Flat-pebble conglomerates: a local marker for Early ...hera.ugr.es/doi/14995542.pdf · Flat-pebble conglomerates: a local marker for Early Jurassic seismicity related to syn-rift

Flat-pebble conglomerates: a local marker for Early Jurassicseismicity related to syn-rift tectonics in the Sesimbra area

(Lusitanian Basin, Portugal)

J.C. Kullberga,*, F. OloÂrizb,1, B. Marquesa,2, P.S. Caetanoa,2, R.B. Rochaa,2

aCentro de InvestigacËaÄo em GeocieÃncias Aplicadas da Universidade Nova de Lisboa, Quinta da Torre, 2825-114 Caparica, PortugalbDepartamento de EstratigrafõÂa y PaleontologõÂa, Facultad de Ciencias, Universidad de Granada, Campus de Fuentenueva s.n., 18002

Granada, Spain

Received 22 February 2000; accepted 28 August 2000

Abstract

Flat-pebble conglomerates have been identi®ed in the Lower Toarcian (Levisoni Zone) carbonates of the Sesimbra region

(30 km south of Lisboa, Portugal) and related to submarine mass movements. Their origin is explained through a three-stage

model based on the comparative analysis of potential generating mechanisms taking into account timing and type of geody-

namic evolution in the Lusitanian Basin: (a) differential lithi®cation of thin carbonate and non-bioturbated horizons embedded

within a more argillaceous matrix; (b) disruption by seismic shocks related to active extensional faulting and block tilting; and

(c) gravity sliding mixing material resulting from broken lithi®ed horizons. This sequential process originated ¯at-pebble

conglomerates during early Jurassic phases of syn-rift evolution in the southern Lusitanian Basin. q 2001 Elsevier Science

B.V. All rights reserved.

Keywords: Flat-pebble conglomerates; Seismites; Syn-rift tectonics; Toarcian; Portugal

1. Introduction

Liassic outcrops in Portugal are found in three areas

(Fig. 1A). The northern one extends from ArraÂbida to

Porto (Lusitanian Basin) showing palaeogeographic

and palaeobiogeographic af®nity with West European

basins and the Subboreal province. The southern one

is con®ned to the Algarve Basin and shows palaeo-

biogeographic af®nity with the Submediterranean

province of the Tethyan Realm. An intermediate

basin exists restricted to the region south of ArraÂbida

(Santiago do CaceÂm) that would have served as an

offshore-barrier system between western European

and Tethyan basins (Mouterde et al., 1972).

In the Lower Jurassic (Lias) of ArraÂbida, three

lithostratigraphic units have been de®ned from bottom

to top:

(a) The Dagorda Formation. A thick series of red

pelites with dolomitic intercalations and evaporites

attributed to the Triassic±Lower Liassic (Hettan-

gian).

(b) The Volcanic±Sedimentary Complex. Probable

Sedimentary Geology 139 (2001) 49±70

0037-0738/01/$ - see front matter q 2001 Elsevier Science B.V. All rights reserved.

PII: S0037-0738(00)00160-3

www.elsevier.nl/locate/sedgeo

* Corresponding author. Fax: 1351-21-2948556.

E-mail addresses: [email protected] (J.C. Kullberg),

[email protected] (F. OloÂriz), [email protected]

(B. Marques), [email protected] (P.S. Caetano),

[email protected] (R.B. Rocha).1 Fax: 134-958-243345.2 Fax: 1351-21-2948556.

J.C. Kullberg et al. / Sedimentary Geology 139 (2001) 49±7050

Fig. 1. (A) Location of Portuguese Mesozoic Basins. (B) Synthetic geological sketch of the ArraÂbida sector and location of studied outcrops

(CA� Cabo de Ares; CM� Cova da Mijona; S� Sesimbra).

Early Liassic in age, identical to that known from

Santiago do CaceÂm and Algarve (these two units

crops out only around the Sesimbra diapir, their real

thickness being quite impossible to estimate).

(c) The lower part of the Achada Dolomites and

Limestones (ServicËos de Portugal (SGP), 1992).

Sinemurian±Toarcian in age.

Within the upper part of these Lower Jurassic

series, an alternation of argillaceous limestones and

marls, dated as Carixian±Toarcian, includes the

studied deposits.

The present paper focuses on the description of

particular deposits including tabular clasts, and their

interpretation taken into account assumed paleogeo-

graphy and evolution of the southern Lusitanian Basin

during the late Early Jurassic.

2. The studied sections

Pioneer studies on outcrops in the ArraÂbida region

(Fig. 1) were carried out by Ribeiro and Nery

(1866±1867), and their data were later used for the

elaboration of two geological maps (scale 1/100 000)

published in 1866 and 1867. Choffat (1903, 1905)

revisited the area, worked a 1/20 000 scale geological

map of the region, and interpreted tectonics in the

ArraÂbida Chain (Choffat, 1908). The latter work by

Choffat includes the ®rst reference to the outcrops

understudy at Cabo de Ares, Sesimbra and Cova da

Mijona. Later, these sections were cited in mapping

works (Zbyszewski et al., 1965), palaeogeographic

analysis (Mouterde et al., 1972), synthetic revisions

of Portugal's geology (Ribeiro et al., 1979; Intern.

Geol. Congress, Paris, 1980), sequence analysis of

the Triassic±Middle Jurassic in the Lusitanian Basin

(Watkinson, 1989), and recently by Marques et al.

(1990, 1994) who identi®ed deposits related to sub-

aqueous mass movements in the Upper Liassic (Toar-

cian) of western ArraÂbida.

Three geological sections belonging to the Achada

Dolomites and Limestones have been studied in the

area of Sesimbra, at Cabo de Ares, Sesimbra, and

Cova da Mijona outcrops (Fig. 1B). Intense dolomiti-

zation precluded precise petrography in the Cabo de

Ares section. The other two sections (Fig. 2), which

were studied by Choffat (1903, 1905), were

favourable for analysing the presence of ¯at pebble

conglomerates. The Sesimbra section is located on

hills east of Sesimbra (topographic sheet 464,

U.T.M.: 29 S MC 923 553) and shows a thick, mainly

carbonate succession with pronounced dolomitization

towards the top. The same sedimentary package

occurs in the Cova da Mijona section located on cliffs

6 km West of Sesimbra (topographic sheet 464,

U.T.M.: 29 S MC 851 535).

