental shelf and slope

opposite Carrizalillo also shows that the valley con-

tinues offshore to a depth of about 4000 m, termi-

nating in what seems to be a submarine fan on the

continental slope (Bundesanstalt fur Geowissenschaf-

ten und Rohstoffe, 2001).

Water depths at the foot of the SSW-trending scarp

could have reached 50–500 m, as suggested by the

presence of sinistrally coiled N. pachyderma within

the deposits. This planktonic foraminifer is a deep-

water species that lives in the photic zone during its

juvenile stage but, on maturing, descends to depths

normally exceeding 50 m, preferably below the ther-

mocline (Be and Tolderlund, 1971; Sautter and Thu-

nell, 1991). West of the NW-trending basement scarp,

water depths were on the order of 200–500 m, as

indicated by the foraminifera and diatom species

mentioned in Section 2.

4. Facies description

Three key outcrops in the vicinity of the submarine

scarps are described in this paper. The first is located

about 1 km northeast of the Chanaral de Aceituna

homestead (Fig. 1), where the SSW-trending scarp is

beautifully exposed. South and southwest of the

homestead, the basement dips gently southeastward,

where it is overlain by deposits including several

channelised features. The largest of these forms the

second locality described here (Fig. 1). The third

exposure is located about 800 m from the mouth of

Quebrada Chanaral below the NW-trending scarp,

where inclined macroforms are developed around

obstacles formed by gravelly deposits.

4.1. Facies association A: shelly sandstone interbed-

ded with poorly sorted conglomerate

At locality 1, the Coquimbo Formation dips 28jtowards the southeast against the basement scarp, but

flattens out rapidly away from the cliff (Fig. 3). A

total thickness of 35 m of deposits was measured (Fig.

4), representing the minimum height from the top of

the scarp to the basement floor (which is not exposed).

The platform edge here lies at an elevation of 150 m

above present sea level.

The age of the exposed succession is uncertain, but

it lies below unit S4 at locality 2 (Fig. 4), which field

relationships indicate is older than another unit dated

by 87S/86Sr at 5.6 Ma (Le Roux et al., in press). A

maximum age of 11 Ma is given by the presence of

the foraminifer N. pachyderma (which made its first

appearance in the Tortonian) at the base of the

succession.

4.1.1. Facies A1: shelly sandstone (biocalcarenite)

with lenses of coquina

More than 6 m of shelly sandstones (S1 in Fig.

4) are exposed below the first conglomerate (C1),

whereas the thickest conglomerates (C3/C4) are

overlain by about 8 m of similar deposits (S2).

These consist of brownish yellow, generally fine-

grained, poorly consolidated sandstone with frag-

mented to whole shells dominated by oysters

(Ostrea sp.) and pecten (Chlamys sp.). The frag-

ments are oriented parallel to the horizontal lami-

nation, except around fallen basement blocks where

they assume an orientation parallel to the sides of

the latter. Lenticular beds of coquina with larger

fragments and occasionally whole shells of oysters,

pecten, and barnacles (Balanus sp.) occur within

the shelly sandstone. Although the larger shell

fragments are parallel to the bedding, they show

no preferred convex-up or convex-down orientation.

The coquina lenses generally display irregular,

sharp basal contacts and also contain angular base-

ment clasts mostly less than 5 cm but up to 40 cm

in diameter. Towards the top of the measured

Fig. 3. View, looking NW, of scarp wall (basement) at locality 1, with biocalcarenites and coquina (S1–S3; light-coloured deposits at base of

cliff), as well as rock fall and debris flow conglomerates (C1–C5; dark deposits containing boulders and clasts). Human figure (encircled) for

scale (see also Fig. 4).

J.P. Le Roux et al. / Sedimentary Geology 165 (2004) 67–9272

succession, there are more bivalves and gastropods,

including Turritella sp. The foraminifera Textularia

gramen, G. bulloides, and the sinistral form of N.

pachyderma were identified within these deposits.

4.1.2. Facies A2: poorly sorted conglomerate

The lowermost conglomerate (C1) forms lenses up

to 30 cm thick and consists of very poorly sorted,

closely packed, angular to subangular basement clasts

that grade laterally into openwork debris. Isolated,

angular boulders or cobble clusters are present along

the same horizon, protruding into the overlying co-

quina. Towards the south, obliquely away from the

basement wall, this deposit cuts down into the under-

lying units at a low angle. Where the basement clasts

disappear in the same direction, the horizon continues

as a 5-cm-thick massive sandstone with only very

small shell fragments.

The second conglomerate (C2) is separated from

C1 by 1 m of coquina with a yellow sand matrix,

similar to the coquina lenses described above. This

unit is also thin and lenticular, with small, densely

packed basement pebbles filling 20-cm-deep, 30- to

80-cm-wide pockets in the underlying coquina. It is

Fig. 4. Measured stratigraphic columns at localities 1 and 2. On the left are the deposits of facies association A, with facies association B

deposits on the right. Dotted line shows probable correlation of units. Vertical scale applies to both columns.

J.P. Le Roux et al. / Sedimentary Geology 165 (2004) 67–92 73

overlain by 20 cm of fine shelly sandstones with

scattered, larger fragments of oysters.

The first major conglomerate (C3) has a sharp,

somewhat irregular basal contact. It exceeds 7 m in

thickness near the basement cliff but pinches out

about 100 m further south, oblique to the west–

southwestward transport direction indicated by the

clast imbrication. In the northernmost section, close

to the basement wall, this unit can be divided into

three subunits (Fig. 4). At the base, against the

cliff, is a boulder conglomerate thinning rapidly to

a 160-cm-thick zone of mostly small-pebble debris

with a fine sand matrix, but with large cobbles and

boulders up to 3 m in diameter protruding into the

middle sandstone subunit. Its upper contact is

mixed gradational. The clasts are mostly angular,

but they are intermixed with well-rounded basement

cobbles. Both clast types are oriented with their

long axes dipping up to 12j towards the basement

wall and varying in orientation between 240j and

270j. The middle subunit is a normally graded,

pebbly biocalcarenite, clasts decreasing upward

from a few centimeters in diameter to less than 1

cm. The upper subunit overlies an irregularly

scoured surface and contains cobbles and boulders

(up to several meters in diameter), which are

generally angular but intermixed with well-rounded

cobbles. The clasts are imbricated and indicate

transport between 250j and 270j. The upper part

of this subunit is normally graded, with outsized

clasts protruding into the overlying, brownish yel-

low sandstone.

About 15 m south of the thickest part of the

wedge, the lower and upper subunits merge. From

J.P. Le Roux et al. / Sedimentary Geology 165 (2004) 67–9274

this point, the conglomerate to all outward appear-

ance represents a single, uninterrupted bed gradual-

ly coarsening upward from pebble debris to cobble

and boulder debris. The basal part of the conglom-

erate in this section shows steep-sided, slightly

undercut scours filled with closely packed cobbles

and boulders (Fig. 5). The chutes are up to 80 cm

deep and 110 cm wide, with their sides parallel to

the orientation of clast long axes. The latter are

oriented between 228j and 258j, with maximum

dips of 25j towards the canyon wall. Shell frag-

ments directly below the scours are oriented parallel

to the contact.

The second major conglomerate (C4) varies in

thickness from 80 to 180 cm, showing a sharp,

erosional basal contact that incises into C3 towards

the south, where the contact between the two units

is formed by a shell-rich sandstone with horizontal

lamination. In the northern section close to the

basement cliff, the basal part of this unit consists

of small, densely packed pebbles with scattered

outsized clasts, showing inverse grading changing

upward into normal grading, where the unit has a

diffuse, gradational contact with the overlying lime-

rich sandstone. The clasts generally increase in size

towards the south, where one boulder reaches a

diameter of more than 4 m. Their orientation varies

between 242j and 295j.

Fig. 5. Boulder indentation chute, possibly undercut slightly by turbulent c

1). Note that the large boulder on the right did not produce loading, which

chute) for scale.

Unit C5 is a lenticular, poorly sorted conglom-

erate with a well-cemented, shelly matrix. It is

clast-supported. In contrast, unit C6 has an open

framework of small pebbles with some scattered

large, angular clasts, together with mostly whole

bivalve shells and barnacle fragments.

4.2. Facies association B: channelised coquina and

sedimentary breccia

This facies association is characterised by the

fact that it occurs in wide, relatively shallow

channels eroded into fine-grained sediments. These

channels trend parallel to the SSW-trending base-

ment scarp, the largest being located in Quebrada

Algarrobo southeast of Chanaral de Aceituna (Fig.

1). The base of the channel, where exposed along

the northeastern margin of the quebrada, lies about

125 m above sea level. This channel has a maxi-

mum depth exceeding 14 m (the base not being

exposed in its central, deepest part) and its width is

at least 200–300 m, but the eastern margin is

covered by surface sediments. The channel was

incised into ochre-coloured, silty sandstone, which

is horizontally laminated with thin, light green

argillaceous bands showing ripple lamination and

horizontal bioturbation. Lenses of brown-weathered

coquina with Turritella and other gastropods, as

urrent caused by water displacement in front of debris flow (locality

argues against such an origin for the chutes. Geological hammer (in

J.P. Le Roux et al. / Sedimentary Geology 165 (2004) 67–92 75

well as barnacle fragments and basement clasts, are

developed within this fine-grained unit. These

lenses show erosional basal contacts, low-angle

cross-stratification, and fining-upward trends. Hum-

mocky cross-lamination is present in sandstone unit

S4 (Fig. 4) about 12 m above the main channel

deposits. The foraminifers G. bulloides and Globi-

gerinoides ruber were identified in overlying silty

sandstones. An age of about 8–5.6 Ma is suggested

by field relationships and correlation with nearby

dated units.

The western margin of this channel apparently

trends 175j across Quebrada Algarrobo, but is not

very well exposed. The base here is sharp, being

characterised by megaflutes up to 60 cm wide, 15

cm deep, and more than 80 cm long. Their trends vary

between 179j and 208j with a mean of 193j (26

measurements), approximating the 204j trend of the

basement wall described in Section 4.

A second, younger channel is progressively incised

into the basal channel, especially on the southern side

of Quebrada Algarrobo (Fig. 6). This unit has a shell-

rich conglomerate at its base, containing whole

bivalves, barnacles, Turritella, and other gastropods.

Higher up, it grades into open framework conglom-

erate composed mostly of very poorly sorted base-

ment clasts, overlain with a gradational contact by

coquina with medium-sized, broken bivalves, Turri-

tella, and barnacles. Towards the east, the unit dis-

plays large-scale trough cross-bedding, with low-

angle cross-bedding also present at the top.

4.2.1. Facies B1: basal, thin, pebbly sandstone

This facies is represented by a very poorly

sorted, pebbly sandstone layer with a consistent

thickness of 1–2 cm (‘‘shear carpet’’ in Fig. 7)

occurring at the base of unit C6. The mean grain

size is between 0.25 and 0.35 mm, but scattered

pebbles up to 2 cm in diameter are present within

and especially towards the top of the bed, where

they protrude into the overlying pebbly coquina.

The sandstone continues uninterruptedly through the

erosive flutes at the base of the channel with no or

only a minor change in thickness (Fig. 8), but

contains fewer clasts and shell fragments along

the steep sides of the scours.

The basal part of the sandstone shows no clear size

grading and is composed of densely packed quartz

grains with bioclasts including shell and echinoid

spine fragments. The upper few millimeters, however,

display a higher concentration of slightly coarser shell

hash (in places showing an imbricated structure) and

some outsized clasts, giving the layer an overall

coarsening-upward appearance. The bed has a sharp,

undulating upper contact.

4.2.2. Facies B2: stratified pebbly coquina

Facies B2 (C6; Fig. 4) occurs at the base of the

succession in the central, deepest part of the main

channel. It is a well-stratified, pebbly coquina con-

sisting of bivalves, Turritella, and other gastropods,

together with scattered basement clasts in a coarse-

grained, very poorly sorted sandstone matrix. Bed

thickness varies between 2 and 15 cm. Shallow

troughs reaching about 15 m in width merge laterally.

These merge laterally across the channel with crude

horizontal stratification formed by conglomerate

lenses 40–70 cm thick, in which the clasts reach a

concentration of about 70%. The bivalves are medi-

um-sized, disarticulated, and largely broken, the ma-

jority being oriented with their convex sides up, which

is the most stable orientation for current flow. The

Turritella shells also have a preferred orientation

parallel to the current, with their apices pointing

upstream. Large rip-up clasts of soft, silty sandstone,

exceeding 1 m in diameter, occur near the base. This

unit grades upward into the unstratified, pebbly co-

quina of facies B3.

4.2.3. Facies B3: massive, pebbly coquina

These deposits represent most of the channel fill

along its western margin, where they overlie facies

B1, but towards the channel axis, they overlie facies

B2 with a gradational contact. Although the overall

lithology is similar to that of B2, there is a general

absence of stratification or sedimentary structures and

the basement clasts are, on average, somewhat larger

(with some outsized clasts reaching 50 cm in diame-

ter). The conglomerate has a closely packed frame-

work and the clasts are generally horizontally oriented

near the base, but there is no sign of imbrication. The

bivalves are conspicuously larger, disarticulated but

generally unbroken, and most are oriented with their

concave sides up (Fig. 7). Turritella shells lie mostly

horizontal, but lack a directional orientation. Sam-

pling for microfossils indicates that this unit is appar-

Fig. 6. Panoramic view (looking SW) of secondary channels south of Quebrada Algarrobo. Note erosional nature of younger channel (C7) and

elutriation sandstones overlying basal channel (C6). Unit C6 has a maximum thickness of about 3 m on photograph.

J.P. Le Roux et al. / Sedimentary Geology 165 (2004) 67–9276

ently sterile, compared with other stratigraphic units

where abundant foraminifers were encountered.

4.2.4. Facies B4: massive, fining-upward

biocalcarenite

Overlying facies B3 (unit C6) in the central part

of the channel with a generally sharp contact is a

buff, medium to very coarse sandstone with scat-

tered, millimeter-scale shell fragments, very similar

to the matrix of the underlying coquina. The unit is

massive, but grades upward into fine-grained, argil-

laceous sandstone. It is eroded by the overlying

pebbly coquina, so that a maximum thickness of

about 350 cm is preserved on the southwestern side

of Quebrada Algarrobo. Towards the east, this sand-

stone becomes more shell-rich, with many broken,

small fragments of barnacles. Here it is horizontally

stratified into more and less shell-rich beds 10–15

cm thick.

4.2.5. Facies B5: sedimentary breccia

A second, much smaller channel occurs a few

hundred meters south of the Chanaral de Aceituna

homestead at an elevation of about 100 m. The deep-

est part of the channel is filled by shelly conglomerate

containing many large, unoriented clasts of the sandy

Fig. 7. Shear carpet at base of debris flow deposit (thin white sandstone layer) at locality 2. Concave-upward bivalves in the debris flow

conglomerate indicate fluid escape.