Flat-pebble conglomerates are found in the upper

part of the Achada Dolomites and Limestones in both

of these studied sections (Facies E in Fig. 2). Choffat

certainly observed these structures in the Sesimbra

region when referred to ªpetites concreÂtions arron-

diesº in the upper part of his Bed K (Choffat, 1905,

p. 137). Choffat's observations were considered by

Zbyszewski et al. (1965, pp. 105±106, Bed 10) who

cited unpublished notes from Cova da Mijona made

by Choffat: ªcalcaÂrio dolomõÂtico rijo, cinzento,¼

imitando em parte uma brecha com elementos angu-

lososº (hard, grey dolomitic limestone,¼ partly simi-

lar to a breccia with angular elements). This

stratigraphic interval (Choffat Bed K) with ªconcreÂ-

tions arrondiesº overlies Bed J labelled by Choffat

(1903, p. 82; 1905, pp. 136±137� Bed 8) in Zbys-

zewski et al. (1965, pp. 106±107), which yielded three

fragments of Ammonites communis. The overlying

Bed L of Choffat (Bed 11 in Zbyszewski et al.,

1965, p. 106) provided a fragmentary cast of a large

specimen of Ammonites bifrons?

No specimens collected by Choffat have been iden-

ti®ed in collections housed at the Instituto GeoloÂgico e

Mineiro in Lisboa. However, Mouterde and Rocha

collected fragments of Dactyliocerasgr. semicelatum,

initially referred to as DDactylioceras communis

(Zbyszewski et al., 1965, p. 107), from the same hori-

zons sampled by Choffat. These ammonites belong to

the base of the Lower Toarcian (Polymorphum Zone).

Since Choffat's work, no other ammonites have

been collected from horizons overlying those showing

¯at pebbles. Based on the accurate knowledge by this

author of Toarcian ammonites, we assume that the

large incomplete ammonite cast mentioned above

belonged to Hildoceras sp. Given the fact that Hildo-

ceras bifrons is rarely reported from the Lusitanian

Basin, the reference to this species proposed by

Zbyszewski et al. (1965, p. 106) is less probable.

Whatever the case, the identi®cation of this ammonite

J.C. Kullberg et al. / Sedimentary Geology 139 (2001) 49±70 51


Fig. 2. Studied sections. Lithologic columns showing selected stratigraphic features. Encircled numbers refer to the location of respective

®gures.

at the genus level is enough to recognize the Bifrons

Zone of the Middle Toarcian. Therefore, horizons

containing ¯at-pebble conglomerates, which are

unknown elsewhere in the Lusitanian Basin, would

belong to the Levisoni Zone, and hence correlated

with the ªCalcaÂrios em plaquetasº that are found

throughout the basin north of the Tejo river (Mouterde

et al., 1972; Duarte, 1990; Rocha et al., 1996).

2.1. Type-facies

In spite of intense dolomitization that make thin


Fig. 3. Type-facies: (A) pelletoidal wackestone showing benthic foraminifera recrystallized and fragmented pelecypod (oyster) with encrusting

micrite-walled microproblematica; (B) micritized grain coated by algae in intraclastic pelletoidal wackestone; (C) pelletoidal intra-bioclastic

lime grainstone. Note aggregate grains (lumps) and peloids, encrusting algae on micritized ooids and/or lumps, and transverse sections of

echinoid spines; (D) longitudinal section of crinoid stem coated by algae in pelletoidal and intraclastic wackestone; (E) imbricated laminites

(ªmicro ¯at pebblesº) made of mudstones in horizon with ¯at pebbles.

section analysis dif®cult, recognition of six marine

facies (A±F) was possible.

Type-facies A (Fig. 3A) Ð Microsparitic wackes-

tones and packstones. Bioclasts are mainly bivalves

such as pectinids (Chlamys), cardiids (Cardium) and

oysters. Gastropods, belemnites, brachiopods and

echinoderms are rare. Ostracods and benthic forami-

nifera also exist. Bioclasts are sparsely distributed or

locally concentrated in horizontal layers 1 cm-thick.

Non-skeletal grains consist of mud peloids, faecal

pellets, algal peloids, and very rare aggregate grains.

Lumpy deposits presumably related to pervasive but

indeterminable bioturbation exist. Laminations and

gradations are the most common sedimentary struc-

tures. Bed thickness varies between 15 and 70 cm

(40 cm in average).

Type-facies B (Fig. 3B) Ð Wackestones and pack-

stones similar to those of Type-facies A, but skeletal

debris are mainly made up of bivalves, foraminifera

and algae (Dasycladaceae and Gymnocodiaceae)


Fig. 3. (continued)

accompanied by belemnites, brachiopods, echino-

derms, annelids (serpulids) and bryozoans. Scarce

hermatypic corals and sponges occur. Bioclasts are

disseminated or concentrated in horizons parallel to

bedding. Non-skeletal grains occur as abundant

peloids and faecal pellets. Intraclasts are rare. Discon-

tinuity surfaces, some of them ferruginized, exist.

Bioturbation similar to that reported for Type-facies

A is common and locally pervasive. Abundant lami-

nations and rare ripple marks are present. Stylolites

are usual. Bed thickness varies between 30 and 70 cm

(40 cm in average).

Type-facies C (Fig. 3C) Ð This facies is repre-

sented by well sorted, ®ne to medium grainstones,

which are more or less oolitic and bioclastic. Skeletals

are very abundant and diversi®ed: bivalves, gastro-

pods, brachiopods, echinoderm plates and spines,

annelids (serpulids), benthic foraminifera, ostracods,

and algae. Ooids have nuclei of foraminifera, ostra-

cods, bivalve and echinoderm fragments, and

commonly show concentric and radial structure

more or less masked by micritization. Mud peloids,

faecal pellets, pellets and rare aggregate grains exist.

Bioturbation is common, including vertical to sub-

vertical burrows (Skolithos). Quartz grains are sparse.