J.P. Le Roux et al. / Sedimentary Geology 165 (2004) 67–92 77

and silty substrate. Just to the north of the channel

axis, a ruptured 1-m-thick bed of sandstone is appar-

ently enclosed in a massive, coarser-grained sandstone

bed. The latter becomes very thin at the base of the

block, disappearing in places, but is apparently linked

to its upper part by large cracks traversing the block

(Fig. 9).

Fig. 8. Megaflute in fallen block representing base of channel at locality 2

and the overlying debris flow (facies B3) can be clearly seen, the former be

shear carpet does not change significantly within the flute, indicating its d

4.3. Facies association C: fining-upward sandstones,

poorly sorted conglomerates, inclined sandy macro-

forms, and diatom-containing tuffs

The outcrop from locality 3 described here lies at

an elevation of 20 m above sea level on the southern

side of Quebrada Chanaral below the NW-trending

a. The textural difference between the basal shear carpet (facies B1)

ing much finer (smoother appearance on photo). The thickness of the

eposition during a subsequent, nonturbulent event.

Fig. 9. Sandstone bed traversed by cracks injected with coarser-grained sand from above, locality 2. Similar cracks apparently injected from the

base occur in the same bed. For explanation, see text and Fig. 13.

J.P. Le Roux et al. / Sedimentary Geology 165 (2004) 67–9278

scarp (Fig. 1), along the shallow trough formed by the

basement topography (Fig. 2). It shows a basal suc-

cession of sandstones and conglomerates dated at

7.3–5.2 Ma by Sr isotopes and the foraminifer B.

aenariensis (Le Roux et al., in press). These deposits

were truncated by a land- and seaward-dipping ero-

sional surface, forming a prominent ‘‘ridge’’ that was

subsequently onlapped by argillaceous, very fine-

grained sandstones as well as diatom-containing vol-

canic tuffs (Figs. 10 and 11). The onlap succession

was dated at between 4.5 and 3.4 Ma by the joint

presence of B. aenariensis and G. calida (Le Roux et

al., in press). A younger conglomerate directly over-

lies the basal succession as well as the upper onlap

units with an erosional contact. It has a probable age

of 3.4–2.6 Ma as indicated by the joint presence of

the pelecypod Chlamys hupeanus (Late Pliocene) and

the gastropods Chorus blainvillei (Late Miocene to

Late Pliocene) and Chorus giganteus (Late Pliocene

to middle Pleistocene). The basal and onlap sediments

are also overlain by shales containing the foraminifer

G. inflata, indicating a probable age of less than 2.6

Ma, in turn overlain by a coquina dated by 87Sr/86Sr at

2.0–1.8 Ma (Le Roux et al., in press).

4.3.1. Facies C1: massive, bioturbated biocalcarenite

The lowermost unit exposed below the onlap

surface is a massive, fine-grained, intensely biotur-

bated (Thalassinoides) sandstone with scattered shell

fragments, exceeding 4 m in thickness. Its uppermost

20 cm is rich in organic material, showing 1-m-wide

troughs and linguoid ripple cross-lamination as well

as small slumps. This unit is sharply overlain by

massive, bioturbated sandstone grading upward into

very poorly sorted, coarse-grained sandstone contain-

ing angular quartz, feldspar, and volcanic clasts ex-

ceeding 1 cm in diameter. The grain size varies

irregularly, but there is a generally coarsening-upward

trend up to about 90 cm from the base, from where

this second unit passes gradually into a meter-thick,

medium-grained sandstone with pockets of very poor-

ly sorted sandstone containing clasts up to 5 mm in

diameter. No sedimentary structures except bioturba-

tion are visible.

A prominent (at least 5 m deep) narrow chute cuts

through the upper sandstone into the lower sandy unit.

It has very steep edges (76j) and trends between 178jand 208j, being filled by a succession of 3- to 15-cm-

thick beds of gritty sandstone grading upwards into

fine, massive sandstone (Fig. 11). Contacts between

the units are sharp and bedding dips increase to about

48j against the chute edges.

4.3.2. Facies C2: coarse, pebbly coquina

The units described above are truncated and over-

lain by up to 1 m of medium to coarse, pebbly coquina

Fig. 10. Outcrop showing basal succession of sandy debris flow deposits overlain by conglomerate deposited by turbulent flow at locality 3

(looking E). Geological hammer (encircled) for scale. Outlined is a deep scour chute truncating the sandy debris flow deposits, which is filled by

several subsequent sandy turbidites. On the right, the erosional surface is onlapped from the southwest by argillaceous sandy turbidites. The

diatom-containing tuffs were deposited by upwelling currents.

J.P. Le Roux et al. / Sedimentary Geology 165 (2004) 67–92 79

with very poorly sorted, angular to well-rounded

cobbles and boulders of basement, intrusive rocks,

lava, phosphate, and yellowish sandstone. No imbri-

cation is observed, but the largest boulders (reaching

110 cm in diameter) tend to fill the deepest chutes.

The shelly matrix is dominated by barnacle and pecten

Fig. 11. Measured stratigraphic columns at locality 3. The lower left column

right shows details of the onlap units. Dotted lines show stratigraphic rela

fragments, and is very poorly sorted. Towards the

northeast, the basal contact of the conglomerate cuts

into the underlying units, where the unit presents

large, shallow troughs (trending 188j) lined with

scattered cobbles and filled by low-angle cross-bed-

ded coquina with clasts and boulders. The coquina

shows the succession below the erosional onlap surface. That on the

tionships. Vertical scale applies to both columns.

J.P. Le Roux et al. / Sedimentary Geology 165 (2004) 67–9280

grades upward into 35 cm of massive to horizontally

and low-angle cross-stratified, medium-grained bio-

calcarenite with fine-grained shell fragments oriented

parallel to the bedding.

4.3.3. Facies C3: inclined sandy macroforms

The erosional onlap surface strikes 293j with a dip

decreasing downward from 20j to 12jS. The onlap

succession consists of a series of 20- to 180-cm thick

sandstone beds, the thickness of each bed increasing

away from the ridge formed by the onlap surface (Fig.

10). All the sandstones have sharp bases in places

overlain directly by isolated shells, cobbles, pebbles,

or pebble lenses, but there is a scarcity of floating

larger clasts within them. The grain size of all beds

increases conspicuously towards the onlap surface

(e.g., from very fine-grained to coarse-grained over

a distance of 10–15 m). The sandstones consist of

very poorly sorted grains in a highly argillaceous

matrix that generally decreases in percentage upward,

with some exceptions where the opposite trend is

observed. A conspicuous coarse-tail normal grading

is present, however, with sorting improving notice-

ably towards the bed tops. Although most beds are

massive, some show horizontal lamination or small-

scale trough cross-lamination at the top, where bio-

turbation is also present. The burrows are filled by

coarser material from the base of the overlying unit.

On the northeastern, poorly exposed side of the

basal ridge, onlap relationships are also evident. The

southwestward-onlapping units are fine- to very fine-

grained sandstones with horizontal and ripple lamina-

tion, varying in thickness between 40 and 50 cm.

They also have sharp basal contacts.

4.3.4. Facies C4: diatom-containing tuffs

Interbedded with the sandy beds are a few white-

weathering, tuffaceous, argillaceous sandstone beds

containing diatoms, including Chaetoceros resting

spores, Stephanopyxis, Thalassiosira, and Gramma-

tophora species, as well as P. sulcata. These units

vary in thickness between and 180 and 280 cm, have

sharp basal contacts, and, in some cases, display

planar cross-lamination dipping at angles of 3–13jtowards the northeast (038j) (i.e., against the general

current flow). Horizontal and ripple lamination (also

oriented towards the northeast) and extensive biotur-

bation (Megagrapton) consisting of branched hori-

zontal tubes about 1 cm in diameter and up to 30 cm

in length are also present.