Planar and festoon cross-laminae are common, and

ripple marks secondary. Type-facies C has medium-

thick beds (30 cm at Cova da Mijona and 40 cm at

Sesimbra sections).

Type-facies D (Fig. 3D) Ð Inter-bedded argillac-

eous and silty limestones, bioclastic wackestones,

lime mudstones and marlstones. Skeletal debris are

mainly belemnites and brachiopods (terebratulids

and rhynchonellids). Bivalves and gastropods are

common in some horizons, and ammonites are rare.

Abundant carbonaceous plant remains and secondary

intraclasts are present in argillaceous limestones.

Intercalated marls yielded fossils, carbonate nodules

and carbonaceous plant remains. Limestone beds

show planar top and bottom surfaces, and their thick-

ness varies from 5 to 50 cm (23 cm in average); marly

horizons are 2±60 cm-thick (20 cm in average).

Type-facies E (Fig. 3E) Ð Laminated mudstones

and bioclastic wackestones rich in fossils similar to

those in Type-facies B. Skeletal debris predominantly

consist of bivalves, rare gastropods and brachiopods,

echinoderm plates and spicules, annelids (serpulids),

benthic foraminifera, ostracods, and algae. Non-skele-

tal components consist of silt-sized pellets. Lamina-

tions are abundant, but neither planar cross-bedding

nor asymmetric ripples were found in Type-facies E.

Stylolites are frequent. Bed thickness varies between

10 and 130 cm (65 cm in average) at the Cova da

Mijona section and between 15 and 80 cm (58 cm in


Fig. 3. (continued)

average) at the Sesimbra section. Type-facies E is the

only facies showing ¯at pebble conglomerates (see

below).

Around twenty depositional units have been recog-

nized, the lower fourteen being better preserved

(enlarged section in Fig. 2), some of them include

minor ones, and the majority are bounded by thin

marly inter-beds. In these depositional units, pebbles

concentrate in their lower parts or alternatively appear

from bottom to top. When pebbles show low dense

packing (scattered within the matrix) they are more

rounded. Depositional units that do not show pebbles

are rare.

Type-facies F Ð Dolomitized limestones mostly

formed by secondary alteration of peletoidal pack-

stones, grainstones and wackestones. Bioclasts are

bivalves, gastropods and brachiopods. Peloids and

ooids are also common. Cross, planar, and wavy lami-

nae and fenestral structures are usual.

2.2. Synsedimentary deformation structures

In the lower part of the Achada Dolomites and

Limestones, common synsedimentary deformation

structures affecting 10 to 20 m-thick sections of

considerable lateral extension (.500 m) were

observed in the three outcrops studied. Disturbed

set-beds show plastic or brittle deformation of

bedding, as well as reworking, the most common

structures being ¯at pebbles, slump-sheets and intra-

formational conglomerates, and breccias.

2.2.1. Flat pebbles

These structures are the most common in the

Achada Dolomites and Limestones. They occur in

regular layers 30 cm to mainly 80 cm-thick. The

most striking feature of these layers is the abundance

of ¯at pebbles, many of which are highly tabular or

subrounded reaching 6 cm in length and 2 cm in thick-

ness. Thickness and angular to subrounded shape

suggest that pebbles originated through reworking of

consolidated to semi-consolidated thin-bedded calcar-

eous deposits outcropping in the area. More dense

packing of pebbles relates to more angular and

sharp edged pebbles. Some micropebbles (,1 cm)

embedded in the matrix show plastic deformation.

Texture in pebbles is pelmicrite and most commonly

micrite (unfossiliferous pelitic limestone) in accor-

dance with the nature of limestones directly underly-

ing these intra-formational calcirudites. Pebbles rarely

contain calcarenite intraclasts.

The arrangement of microclasts shows intraclast

orientations from bedding-parallel (Fig. 4A) to

random (Fig. 4B), occasional imbrication of pebbles,

and edgewise breccia fabrics. Pebbles are con®ned to

the lowermost part of the calcarenitic layer or, alter-

natively, distributed throughout the whole bed. In

addition, small and irregular concentrations of

pebbles occur locally. Matrix embedding the ¯at

pebble conglomerates is calcarenitic.

2.2.2. Slump sheets and intra-formational

conglomerates and breccias

Slump sheets and intra-formational conglomerates

and breccias are common at Cova da Mijona and

Sesimbra sections. They consist of broken slump-

folds forming slabs accompanied by isolated lime-

stone blocks embedded in the intra-formational

conglomerate, which is mainly composed of ¯at

pebbles and sparse skeletals (Fig. 5). Layers contain-

ing slump structures are 80±100 cm thick. The

passage of slump structures into intra-formational

conglomerates is observed locally. Sedimentary

slabs conformable with the original bedding make

the understanding of any relationships with coastal

erosion unclear. The bottom surface in intra-forma-

tional breccias and conglomerates is normally ¯at,

but sometimes is uneven, showing crests and furrows

due to scouring on the substratum with resulting

deposition of intraclasts over the erosional surface.

However, intra-formational breccias laterally pass

occasionally into pure calcarenites. Lithological

homogeneity between pebbles and the substratum of

some breccias was also identi®ed. The matrix of brec-

cias is calcarenitic and mainly made up of organic

detritus that contains fossils bearing traces of

mechanical wear reworked from shallower environ-

ments. Skeletals are much rarer in intra-formational

breccias.

Flat pebbles are derived from the underlying slump

sheet, and/or contemporaneous equivalents, their

morphology and thickness being identical to the

subjacent layers. They also display parallelism to

the lamination. Tops of slump sheets are indistinct

as these sheets grade into laminated limestone. Strata

containing slump sheets (reaching 60 cm in length and


20 cm in thickness) shows these to be parallel to

bedding at the base, followed by breccia-like horizons

due to the presence of ¯at pebbles, and then lamina-

tion towards the top. All this re¯ects decreasing

energy in depositional conditions. This stacking

pattern could also occur during reworking of mud¯ow

deposits experiencing internal differential move-

ments.

2.3. Depositional history

Deposits containing ¯at pebbles crop out in a

restricted area and are unknown in any other region

and age in the Lusitanian Basin. They therefore are

particular deposits that may be the expression of

special events restricted in space and time, but accord-

ing with the style and timing of basin structuring.


Fig. 4. Flat-pebbles arrangement parallel to bedding (A) and random (B). Coin size in 4A is 3 cm.

According to Watkinson (1989), during Triassic

and earliest Jurassic (Hettangian) times, the ArraÂbida

area was a saline basin, the dolomitic member of the

Dagorda/Pereiros Fm. representing tidal-¯at deposi-

tion on the basin margins. Initial deposition in carbo-

nate-ramp conditions corresponded to the lower part

of the Achada Dolomites and Limestones (Sinemur-

ian±Toarcian). Subsequently, the beginning of open

marine regimes was heterochronous throughout the

Lusitanian Basin.

The type-facies assemblage differentiated above

clearly points to a shallow-marine shelf lagoon as

the depositional environment for the lower part of

the Achada Dolomites and Limestones. In some

cases, the micritic matrix and the high content of

micro-bored bioclasts, microfossils (ostracods, fora-

minifera), algae and dominant molluscs, are charac-

teristic (Type-facies A). Bivalves such as pectinids

(Chlamys), cardiids (Cardium) and oysters indicate

warm and nutrient-rich shallow waters in moderate-

energy coastal areas showing local cohesive

substrates. The presence of rare belemnites, echino-

derms and brachiopods indicates that lagoonal condi-

tions were not fully restricted. In addition, the record

of foraminifera (Lituolidae, Globigerina-like forms

and Dentalina) indicating shallow marine biotopes

accords with occasional connections with marine

waters of around normal salinity.

Low sedimentation rates and bottom conditions

favoured intense bioturbation (Rhizocorallium).

Bioclast layering and laminations covered by ferrugi-

nized surfaces (minor discontinuities), and common

peloids indicate episodes of decreasing energy

between minor depositional events unaffected by

burrowers. Thus, Type-facies B deposits are inter-

preted as resulting from erosion affecting carbonate

build-ups growing on the shelf, which were

dominated by coral±algal debris-facies (dominantly

Dasycladaceae and Gymnocodiaceae), as well as

sponge spicules and fragments (Type-facies A and

B). Besides coral and algae remains, the mixed

bioclastic facies contains spines and ossicles of

crinoids, complete and incomplete shells of brachio-

pods, belemnites and common bivalve fragments

(particularly oysters and Lithophaga) as well as serpu-

lids. Gastropods, bryozoans and foraminifera also

occur.

Non-skeletal elements are intraclasts, peloids and

aggregate grains. Cephalopods (belemnites), brachio-

pods and crinoids (non-reefal elements) probably

derived post-mortem from somewhat deeper or more

open-sea settings. Further components, like gastro-

pods and bivalves, seem to be admixed from a differ-

ent environmental zone. Nuclei in ooids mainly

consist of bioclasts (Type-facies B and E).

The bioclastic, oolitic and peloidal packstones and

grainstones, with cross-laminae (planar and festoon)

and some ripple marks, were deposited under variable


Fig. 5. Slump sheets. Note convolute bedding caused by displacement of one of the torn-out slabs and a relatively gradual upward transition

from ¯at-pebble to laminated horizons and then bioclastic wackestones.

but relatively high energy in the sub-tidal zone. Bioclas-

tic nuclei in ooids dominate, although grains of detrital

quartz also occur. Ooids with composite-multiple nuclei

are uncommon. The high rate of admixed cortoids

suggests ooids originated in a relatively low-energy

environment, e.g. con®ned lagoons. The occurrence of

sponges, brachiopods, bryozoans, dasycladaceans and

rare coral fragments, points to the vicinity of para-reefal

environments. On the other hand, variable fragmenta-

tion and terrigenous (quartz) input are recorded. The

oolitic bioclastic packstone/grainstone complex shows

low-angle cross lamination and rapid wedging of indi-

vidual laminae-sets, the better-sorted sediments prob-

ably representing bar-systems within this lagoon

(Type-facies C). These features are in accordance with

lateral accretion deposits, and could relate the cortoid-

oolitic packstones and grainstones with axial channel

deposits within shallow lagoonal environments.

However, outcrop limitations make dif®cult any conclu-

sive interpretation.


Fig. 6. Conjugated sets of within-bed synsedimentary extensional faults. Note bottom-wards bending of more spaced faults that become parallel

to bedding (ªlistricityº), as well as ductile deformed horizon sealing these structures.

Fig. 7. Close-up view of top-horizon in Fig. 6 showing reworked deposits embedded in laminated horizon below a micro-graben sealed by

horizon with synsedimentary ductile deformation.

The presence of pelagic cephalopods (Type-facies

D) indicates relative open seas nearby the area

studied. Off-shore facies in the area are characterized

by the presence of lime mudstones and bioclastic

limestones inter-bedded with argillaceous limestones

and marls. Pelagic bivalves and cephalopods (belem-

nites and ammonites), together with echinoderm

bioclasts, indicate the increasing in¯uence of more

pelagic and normal marine conditions, which

occurred just before and after deposition of Type-

facies E containing ¯at-pebble conglomerates.

Environmental factors controlling deposition of

Type-facies E were related to basin instability.

Three sets of depositional units (E-1 to E-3 in

enlarged section in Fig. 2) are identi®ed, the middle

one showing the most intense synsedimentary defor-

mation. Outcrop limitations make dif®cult the precise

interpretation of the lowermost E-1 stratigraphic inter-

val, but similarity with E-3 is envisaged from local

observations. Therefore, the roughly symmetrical

pattern accords with a paroxysmal tectonic distur-

bance (E-2) that began and ended progressively.

Depositional units with ¯at pebbles are simple (E-1,

E-3) and secondarily complex (E-2), the latter show-

ing olistoliths and/or boulders that also contain ¯at

pebbles. No sedimentary structures related to currents

accord with mass-¯ow conditions from near in situ

redeposition (lower part of Fig. 4A) to low-order

displacements (upper part of Fig. 4B). Dense packed

to ¯oated pebbles indicate variations in the affected

number of sedimentary horizons, the amount of

energy involved in a particular event, and the resulting

degree of mobilization affecting deposits undergoing

advancing but variable lithi®cation. Given fossil

content and distribution in Type-facies E, the lack in

shell-beds indicates that no re-suspension events

affected the substrate, at least as a prime factor

controlling deposition as recognized from the mature

record analysed in the sections studied.