5. Facies interpretation

5.1. Facies association A: foot-of-scarp traction

current, rock fall, and debris flow deposits

Field relationships in this spectacular outcrop

clearly indicate that the poorly sorted, unstratified

conglomerates of this facies formed by sporadic

collapse of the basement wall onto the channel floor,

which consists of finer-grained shelly sandstones and

coquinas.

5.1.1. Facies A1: foot-of-scarp traction current

deposits

The fine-grained, shelly sandstones are similar to

the deposits filling modern submarine canyons. Scott

and Birdsall (1978), for example, described samples

collected at water depths between 14 and 400 m along

the Hueneme Canyon on the Californian coast. Sedi-

ments occurring along the canyon wall in this case

consist of thinly laminated, silty sands with pelecypod

and gastropod shells and few biogenic structures,

contrasting with the highly bioturbated central canyon

floor deposits.

The presence of mostly disarticulated, fragmented

shells of oysters (Ostrea sp.), pecten, barnacles, and

Turritella in facies A1 suggests that these sediments

were probably derived from shallower water and were

transported along the scarp by currents. The forami-

nifera also seem to represent a mixture of sedimentary

environments: T. gramen is a robust species found in

beach deposits (Boltovskoy, 1955), G. bulloides lives

preferentially in intermediate water depths between 50

and 100 m (Be and Tolderlund, 1971), and sinistrally

coiled N. pachyderma is essentially a deep-water

species (Be and Tolderlund, 1971; Sautter and Thu-

nell, 1991), although it may also occur in shallower

water. The first two species were therefore probably

derived from shallow coastal areas, whereas the last

may represent an actual water depth of between 50

and 500 m.

The orientation of shell fragments parallel to the

borders of basement blocks indicates sedimentation

from suspension, possibly representing material

J.P. Le Roux et al. / Sedimentary Geology 165 (2004) 67–92 81

churned up by blocks falling and sliding from the

scarp. The coquina lenses were probably deposited by

stronger currents capable of transporting coarser ma-

terial and eroding the sandy substrate.

5.1.2. Facies A2: rock fall and debris flow deposits

All six conglomerate units (C1–C6) are interpreted

as rock fall and debris flow deposits originating

directly from the basement cliff against which they

terminate. This is suggested by their poor sorting,

generally angular clasts, absence of stratification, and

the presence of outsized clasts along the upper contact

of the beds, as well as the imbrication dipping

obliquely towards the basement wall.

The thin sandstone layer continuing laterally

from unit C1 may represent material churned up

by and/or expelled from the debris flow and

deposited beyond the stalled head and margins of

the latter, thus representing surface transformation

and elutriation deposits (Hampton, 1972; Fisher,

1983).

The pebble pockets at the base of unit C2 have

been described as typical of slurried beds, interpreted

to be deposits of debris flow (Mutti et al., 1978;

Shanmugam, 2000). However, the small scale of these

lenses, as well as the tight packing of the pebbles, may

indicate an origin as debris falls rather than flows.

In unit C3, the juxtaposition of the conglomerate

and basement wall in the northern part of the outcrop

implies that sedimentation processes must have been

somewhat different from those normally assumed for

mass flows. Imbrication in both units dips obliquely

towards the basement cliff, indicating the latter as the

principal source of the angular clasts, whereas the

well-rounded cobbles could not have been shaped by

transport over a distance of only a few meters. They

must therefore have originated by falling from the

scarp edge, where they had been rounded by wave

action at the beginning of marine transgression over

the basement platform (Fig. 12). The fact that they are

intermingled with angular scarp debris indicates that

at least the upper part of the wall collapsed. Due to the

close proximity of the debris flow deposits to the wall,

this in turn implies that some of the clasts fell freely or

bounced to their present position. The middle sandy

subunit in C3 may thus represent material churned up

by these falling basement cobbles and boulders,

forming a suspension cloud settling out on top of

the debris, similar to that observed around isolated

basement cobbles and boulders in the traction current

foot-of-scarp deposits. This is supported by its fining-

upward trend.

The upper subunit of C3 overlies an irregular

contact scoured into the underlying sandy subunit,

which was apparently caused by turbulent flow. Such

flows may originate from water displaced away from

the foot of the cliff by the large volume of material

free-falling and sliding from above.

The apparently uninterrupted nature of unit C3

away from the basement wall suggests that caution

should be exercised when interpreting inverse grading

in such deposits. Obviously, some time elapsed be-

tween the deposition of the middle sandy subunit from

suspension and the second major collapse of the

canyon wall. It is therefore unlikely that the apparent-

ly merged, coarsening-upward basal and upper debris

flows represent one event at a greater distance from

the source.

The scoured and undercut basal contact of unit C3

at this locality obviously did not originate from

turbulent flow alone, as indicated by the contact-

parallel orientation of shell fragments in the silty

sandstones below the deeper chute forms. The latter

were probably formed by large blocks bouncing and

sliding ahead of the debris flows, indenting the soft

substrate and reorienting the shell fragments by their

weight. A load cast origin in this case seems unlikely

because large basement boulders lying just above the

contact are not associated with a deepening basal

contact directly below them, as would have been

expected in this case (Fig. 5). However, minor subse-

quent modification of the boulder indentation chutes

by turbulent flow is indicated by their slightly under-

cut walls, unlikely to have been caused by the

boulders. These flows may have originated from water

displaced by the falling debris and accelerating along

the indentation chutes.

The erosional basal contact of unit C4 suggests the

passage of an earlier nondepositing turbulent flow.

The inverse-to-normal grading indicates basal shear-

ing accompanied by surface flow transformation,

perhaps with the elutriation of fines by escaping fluids

(Fisher, 1983). An increase in clast size along the flow

path accompanied by the presence of some very large

boulders is also typical of debris flow fronts, where

debris fall blocks are common (Bagnold, 1968; Suwa,

Fig. 12. Wave-reworked boulders overlying basement platform, which formed during a marine transgression between 16 and 15 Ma. These are

‘‘broken rounds’’ [i.e., well-rounded clasts that have subsequently been broken by (a) very-high-energy event(s)]. Geological hammer (next to

boulder in centre of photograph) for scale. At locality 1, similar boulders were incorporated into the rock falls, mixing with the angular clasts of

the latter.

J.P. Le Roux et al. / Sedimentary Geology 165 (2004) 67–9282

1988; Sohn, 2000). These features occurring within a

relatively short distance from the debris source sug-

gest that the flow rapidly froze.

In spite of the proximity of these conglomerates

to their origin, it is interesting that they already

exhibit many of the characteristics of typical debris

flow deposits, such as protruding outsized clasts. As

it is difficult to visualise shear sorting taking effect at

a distance of only meters from the source, it is

possible that many such clasts owe their position at

the top of debris flows to the fact that they fell

directly from the collapsing wall onto the developing

flow, being unable to sieve through to the bottom as

would have been the case for smaller clasts. Such

outsized clasts may ride for long distances down

submarine canyons, perhaps in the process advancing

towards the snout of the flow by virtue of their

greater momentum.

The presence of imbrication so close to the

canyon wall also indicates that laminar shearing

already commenced at the collapse-and-slide stage,

in this case probably enhanced by the up to 28jangle of the beds against the wall. Imbrication in

debris flow deposits has been reported by various

authors (e.g., Nemec, 1990; Sohn, 2000), but it

may be that its development requires a high shear-

ing rate, as this structure was not noted in any of

the more distal debris flow deposits described in

Section 5.2.

5.2. Facies association B: detachment, shear carpet,

debris flow, elutriation, and rip-up raft deposits

The general setting of these channels is shown by

the nature of the underlying and overlying deposits.

The underlying silty sandstone with horizontal bio-

turbation obviously represents a low-energy environ-

ment, whereas the fining-upward lenses of coquina

indicate minor, waning currents transporting coarser

material and disturbing the tranquil environment from

time to time. Water depths were probably on the order

of 100–140 m, as indicated by the presence of

hummocky cross-lamination as well as the foramini-

fers G. bulloides and G. ruber, which according to

Phleger (1960) are typical of the inner continental

shelf. The presence of megaflutes at the base of the

channel indicates that it was eroded by a highly

turbulent flow, but the channel-fill facies B1–B5

apparently represent different stages of deposition.

5.2.1. Facies B1: shear carpet deposits

The fact that this sandstone layer was apparently

unaffected by the megaflute topography indicates that

its deposition postdated the formation of the flutes and

J.P. Le Roux et al. / Sedimentary Geology 165 (2004) 67–92 83

that its nature was not that of a typical turbulent flow,

which would have filled the scours more evenly as a

result of the sudden drop in current velocity in areas of

expanding flow depth.

The sandstone layer is considered to represent a

basal shear carpet. The term ‘‘traction carpet’’ has

been applied to such deposits by various authors (e.g.,

Dzulynski and Sanders, 1962; Sohn, 1997), but as

correctly pointed out by Shanmugam (2000), such

dense flow layers have a plastic rheology and laminar

flow state, so that the term ‘‘traction’’ in this case is

misleading. These deposits are commonly only a few

centimeters thick and normally develop beneath rap-

idly dissipating, turbulent sediment flows including

turbidity currents (Yagishita, 1994), pyroclastic flows

(Chough and Sohn, 1990), and subaerial hypercon-

centrated flows (Todd, 1989). Energy is provided to

basal shear carpets by the shearing action of the

overlying flow, so that grain-to-grain interactions

dominate. Although some authors have attributed

the inverse grading typical of basal shear carpets to

the dispersive pressure characterising such flows (e.g.,

Todd, 1989; Hiscott, 1994), this concept as originally

proposed by Bagnold (1954) has been challenged

recently. Sohn (1997) considered a variety of process-

es, including kinematic sieving (Middleton, 1970), a

geometrical mechanism whereby large grains climb

over smaller grains (Haff, 1991), as well as the

drifting of larger grains toward the zone of least

dispersive stress at the top of the collision zone, which

overlies a basal friction zone in basal shear carpets.

Legros (2002) argued that dispersive pressure in grain

flows or basal shear carpets cannot be invoked as a

mechanism for the upward segregation of large par-

ticles because an increase in dispersive pressure

would cause an immediate expansion of the flow until

it equals the applied normal stress again. Only par-

ticles lighter than the bulk density of the flow can

therefore be pushed upward. Le Roux (2003), how-

ever, challenged this concept, arguing that basal shear

carpets represent a stage of disequilibrium before

expansion can dissipate the dispersive pressure. He

proposed a combination of kinematic sieving in the

upper, expansional collision zone and kinematic

squeezing in the basal, compressive friction zone as

a mechanism for inverse size grading.

In the present case, the nonorganic fraction of the

basal shear carpet is not clearly size-graded, with the

exception of some outsized clasts (up to 2 cm in

diameter) floating within or occurring at the top of the

layer and protruding into the overlying debris flow.

However, the upper few millimeters do contain more

shell hash, which is slightly coarser (in length, but not

in volume) than the nonorganic grains. The absence of

large shell fragments so typical of the overlying debris

flow deposit also indicates that these were in some

manner eliminated from the basal shear carpet. To

explain this observation, the following mechanism is

envisaged: As the shear carpet developed, expansion

in the upper collision zone caused kinematic sieving

(i.e., the momentary creation of holes in the agitated

granular mass, through which the smaller particles

could fall more easily than the larger ones) (Middle-

ton, 1970; Savage and Lun, 1988). Due to the shear-

ing action, shell fragments would be oriented

horizontally, presenting their largest surface areas to

the holes and consequently being unable to penetrate

to the base of the flow. The concentration of small

grains in the developing, basal friction zone would

also lead to fluid expulsion, perhaps helping to

suspend the small shell fragments in the overlying

collision zone.

Sohn (1997) suggested that basal shear carpets

should increase progressively in thickness with con-

tinued sedimentation. If the base of the collision zone,

which is initially thicker than the friction zone (Fig. 3;

Sohn, 1997), shifts upward simultaneously with the

growth of the latter (Le Roux, 2003), this would lead

to an ungraded basal friction zone showing only a

slight change in grain size or in the percentage of shell

fragments at the top where it grades into the collision

zone. The relative thickness of the friction and colli-

sion zones may depend therefore on the duration of

the flow more than on the type of flow, with the

friction zone increasing in importance during more

sustained flows.

5.2.2. Facies B2: detachment deposits

The occurrence of traction structures and large

sandstone rafts within this facies indicates turbulent

flow capable of eroding the soft, but cohesive sandy

substrate and incorporating chunks of it into the

resulting deposits. However, the fact that these giant

rip-up clasts do not rest on the base itself suggests

hindered settling owing to a high concentration of

suspended particles.

J.P. Le Roux et al. / Sedimentary Geology 165 (2004) 67–9284

These deposits are interpreted as having formed by

detachment of debris from a hydroplaning debris flow

front, its disintegration and partial dilution by mixing

with the ambient seawater, followed by deposition

well ahead of the more slowly advancing main debris

flow. Similar processes have been described in lahars

by Pierson and Scott (1985), Smith (1986, 1987), and

Scott (1988), who referred to sediments of this origin

as hyperconcentrated flow deposits. They have char-

acteristics intermediate between debris flows and

streamflows, producing ungraded or normally graded

gravel and crudely stratified gravelly sand deposits

(Harrison and Fritz, 1982). However, the term is

currently ambiguous, having been applied to both

Newtonian fluids characterized by a turbulent state

in which coarse and fine particles settle together, or to

Bingham fluids characterized by a high sediment

concentration and laminar state, in which coarse and

fine particle are deposited together by freezing (Qian

et al., 1980). The genetic term ‘‘detachment deposits’’

is preferred here to avoid confusion.

The fact that this facies only occurs in the deepest

part of the channel and not along its margins may

reveal an important concept concerning the processes

involved. It is to be expected that a debris flow would

attain its highest velocity along the channel axis. One

of the conditions for hydroplaning is that the densio-

metric Froude number Frd be higher than 0.4 (Mohrig

et al., 1998). Frd is given by:

Frd ¼ V=M½ðqd=qf � 1Þghcosh� ð1Þ

where V is the velocity of the debris flow, qd is the

density of the debris flow, qf is the density of the

ambient fluid, g is the acceleration of gravity, h is the

debris flow thickness, and h is the slope angle. It

follows that Frd would be highest in the channel

centre, so that 0.4 may be exceeded here and not

along the slower-moving channel margins.