2.4. Structural analysis

Apart from sedimentological observations, the

recognition of normal faults, tension gashes in

pebbles, micro-fracturing in ¯at-pebble horizons,

and olistoliths shows relationships with extensional

tectonics.

2.4.1. Extensional faults

Macroscopic and mesoscopic extensional faults are


Fig. 8. Filling of extensional fracture with mud and sparse ¯at pebbles (arrow) showing evidence of synsedimentary faulting.

Fig. 9. Stereograms (Schmidt±Lambert projection, lower-hemisphere) of bedding and tectonic structures, with removal of the late effect of

tilting caused by halokinetic movements of the Sesimbra diapir (348, as shown in stereogram A): (A) bedding; (B) extensional faults; (C)

tension gashes; (D) micro-faults; (E) micro-faults within the major olistolith. Azimuths are expressed through Right Hand Rule.


observable in the area. The ®rst are mappable but not

represented in the vertical section of Fig. 2. They

probably worked when ¯at pebble conglomerates

originated, but it is dif®cult to strictly relate them to

the ¯at-pebble deposits since they reactivated

throughout the whole Jurassic.

Mesoscopic extensional faults are observable at the

outcrop, some of them affecting more than one bed. At

the top of the section they occur, generally, restricted to

one bed and controlled by bed thickness (Fig. 6). These

extensional faults are of metre±decimetre order,

frequently conjugate and show listric geometry. Some-

times differential deformation within beds can be

observed: centimetre-order displacements in the lower

parts due to brittle deformation, while ductile deforma-

tion with micro-bending in the upper part affected

directly overlaying semi-lithi®ed carbonate horizons

(Fig. 7). The in®lling of narrow micro-grabens by thin

and semi-lithi®ed carbonate horizons determined their

local failure and the variable orientation of clasts, verti-

cal positions included. In®lling of some fault-related

pockets includes ¯at pebbles (Fig. 8). On the whole,

fault orientation is rather consistent with an approxi-

mately E±W extensional regime (Fig. 9B).

2.4.2. Tension gashes

Tension gashes in ¯at pebbles are always sub-

perpendicular to their larger axes, irrespective of

pebble orientations (Fig. 10). Evidently, these frac-

tures already existed before the reworking events

that caused pebble accumulation (conglomerates).

Furthermore, lithi®cation of the thin carbonate hori-

zons that sourced the ¯at pebbles was faster than that

of the embedding deposits.

As shown in Fig. 10, a younger generation of

tension gashes was recognized in the outcrops studied

through fractures that affected both pebbles and sedi-

ments surrounding them. This demonstrates their late

generation, most probably linked to activity in the

Sesimbra diapir, as proved by their orientation that

®ts well the fracturing induced by halokinetic move-

ments in the area (Fig. 9C).

2.4.3. Micro-fractures

Two types of micro-fracturing were observed in

horizons related with ¯at pebble conglomerates: (a)

more or less straight and continuous micro-fractures,

which affected the matrix exclusively, showing

anastomosed ends; and (b) sharp/angular micro-frac-

tures that show noticeable refraction in more micritic

levels, were more penetrative in argillaceous frac-

tions, resulting in polygonal frames in the whole set

of affected horizons.

Two processes are envisaged to result in micro-

fracturing: (a) the displacement of semi-lithi®ed sedi-

ment; and (b) disturbances affecting semi-lithi®ed

deposits without displacement. The parallelism

between the more angular pebbles and micro-fractures

affecting the sediments accords with displacement

(process a) forcing relative movements of larger

elements along sliding-planes within the mass (Fig.

4B), while more rounded pebbles rolled over. Alter-

natively, carbonate deposits disrupted in place

(process b) and originated two linear and approxi-

mately symmetric patterns showing low angles with

respect to bedding (Fig. 11).

As could be expected (Fig. 9D), the general orien-

tation of resulting structures is very similar to that

found in the extensional faults, their genetic associa-

tion being obvious. Differential phases of deposition

and synsedimentary deformation are also evident in

Fig. 11. In situ deformation during early lithi®cation

preceded deposition of reworked sediments that

included ¯at pebbles. Thus, differences in lithi®cation

micro-fracturing, and transport, affected deposits

directly below and above particular bed-surfaces.

2.4.4. Olistoliths

Two olistoliths were recognized in relation to ¯at

pebble conglomerates.

The ®rst one �2:5 m length £ 2 m wide £ 60 cm

high; Fig. 12) exhibits type-b micro-fracturing as

described above. The comparison between the mean

orientations of micro-fractures in this olistolith (Fig.

9E) and those above referred to as type-b, shows that

this block experienced a 208 sinistral rotation, without

tilting, during displacement. Therefore, micro-fractur-

ing within this olistolith preceded both unrooting and

mobilization.

The second olistolith �60 cm length £ 40 cm wide;

height unknown), exposed on a bed surface, moved

from west to east and left a ªtailº that corresponds to

the shadow-zone in which internal fractures created

by dragging-suction and removal of the embedding

sediment are recognized (Fig. 13). Displacement of

this olistolith was perpendicular to its longest axis


suggesting a relatively short displacement (a few

meters at the most), favoured by its arched front.

The occurrence of olistoliths can be related to set-

beds collapse and then sliding along fault scarps.

Although the majority of faults affected only one

bed, and rarely two or more, with only centimetre-

to metre-order displacements, the existence of larger

faults cannot be disregarded. The orientation of all the

analysed structures (Fig. 9) points to an approximately

E±W extensional tectonic regime. Signi®cantly, this

accords with the orientation of major faults control-

ling the structure of the Lusitanian Basin (Ribeiro et

al., 1979; Soares et al., 1993), thus favouring the inter-

pretation of an eastwards regional tilting of blocks.

3. Discussion

During the Mesozoic, the Iberian Peninsula was

located in a particular hinge situation between devel-

oping rift-systems in the proto-Atlantic (West) and

Tethys (South±Southeast) that formed the Portuguese

Lusitanian and Algarve Basins, respectively. While

Palaeozoic basement rocks of the Iberian Meseta

bordered the Lusitanian Basin on the north, east,

and south, uplifted Hercynian rocks (today the Berlen-

gas islands) lay to the West, and a permanent seaway

connected the Lusitanian and Proto-Atlantic Basins

south-westwards throughout the Sintra/ArraÂbida area

(Ribeiro et al., 1979).

The origin and development of the Lusitanian

Basin under extensional tectonics during the Triassic

rift phase was a subject revisited during eighties (from

Lancelot, 1982 to Montenat et al., 1988; Wilson et al.,

1989). This basin evolved as a passive Atlantic-type

basin (Ribeiro et al., 1990), basically structured by

tardi-Hercynian faults orientated N±S and NNE±

SSW, which largely were responsible for basin struc-

turing. During Mesozoic distension, these faults reac-

tivated as approximately E±W extensional trends

(Ribeiro et al., 1990; Kullberg and Rocha, 1991).

The area studied, near the eastern margin of the Lusi-

tanian Basin, as well as the entire basin (Wilson et al.,

1989), was under the in¯uence of deep listric faults,

that forced half-graben tilting to the east. Overall, the

half-grabens deepened progressively westwards, as

evidenced by the distribution of carbonate facies indi-

cating a westerly-dipping ramp (Wilson et al., 1989);

low subsidence favoured the gradual sinking of the

basin during the Triassic and Early Jurassic, with no

signi®cant local tectonics. Mouterde et al. (1972)

concluded that, during the Toarcian, depocentres

receiving the thickest deposits were elongated

NNE±SSW and located west of Coimbra.

The vicinity of the Sesimbra diapir to the sections

studied supports the hypothesis of halokinetic


Fig. 10. Tension gashes: the earliest generation only affected pebbles, being perpendicular to their largest axes; the latest generation is larger,

affected the whole bed and is perpendicular to bedding.

movements in¯uencing deposition in the area. Monte-

nat et al. (1988) considered that halokinetic phenom-

ena started at least during the Toarcian, this being the

only autocyclic factor that signi®cantly disturbed

background deposition. However, this hypothesis

does not apply in our case study, because the centre

of the Sesimbra diapir was located SW of the Sesim-

bra section. Therefore, its in¯uence on sedimentation

in nearby areas, would follow a NE to ENE trend; this

is clearly oblique to the orientations of bedding and

structures in the area as shown on the stereograms

(compare Fig. 9A and B). In addition, Kullberg and

Rocha (1991) reported evidence pointing to a Cretac-

eous age for the ®rst halokinetic movements of the

Sesimbra diapir. Thus, the in¯uence of early tectonic

pulses as factors determining the origin of the studied

deposits seems unequivocal. Some traces of tectonic

events were directly recorded, but the occurrence of

fracturing is not self-explanatory in terms of the

depositional dynamics, especially that related to ¯at-

pebble production.

Flat-pebble conglomerates are rather particular


Fig. 11. Example of disrupted bed exhibiting two linear patterns of fractures (upper left to lower right dominant) at a low angle to bedding. Note

overlying undisturbed sediments showing ¯oated pebbles separated from reworked and pebble supplying horizons.

Fig. 12. Large olistolith exhibiting type-b micro fracturing and forcing intense deformation on sliding surface.

deposits that have been described from various sedi-

mentary environments. Braun and Friedman (1969)

interpreted the origin of ¯at-pebble conglomerates in

tidal muds through early lithi®cation and disruptive

desiccation due to subaerial exposure, followed by

reworking and transportation of pebbles during subse-

quent immersions, storm action included (Shinn,

1983). Many authors considered that intra-forma-

tional ¯at-pebble conglomerates resulted from storms

(Jansa and Fischbuch, 1974; Jones and Dixon, 1976;

Seilacher, 1991; Sepkoski et al., 1991). Sepkoski

(1982) pointed out the conditions for the origin of

¯at pebbles: episodic deposition and rapid cementa-

tion of thin permeable carbonate layers, separated by

muddy intercalations, and then erosion and reworking

due to intense storms.

Sedimentary structures somewhat similar to those

analysed in this paper were noted by Valenzuela et al.

(1986), who interpreted them as resulting from storm

action. Intra-formational ¯at-pebble conglomerates

and breccias, and slump sheets, very similar to those

found in the Sesimbra area were described by

Szulczewski (1968) from the Upper Devonian lime-

stones of the Holy Cross Mts. in Poland. These depos-

its were originated by sub-aqueous sliding, according

to this author, who considered them, dynamically, as

slide conglomerates that were involved in mass move-

ments associated not only with mechanical wear of

reefs but also with earthquakes during early phases

of the Variscan orogeny. Kazmierczak and goldring

(1978) revisited these deposits, favouring the in¯u-

ence of storm or tsunami action as trigger factors,

but they also considered the possibility of seismic

shocks as a mechanism that may have contributed to

¯at-pebble production through loosening along less

cohesive laminae.

Debris-¯ow deposits described by Cook and

Mullins (1983) are similar to ¯at-pebble conglomer-

ates recognized in the Sesimbra area. Mud ¯ows were

also considered as an explanation for the origin of

intra-formational conglomerates containing ¯at

pebbles in the Middle Triassic of the Holy-Cross

Mts. in Poland (Bialik et al., 1972).

In summary, four hypotheses could explain the

origin of ¯at-pebble conglomerates resulting from

reworking of lithi®ed thin carbonate layers: (a)

subaerial desiccation and remobilization after immer-

sion; (b) storms; (c) tsunami events; and (d) seismic-

shocks.

(a) The absence of sedimentary structures such as

mud-cracks, fenestrae, and vugs is inconsistent with

the origin of the studied ¯at pebbles by desiccation

through subaerial exposure. The occurrence at Sesim-

bra of olistoliths in horizons containing ¯at pebbles

also refutes this hypothesis.