A second factor conditioning hydroplaning is the

time scale of pore pressure decay R (Iverson and

LaHusen, 1989; Mohrig et al., 1998; Sohn, 2000):

R ¼ d2l=kE ð2Þ

where d is a characteristic length identified with the

average flow depth, l is the viscosity of the interstitial

fluid, k is the permeability of the debris, and E is the

uniaxial compression modulus or stiffness of the

debris. R would thus be higher in the channel centre

due to the increased flow depth, enhancing the possi-

bility of hydroplaning. It is thus conceivable that the

fastest-moving central part of a channelised debris

flow may hydroplane and produce detachment depos-

its directly downslope, whereas the margins may be

nonhydroplaning and lack such underlying deposits.

5.2.3. Facies B3: cohesionless debris flow deposits

The massive, ungraded nature of this facies,

together with its poor sorting and sandy matrix, is

typical of cohesionless debris flow deposits domi-

nated by frictional grain interactions (Kim et al.,

1995; Sohn, 1997). The horizontal clast alignment

suggests laminar shear (Fisher, 1971), which is

supported by the presence of a basal shear carpet

normally associated with laminar, (pseudo)plastic

flow (Postma et al., 1988; Shanmugam, 2000). It

therefore probably behaved as a Bingham fluid. The

concave-up orientation of the bivalves indicates

fluid escape, which may be supported by the

absence of microfossils. Although foraminifera have

the density of calcite (2.71), their hollow interiors

would have diminished their settling velocity and

thus facilitated their expulsion. According to Shan-

mugam (2000), water escape is common in sub-

aqueous debris flows. In experimental studies, sandy

debris flows showed water entrapment beneath the

flow, followed by water escape to form dish struc-

tures, vertical pipes, and sand volcanoes (Mohrig et

al., 1998). Owing to the coarse nature of the debris

flow deposits described here, however, these types

of structures would be difficult to observe, if

formed at all.

5.2.4. Facies B4: elutriation deposits

The similarity between the sandstone and the

matrix of the underlying debris flow deposit suggests

that it may have been derived from the latter by

surface transformation and elutriation (Fisher, 1983).

Dilution and stripping of sand from the top of the

debris flow (Hampton, 1972) and expulsion of fines

from within the latter by escaping fluids would have

generated an upper turbulent cloud, in this case

lagging behind the faster-moving debris flow deposits.

Sohn et al. (1999) interpreted thinly stratified, coarse

to fine-grained sandstone overlying debris flow con-

glomerate as the deposits of hyperconcentrated flows

J.P. Le Roux et al. / Sedimentary Geology 165 (2004) 67–92 85

generated behind the debris flow head by surface flow

transformation. This is exhibited towards the western

channel margin in the present case. According to

Shanmugam (2000), sandy turbidites show a general

lack of cross-bedding and a characteristic normal

grading, which is displayed in the central part of the

channel. The difference in sedimentation style may be

due to larger volumes of sand having been expelled in

the central, faster-moving, and hydroplaning part of

the debris flow, leading to denser turbidity currents

and rapid dumping of the sediment load (as in the

basal unit of the Bouma cycle), whereas the channel

margin may have experienced less elutriation, forming

more dilute turbidity currents and depositing horizon-

tal beds (second unit of the Bouma cycle).

5.2.5. Facies B5: rip-up raft deposits

The fact that sandstone rafts occur throughout the

channel thickness of several meters may suggest that

they originated by undercutting and collapse of the

channel margin, but an alternative mechanism is

suggested by the cracks described above (Fig. 9).

The medium to coarse sandstone filling the cracks

appears to have been injected into the latter from both

above and below (not shown in Fig. 9), apparently

under high pressure. As the fractures are obviously

not mudcracks, a likely explanation is that they

formed during an earthquake, followed soon after by

a tsunami, causing a strong high-density backflow

accentuated along the scarp and secondary channel

axes. Experiments have shown that high-concentra-

tion beds may develop at the base of currents under-

going rapid deposition by sediment falling out from

suspension (Middleton, 1967; Vrolijk and Southard,

1997). They are driven by shear from the overriding

flow. The high dispersive pressures developing within

such pseudoplastic quick beds could cause injection

of sand into fractures, widening the latter and even

penetrating the entire bed and proceeding along its

basal contact, as seems to have been the case here.

Similar features have been reported in submarine

canyons by Morris and Busby-Spera (1988), as well

as in subaqueous volcaniclastic gravity flows (Bal-

lance and Gregory, 2001). Sand-injected cracks may

thus be a characteristic feature of high-pressure hyper-

concentrated flows and especially of tsunami events.

Fractured bed blocks can be lifted and separated from

the substrate in this manner, riding down the channels

on a cushion of high-pressure quicksand. Down-

current tilting of some fractured blocks would expose

their up-current margins to the upward-increasing,

laminar shearing action of the quick bed and also

assist in their being plucked from the substrate (Fig.

13). Camacho et al. (2002) show such a tilted margin

frozen in place (their Fig. 8A). Due to the upward

increase in the shear velocity of laminar flows, some

blocks thus separated from the substrate may be

transported down-current in a tumbling fashion, pro-

gressively climbing upward into the flow as assisted

by kinematic sieving. The high density of the quick

bed would prevent them from sinking.

5.3. Facies association C: continental slope deposits

An upper continental slope environment is indi-

cated for this facies (Figs. 10 and 11) by the

presence of sinistrally coiled N. pachyderma, which

is often encountered at depths greater than 200 m

on the upper continental slope (Lagoe, 1984) as

well as G. inflata and O. universa, foraminifers

common in hemipelagic sediments (Natland, 1976).

The presence of rare grass phytoliths may indicate

transport by turbidity currents from shallow coastal

areas.

5.3.1. Facies C1: sandy debris flows

The two units at the base of the measured section

(Fig. 11) are interpreted as sandy debris flows because

of their massive nature, floating clasts, and coarsen-

ing-upward trend in the case of the upper sandstone.

The sedimentary structures in the uppermost part of

the basal unit indicate current flow, possibly due to

surface flow transformation, whereas the small slumps

may be the result of shearing by the overlying sandy

debris flow. Pebbly pockets in the upper part of the

latter flow are also typical of slurried beds and debris

flows (Mutti et al., 1978). The presence of Thalassi-

noides in submarine canyons and deeper marine

environments is well documented (Buatois et al.,

2002).

The massive, fining-upward, chute-filling beds

were probably deposited from suspension, as sup-

ported by the fact that they drape upward against

the steep channel walls. They are interpreted as

turbidites deposited after the channel had been eroded

by a turbulent nondepositing flow.

Fig. 13. Plucking process by quicksand injection into cracks, widening the latter and separating the block from the substrate. Tilting of the block

exposes it to the laminar shearing action of the quick bed and overturns it in a down-current direction, in the process climbing into the flow.

J.P. Le Roux et al. / Sedimentary Geology 165 (2004) 67–9286

5.3.2. Facies C2: gravelly, turbulent flows

The conglomerate was apparently deposited by a

highly competent, turbulent flow, as indicated by the

large-scale traction structures. Komar (1970) pro-

posed that pebble and cobble-size material may be

transported by turbidity currents along confined

environments such as channels, which is thus sup-

ported by the present study. The fining-upward

nature of the conglomerate indicates waning flow

conditions.

5.3.3. Facies C3: inclined turbiditic macroforms

The fact that the dip direction (203j) of the

erosional onlap surface parallels the erosion chute

described above as well as the transport direction

given by megaflutes at the base of the channelised

debris flow deposits implies that it was eroded by a

nondepositing turbulent current flowing in the same

general direction.

The massive, fining-upward nature of the sand-

stones, together with their high clay content, is typical

of turbidites and indicates Newtonian rheology and

fluid turbulence (Pettijohn, 1957; Sullwold, 1960).

According to Shanmugam (2000), sands in true tur-

bidity currents are transported by suspended load and

thus show a general lack of cross-bedding and a

characteristic normal grading. The very last stages of

sand transport by the ‘‘dilute tail’’ end of turbidity

currents, however, may form thin divisions of parallel

lamination by reworking (Kuenen, 1953, 1964), as

also observed here. The fact that larger pebbles only

occur at the base of the beds suggests unhindered

settling, thus distinguishing these sandstones from

sandy debris flows. An unusual feature of these

turbidites is the fact that the clay matrix generally

decreases upward, which may be due to high sediment

concentrations dampening turbulence in the basal

parts of the flows and leading to rapid deposition,

while increased turbulence in the more diluted, upper

parts of the flows kept the finest grains in suspension.

Stanley et al. (1978) referred to similar beds as coarse,

truncated turbidites, which may form the largest

proportion of channelised sequences in the distal

sectors of submarine canyons. Middleton (1967) also

observed that high-concentration turbulent flows can

develop coarse-tail grading as a function of pseudo-

plastic behaviour.

The onlapping bedform architecture is similar to

structures recently described in a submarine canyon in

Spain by Pickering et al. (2001). This new type of

bedform was named ‘‘inclined backstepping macro-

forms’’ by these authors and also occur in the up-

stream depression immediately behind a wave-like

ridge on the surface of cohesive debris flow deposits.

The coarse-grained sandy turbidites composing the

macroforms similarly have significantly coarser up-

stream tails and show partial bioturbation towards the

bed tops. These macroforms are sigmoidal and in-

clined upstream, however, in contrast with the down-

stream-dipping macroforms described here. Pickering

J.P. Le Roux et al. / Sedimentary Geology 165 (2004) 67–92 87

et al. (2001) attribute the origin of these structures to

the deposition of high-velocity gravity flow conglom-

erates overlying erosion surfaces left by nondeposit-

ing turbidity currents. The conglomerates had a wave-

like upper surface, which acted as a seeding structure

for the accumulation of sand from successive turbidity

currents behind the crests. Lower shear velocities over

the depressions behind the crests caused the preferen-

tial deposition of sand behind these obstacles, with

flow stripping leaving pockets of coarser-grained

sediments on the upstream side of the macroforms.

The absence of modification features to the macro-

forms was taken to indicate rapid deposition and

backfilling.

A similar origin is proposed for the inclined macro-

forms described here, with the difference that the

inclined depositional surface at the base of the macro-

forms formed by nondepositing turbulent flows erod-

ing the underlying conglomerate. Subsequent sandy

turbidites were deposited not only on the downstream

side of these obstacles, however, but also on the

upstream side. The depressions on both sides were

filled in by high-density turbidity flows rapidly dump-

ing their loads due to the increase in water depth and

decrease in flow velocity, with successive turbidites

onlapping onto the crest from both sides. The ridge

was initially left bare due to an increase in flow

velocity (decrease in water depth) over its top, but

as the depressions filled up, their effect became less

until the crest itself was buried by successive flows.

Although Pickering et al. (2001) attributed the down-

stream-fining trends of their beds to flow stripping, a

more likely explanation may be simply that the

coarsest sediment load was dumped in the zone of

flow expansion immediately behind the ridge crests,

with successively finer grains settling out as flow

deceleration proceeded down the inclined surface.

The inclined macroforms would thus have grown in

a down-current direction away from the ridge crests

and not up-current from the depression onto the ridge

crest, although successive turbidites do show a back-

stepping architecture onto the latter.

5.3.4. Facies C4: upwelling deposits

The diatom assemblage of these beds is charac-

teristic of highly productive deep-water upwelling

over the continental slope (Le Roux et al., in press).

This is supported by the up-channel directed planar

cross-lamination, although the latter may also be

related to flood tidal currents rising along the sub-

marine canyon.

The upwelling deposits and the bioturbated bed

tops of the turbidites suggest that infilling of the

depression behind the ridge crest was sporadic, with

relatively long periods between successive turbidite

flows.

6. Facies relationships and sedimentological

history

The study area is characterized by numerous

lateral facies changes and units pinching out over

short distances. This is exemplified, for example, at

locality 1, where debris flow conglomerates pinch

out within 100 m from the basement cliff against

which they originated or grade into sandstone in the

direction of transport. Within secondary erosion

channels, there are also rapid facies changes per-

pendicular to the direction of transport, where

detachment deposits showing traction structures

within the channel centres grade into massive debris

flow deposits towards the channel margins. Flow

transformation processes in the down-channel direc-

tion also caused different vertical facies representing

the same event, for example, where basal shear

carpets underlie debris flow deposits or where

elutriation deposits overlie detachment deposits. In

addition, obstacles along the main canyon floor

were responsible for local facies variation by alter-

ing the hydrodynamics of turbidity currents, caus-

ing, for example, backstepping macroforms in their

wake.

Because of these abrupt facies changes and incom-

plete exposures along Quebrada Chanaral, com-

pounded by the irregular basement topography, it is

very difficult to correlate the different stratigraphic

units. At locality 1, for example, units were deposited

simultaneously on the basement platform and the

scarp floor, although the two environments were

separated vertically by at least 35 m. Here, rounded

basement boulders formed along the edge of the scarp

during marine transgression in the early Langhian (Le

Roux et al., in press) were subsequently incorporated

in Tortonian or even younger deposits at the foot of

the scarp when the wall collapsed, as indicated by the

J.P. Le Roux et al. / Sedimentary Geology 165 (2004) 67–9288

presence of N. pachyderma in the immediately under-

lying deposits.87Sr/86Sr dating of the units is complicated by

reworked macrofossils and microfossils, especially

where the latter were sampled within erosion chan-

nels. At locality 3, for example, Sr dates of 7.3,

5.4, and 5.2 Ma were obtained in the first coquina

bed (Fig. 11, left-hand column, 6 m from the base

of the succession). The 7.3-Ma age probably cor-

responds to a fossil reworked from the underlying

nonchannelised unit (right-hand side; Fig. 11),

which has a maximum age of 7.3 Ma given by

the presence of B. aenariensis.