(b) Most authors agree that storms are a major

factor responsible for the genesis of ¯at-pebble depos-

its. However, in the studied case this hypothesis seems

much less evident. According to Scotese and Denham

(1987) Iberia was located around 258N latitude during

the Early Jurassic, when westwards hurricanes were

essentially con®ned to the Tethyan ocean (Marsaglia

and Klein, 1983). The ArraÂbida sector was probably

sheltered from the effect of major storm systems

proceeding from the east, due to the physiography

of the Iberian Meseta. Moreover, the rarity of sedi-

mentary structures like hummocky cross-strati®cation

and grain sorting, as well as shelly beds, which are

considered typical components in shallow tempestites


Fig. 13. Top surface of a bed showing small olistolith (1), with ªtailº

(2) indicating the direction of movement.

(Aigner and Reineck, 1982; Duringer, 1984), points to

factors other than storms for ¯at-pebble production in

the Sesimbra area. In addition, the envisaged shallow-

marine shelf lagoonal setting makes improbable any

relationships between storms and the olistoliths within

the ¯at-pebble deposits at Sesimbra.

(c) Consensus following Tinti (1987) points to

submarine landslides, submarine volcanic activity,

and large submarine earthquakes related to sudden

dip-slip motion affecting sea ¯oors through faulting

as the main causes, in order of incidence, for the

generation of tsunami events.

Evidences of mass transport (submarine slides),

cannot be invoked, simultaneously, as the cause and

the effect of tsunami. In fact, according to Tinti

(1987), tsunami generated by this mechanism alone

are extremely rare. Generally, tsunami may be attrib-

uted to a combined origin, since the generating land-

slides or rock-falls are in turn triggered by volcanic

eruptions or earthquakes (see below). In the case

studied, volcanic activity can be ruled out, since no

volcanic episode of late Early Jurassic age has been

recognized anywhere in the Lusitanian Basin (Barros,

1979; Ziegler, 1988).

In contrast, sudden dip-slip motion seems the most

likely explanation for the occurrence of tsunami

events that eventually would produce ¯at-pebble

conglomerates in the Sesimbra area. However, geody-

namic and sedimentological data make the application

of this hypothesis dif®cult (see below). The over-

whelming majority of historic tsunami, particularly

those with large run-up recorded in the Paci®c

Ocean are due to high-magnitude earthquakes asso-

ciated with large dip-slip movements at active plate

margins.

Comparing magnitude scales for tsunami events

according to Iida (1963) and ML magnitudes in

Murty and Loomis (1980), Table 1 was constructed

to correlate earthquake magnitude with the largest

wave generated. This is an excellent approximation

to Whittow's (1980, in Dawson et al., 1988) simpli®ed

table and also to Abe's (1979) expression. In the ML-

magnitude scale it is not the height of the largest

waves that is considered, but the energy of the tsunami

and qualitative ªremarksº. The 0 (zero) ML-

magnitude corresponds to the occurrence of an

ªobservable transoceanic tsunamiº, but when

converted to a maximum run-up elevation value,

through Table 1, it gives an insigni®cant value

(,0.30 m). It is the 6 ML-magnitude, which corre-

sponds to a ªsigni®cant tsunamiº that correlates with

magnitude 1 in Iida's classi®cation, that shows the


Table 1

Comparison between tsunami and earthquake magnitudes and the corresponding wave run-up elevations

ML�Murty±Loomis tsunami magnitude scale; Iida TM� Iida tsunami magnitude scale. (1) and (2) after Murty and Loomis (1980); (3)

calculated values for earthquake energy assuming values ten to one hundred times larger than tsunami energy; (4) calculated earthquake

magnitudes from (3) according to Bolt (1988). Shadowed area indicates the most probable range of earthquake magnitudes in extensional

regimes (Gubbins, 1990) capable of generating tsunami events; (5) values calculated from (2) using equation m � 2:66 1 1:66 log ��2�=1023�:Values in parenthesis are below the minimum value in Iida's classi®cation, but are considered for discussion purposes; (6) h � 10 E �m=3:32�(after Iida, 1963). The authors emphasize that depositional conditions and the local record of the ¯at pebble conglomerates studied accord with

factors others than tsunami or tidal wave action (see text), the latter being used in oceanographical sense as the reference for long, propagating

shallow-water waves of tidal period generated by tide-generating forces and modi®ed by the Coriolis force, bottom friction, and sea¯oor

topography

1 2 3 4 5 6

ML Tsunami energy (ergs) Earthquake energy (ergs) Earthquake

magnitude

(moment

magn. MW)

Iida TM (m) Maximum run-up (m)

10 1024 1025±1026 8.8±9.5 4.3 . 20

8 1023 1024±1025 8.1±8.8 2.7 6.3±8.0

6 1022 1023±1024 7.5±8.1 1.0 2.0±3.0 Tidal wave run-ups

4 1021 1022±1023 6.8±7.5 20.7 0.63±0.75

2 1020 1021±1022 6.1±6.8 (22.3) 0.20±0.30

0 1019 1020±1021 5.5±6.1 (24.0) 0.06±0.30

24 1017 1018±1019 4.1±4.8 (27.3) 0.01±0.30

signi®cant effects of the largest waves (2±3 m), but

their differentiation from those originated by large

tidal waves, constrained by tidal regime and shoreline

geometry, is dif®cult. Expected earthquake

magnitudes from these higher values are around

7.5±8.1.

Under extensional tectonics in rift-basins, large

earthquakes can occur (Gubbins, 1990). However,

even assuming earthquakes up to magnitude 7, locally

generated tsunami events should be extremely unli-

kely in the Lusitanian Basin during the Early Jurassic.

This conclusion is supported by: (1) low subsidence

during this time in the area, as shown by the sedimen-

tary record; (2) near-absence of other tectonic distur-

bances, implying slow and progressive distension; and

(3) facies attributable to shallow lagoonal environ-

ments, thus diminishing the possibility of locally

generated tsunami in the Sesimbra area. The last

accords with the proposed relationship between

tsunami magnitude and water depth over an earth-

quake epicentre, as Iida (1963) considered: larger

tsunami are generated in deeper waters. Any large

tsunami causing ¯at-pebble conglomerates at Sesim-

bra would have had a distant origin and a wide regio-

nal impact, the record of which is unknown from

contemporaneous deposits elsewhere in the Lusita-

nian Basin.