The sedimentological history is thus complex, as

in probably all rocky shoreline-associated environ-

ments, but the basic scheme of events can be out-

lined as follows: During the Burdigalian to early

Langhian, an irregular coastal platform was eroded

subaerially onto Palaeozoic basement rocks. Marine

transgression shortly thereafter inundated this plat-

form, as attested by the presence of in situ Balanus

and Ostrea sp. associated with a wave-smoothed

boulder bed directly overlying the basement. This

transgression was associated with basin subsidence

caused by a deceleration of plate collision and

culminated in a relative sea level highstand at the

commencement of the Serravalian (14.3 Ma), when a

depth of about 170 m was reached on the continental

shelf (Le Roux et al., in press). However, marine

sedimentation probably took place to the east of the

SSW-striking fault-controlled scarp even before this

transgression, with the scarp possibly constituting a

shoreline cliff. The younger, deeper water deposits

presently exposed against the scarp at locality 1 were

possibly deposited during an eustatic marine trans-

gression at 11–10 Ma, when global sea levels rose

between 20 and 120 m above the present sea level

(Haq et al., 1988; Abreu et al., 1998). At this time,

the basement platform was at an elevation of about

50 m below present sea level (Le Roux et al., in

press), so that a depth of 100 m or more as possibly

indicated by the foraminifer N. pachyderma could

have been reached on the canyon floor. The scarp

had a long history of sporadic collapse lasting well

into the Tortonian (11–7.3 Ma), with coarse, prox-

imal rock falls, slides, and debris flows entering

obliquely into a submarine canyon or half-graben

running parallel to the scarp. The collapse of the

canyon wall may possibly be attributed to tectonic

activity associated with the approach of the Juan

Fernandez Ridge between about 13 and 8 Ma (Le

Roux et al., in press). At the peak of this tectonic

uplift event at around 8.5 Ma, relative marine re-

gression reduced water depths to a minimum of

about 20 m on the shelf. Renewed transgression

due to tectonic subsidence in the wake of the

migrating Juan Fernandez Ridge after 8.5 Ma caused

a return to water depths of about 170 m during the

Zanclean (around 4.9 Ma). It was possibly during

this time (late Tortonian to Messinian) that the

secondary channels at locality 2 were eroded and

filled along the trough-shaped, southwestward con-

tinuation of the half-graben. This coincides with the

age of the smaller chutes carved by nondepositing

turbulent flows at the base of the succession at

locality 3, west of the NW-striking scarp. The

possible association of these erosion channels and

their debris flow deposits with tsunamis suggests

continued tectonic activity. The overlying Zanclean

age succession initially shallows upward from outer

shelf to middle shoreface depths, but subsequent

flooding followed by a highstand period is reflected

in an extensively bored phosphatic hardground87Sr/86Sr dated at 3.9 Ma. This bed represents a

condensed section on the outer continental shelf and

is correlated with the turbidite sandy macroforms on

the upper continental slope at locality 3, where the

cooccurrence of B. aenariensis and G. calida indi-

cates an age range of 4.3–3.4 Ma (Le Roux et al., in

press). The fine-grained nature of these turbidites

contrasts with the underlying conglomerates and

coquinas, supporting a relatively greater distance

from the shore because of the creation of new

accommodation space further inland by rising sea

levels. A subsequent lowstand during the Piacenzian

(3.4–2.6 Ma) is represented by inner shelf deposits,

followed by a return to outer shelf deposition with

contemporaneous upper continental slope sedimenta-

tion (including mudrocks and diatom-containing

tuffs) below the ancient shelf break at locality 3.

The presence of G. inflata in these deposits gives a

maximum Piacenzian or more likely Gelasian ( < 2.6

Ma) age (Le Roux et al., in press). The succession is

capped by Gelasian upper shoreface deposits dated at

2.0–1.8 Ma by 87Sr/86Sr. Reemergence of the shelf

during the Pleistocene formed an extensive coastal

J.P. Le Roux et al. / Sedimentary Geology 165 (2004) 67–92 89

plain covered by fluvio-estuarine, shelly gravel, Sr

dated at 1.0 Ma.

7. Discussion and conclusions

Many of the features described in this paper are

similar to the deposits of rocky shorelines and

associated submarine canyons. In the present case,

a fault-controlled basement scarp apparently

deflected and concentrated seaward-directed debris

flows and turbidity currents, with the development

of secondary channels parallel to the scarp. Al-

though sedimentologically active submarine canyons

have been studied fairly extensively, for example,

off California (Shepard and Dill, 1966; Shepard and

Marshall, 1973; Scott and Birdsall, 1978) and the

northeastern United States (Keller and Shepard,

1978; Drake et al., 1978), they are rarely well

exposed in the rock record (Stanley et al., 1978).

The same can be said of rocky shorelines, which

have not been well studied even in modern exam-

ples because of the difficult and often dangerous

field conditions. The exposures described here rep-

resent a wide variety of shoreline, scarp-, and

canyon-related deposits and sedimentary features

that give further insights into depositional processes

not readily revealed by dives or observations from

submersibles, coring programs, and seismic surveys.

This confirms the observation by Kelling and Stan-

ley (1978) that in some environments, ‘‘the ancient

record still appears to be providing a key to the

present.’’

The existence of turbulent, extremely erosive cur-

rents roaring down submarine channels without leav-

ing any deposits has been postulated by various

authors (e.g., Mutti and Normark, 1987), which is

supported by the presence of wide, large-scale chan-

nels or narrow, deep chutes with deposits representing

subsequent events, such as at localities 2 and 3.

Although the existence of clear water currents has

been postulated by Bouma and Hollister (1973), it

seems more likely that highly turbulent, sediment-

loaded currents carved the erosion surfaces but simply

traveled too fast to dump their load. It is also likely

that these currents are related to tsunami events, which

are fairly common along the Chilean coast and may

have been amplified by the scarp and canyon topog-

raphy. It is to be expected that such tsunamis eroded

the coastline and would have been capable of trans-

porting a fairly coarse sediment load down to the foot

of the continental slope, where a permeable deep sea

fan probably exists.

Local currents may precede debris flows due to the

displacement of water away from the base of scarps or

canyon walls suffering catastrophic collapse, leading

to basal erosion surfaces as recorded at locality 1. The

irregular basal surfaces of some debris flows may also

be related to large fall blocks bouncing, rolling, and

sliding ahead of the flows and indenting the soft

bottom sediments. These indentation chutes may be

subsequently enhanced by local turbulent flows pro-

ceeding ahead of the debris flow by water displace-

ment. Large blocks may also fall onto developing

debris flows and ride down canyons, so that protrud-

ing outsized clasts do not necessarily imply shear

sorting.

Flow transformation by hydroplaning, detach-

ment, and elutriation may be restricted to certain

parts of debris flows, in particular the deepest parts

of channels where the highest velocities are attained

and where the time scale of pore pressure decay is

also the largest. Turbulent flow developing by de-

tachment and dilution of the debris flow front may

proceed ahead of the latter and deposit somewhat

finer material showing sedimentary structures such as

horizontal and trough cross-stratification. In shelly

debris flow deposits, fluid escape structures such as

pillars and dishes may not be visible, but bivalve

shells may be oriented concave-side up. Shear car-

pets apparently also occur at the base of debris

flows, where they show floating clasts and a larger

concentration of small shell fragments in their upper

part (collision zone). This can be attributed to their

orientation (by shearing) parallel to the base and

consequent inability to infiltrate the basal friction

zone.

Sand injected from high-density, pseudoplastic

quick beds into fractures penetrating underlying strata

probably indicates earthquakes followed by tsunamis

and may thus be a characteristic feature of such

events. Plucking of large rafts from the substrate in

this manner may tilt them down-current, exposing

them to the laminar shearing action of the flow. They

may thus proceed down-current in a tumbling fashion

and climb upward into the flow at the same time,

J.P. Le Roux et al. / Sedimentary Geology 165 (2004) 67–9290

assisted by kinematic sieving, dispersive pressure, and

upward-escaping fluids.

Inclined sandy macroforms, such as recently de-

scribed for the first time in Spain, are not necessarily

sigmoidal in shape and can also dip downstream

instead of upstream on the lee sides of obstacles on

the canyon floor. Their fining-downstream trends may

be attributed to flow expansion behind canyon floor

obstacles and do not necessarily imply flow stripping.

They are thus envisaged to have grown in a down-

stream direction, with successive beds backstepping

onto the ridge crests. On the upstream side of

obstacles, similar onlapping beds may be deposited.

Acknowledgements

We are greatly indebted to K.A.W. Crook, Y.K.

Sohn, and F. Ricci-Luchi for their very helpful

comments and suggestions. Petrysia Le Roux and

Carlos Vinegas are thanked for assisting with field-

work. We also appreciate the help of Daniel

Frassinetti and Margarita Marchant in identifying the

macrofossils and foraminifera, respectively. The

research was funded by Project Fondecyt 1010691.

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