Sedimentological data concerning tsunami deposits

point the same way. Throughout a tsunami's path,

ocean bottoms are strongly disturbed and, in the

offshore zone, considerable quantities of sea-¯oor

sediment are carried as suspended matter (Dawson

et al., 1991a). After its passage (or passages, as a

tsunami can be composed of more than a single

wave), reworked material settles according to Stokes

Law (Duringer, 1984; Dawson et al., 1991a). There-

fore, grain-size analysis of sediment accumulations

can provide information not only about the number

of tsunami waves that struck a coastal area, but may

also provide information about the relative magnitude

of the successive waves (Dawson et al., 1991a). On

the other hand, during backwash, very strong erosive

currents produce local channelling with transportation

and considerable re-deposition of sediments seawards

(Dawson et al., 1991b).

Sedimentological observations in the deposits

containing the ¯at pebbles studied (Type-facies E)

show neither grain sorting nor cyclic sedimentation,

nor evidence of signi®cant channelling and basinward

(westward) transportation of sediment. Transport

direction was proved to be eastward. Moreover, the

tsunami hypothesis cannot explain disruptions and

micro-fracturing, as described above.

According to all the above, neither geodynamic

conditions nor the geological record favour the

assumption that large tsunami reached the studied

area.

(d) Compared to a tsunami's energy, the energy

related to associated earthquakes is ten to one hundred

times higher (Iida, 1963). Thus, a 6.5-magnitude

earthquake can be quite destructive, but the poten-

tially generated tsunami would have a wave run-up

of 0.3 m. Seismic activity (earthquakes) is well known

in continental rifts, such as the Lusitanian Basin was

during the Early Jurassic. In such a situation, Boillot's

(1983) statement of weak but constant and super®cial

earthquakes characterizing continental-rift evolution

is signi®cant.

Scheidegger (1975, in Dawson et al., 1988)

proposed a minimum magnitude rarely lesser than

6.5 to induce submarine landslide activity, such as

demonstrated for the Storegga slides on the Norwe-

gian continental margin during the Holocene. Jansen

et al. (1987) indicated that a magnitude 5 earthquake

could also be responsible for that, although other

factors such as ice loading and the presence of gas,

gas hydrates, and higher than normal pressure in

pore-waters, could be also involved (Bugge et al.,

1988).

The existence of synsedimentary tectonics (hori-

zons showing intra-formational extensional faults

and those with micro-sliding planes, and type a±b

micro-fractures) controlling deposition of the above

described ¯at pebbles and related deposits is

unquestionable. The ªstrong break-upº identi®ed

by Duarte (1990) on the basis of signi®cant deposi-

tion of distal turbidites (ªCalcaÂrios em plaquetasº)

during the Early Toarcian (Polymorphum±

Levisoni±Chron boundary) 150 km northwards in

the Lusitanian Basin is signi®cant. Soares et al.

(1993) considered these turbidites to be related

with a ªstrong downing and deepeningº affecting

the North-Lusitanian sub-Basin during the Early

Toarcian Levisoni Chron. Regional correlation

through precise biostratigraphy allows us to inter-

pret the studied deposits at Sesimbra as the local


variant of tectonically induced sedimentary changes

affecting the Lusitanian Basin in different palaeogeo-

graphic settings. Moreover, the Toarcian corresponded

to a rifting phase in western Tethys, so that seismic

activity would be expected. Although global eustasy

(Haq et al., 1988) accords with shallowing-up sequences

during the earliest Toarcian, it cannot explain the docu-

mented deformation in the lowermost Toarcian deposits

at Sesimbra. Therefore, tectonic imprint operated in the

Sesimbra area within the context of tectono±eustatic

interactions envisaged by Soares et al. (1993) for the

northern Lusitanian Basin during the Early Toarcian

Levisoni Chron. Hence, the case studied supports the

interpretation of differential interactions between

tectonics and eustasy in the North- and South-Lusita-

nian sub-Basins at this time. In such a context, ¯at-

pebble conglomerates were the local marker of these

events in the Sesimbra area.

4. Conclusions

Early lithi®cation of thin carbonate layers and then

reworking under tectonic in¯uence originated ¯at-

pebble conglomerates in a shallow shelf environment

subject to extensional tectonics at Sesimbra (ArraÂbida

sector of the South Lusitanian Basin).

A three-stage genesis for the ¯at-pebble conglom-

erates identi®ed at Sesimbra is proposed (Fig. 14): (1)

differential lithi®cation of thin carbonate and non-

bioturbated layers embedded in a more argillaceous

matrix; (2) breaking of these layers by seismic shocks

associated to extensional faulting and block tilting;

and (3) gravity sliding causing mixing of layer frag-

ments in a matrix.

The horizons with ¯at-pebble conglomerates

resulted from paroxysmal events, which occurred

between phases of lower tectonic activity identi®able

through micro-fracturing and incipient mobilization

of semi-consolidated sediments (plastic deformation).

Hence, the interpretation of the analysed deposits as

¯at-pebble conglomerates based on descriptive

criteria does not contradict their interpretation as seis-

mites (Seilacher, 1969) in genetic terms.

The occurrence of ¯at-pebble conglomerates in the

upper Lower Toarcian of the South Lusitanian Basin

indicates bottom instability in the Sesimbra area during

the Early Jurassic and is one of the rare records of these

peculiardeposits in theMesozoicelsewhere in theworld.

Acknowledgements

Financial assistance was provided by the Centro de

Estratigra®a e Paleobiologia da Universidade Nova de

Lisboa (CEPUNL) and Project MILUPOBAS (Proj.

no PL Ð 930191) co-ordinated by the Gabinete para a

Pesquisa e ExploracËaÄo de PetroÂleo (GPEP), Portugal,


Fig. 14. Proposed mechanism for a three-stage genesis in ¯at-pebble conglomerates studied at Sesimbra.

and the EMMI Group (RNM-178 Junta de AndalucõÂa),

Spain. The authors bene®ted by helpful discussions in

the ®eld with S. Cloething, R. Mouterde, S. Phipps, A.

Ribeiro, A.F. Soares, A.R. Soria and P. Terrinha. The

authors are indebted to insightful comments and criti-

cal review made by J. Menzies and an anonymous

reviewer. We appreciate mathematical assistance

from A.F. Mendes.

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