Sedimentary Geology 139 (2001) 49±70
www.elsevier.nl/locate/sedgeo
Flat-pebble conglomerates: a local marker for Early Jurassic
seismicity related to syn-rift tectonics in the Sesimbra area
(Lusitanian Basin, Portugal)
J.C. Kullberg a,*, F. OloÂriz b,1, B. Marques a,2, P.S. Caetano a,2, R.B. Rocha a,2
a
b
Centro de InvestigacËaÄo em GeocieÃncias Aplicadas da Universidade Nova de Lisboa, Quinta da Torre, 2825-114 Caparica, Portugal
Departamento 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 geodynamic 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* 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.
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 (Hettangian).
(b) The Volcanic±Sedimentary Complex. Probable
0037-0738/01/$ - see front matter q 2001 Elsevier Science B.V. All rights reserved.
PII: S 0037-073 8(00)00160-3
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J.C. Kullberg et al. / Sedimentary Geology 139 (2001) 49±70
Fig. 1. (A) Location of Portuguese Mesozoic Basins. (B) Synthetic geological sketch of the ArraÂbida sector and location of studied outcrops
(CA ˆ Cabo de Ares; CM ˆ Cova da Mijona; S ˆ Sesimbra).
J.C. Kullberg et al. / Sedimentary Geology 139 (2001) 49±70
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 paleogeography 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 subaqueous mass movements in the Upper Liassic (Toarcian) 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 dolomitization precluded precise petrography in the Cabo de
Ares section. The other two sections (Fig. 2), which
were studied by Choffat (1903, 1905), were
51
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 arrondiesº 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 angulososº (hard, grey dolomitic limestone,¼ partly similar to a breccia with angular elements). This
stratigraphic interval (Choffat Bed K) with ªconcreÂtions arrondiesº overlies Bed J labelled by Choffat
(1903, p. 82; 1905, pp. 136±137 ˆ Bed 8) in Zbyszewski 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 identi®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 horizons 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 Hildoceras 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
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J.C. Kullberg et al. / Sedimentary Geology 139 (2001) 49±70
Fig. 2. Studied sections. Lithologic columns showing selected stratigraphic features. Encircled numbers refer to the location of respective
®gures.
J.C. Kullberg et al. / Sedimentary Geology 139 (2001) 49±70
53
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.
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
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J.C. Kullberg et al. / Sedimentary Geology 139 (2001) 49±70
Fig. 3. (continued)
section analysis dif®cult, recognition of six marine
facies (A±F) was possible.
Type-facies A (Fig. 3A) Ð Microsparitic wackestones and packstones. Bioclasts are mainly bivalves
such as pectinids (Chlamys), cardiids (Cardium) and
oysters. Gastropods, belemnites, brachiopods and
echinoderms are rare. Ostracods and benthic foraminifera 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 structures. Bed thickness varies between 15 and 70 cm
(40 cm in average).
Type-facies B (Fig. 3B) Ð Wackestones and packstones similar to those of Type-facies A, but skeletal
debris are mainly made up of bivalves, foraminifera
and algae (Dasycladaceae and Gymnocodiaceae)
J.C. Kullberg et al. / Sedimentary Geology 139 (2001) 49±70
55
Fig. 3. (continued)
accompanied by belemnites, brachiopods, echinoderms, 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. Discontinuity surfaces, some of them ferruginized, exist.
Bioturbation similar to that reported for Type-facies
A is common and locally pervasive. Abundant laminations 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 represented by well sorted, ®ne to medium grainstones,
which are more or less oolitic and bioclastic. Skeletals
are very abundant and diversi®ed: bivalves, gastropods, brachiopods, echinoderm plates and spines,
annelids (serpulids), benthic foraminifera, ostracods,
and algae. Ooids have nuclei of foraminifera, ostracods, 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 subvertical 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 argillaceous 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 thickness 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-skeletal components consist of silt-sized pellets. Laminations 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
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J.C. Kullberg et al. / Sedimentary Geology 139 (2001) 49±70
average) at the Sesimbra section. Type-facies E is the
only facies showing ¯at pebble conglomerates (see
below).
Around twenty depositional units have been recognized, 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 packstones, grainstones and wackestones. Bioclasts are
bivalves, gastropods and brachiopods. Peloids and
ooids are also common. Cross, planar, and wavy laminae 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 intraformational 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 thickness. Thickness and angular to subrounded shape
suggest that pebbles originated through reworking of
consolidated to semi-consolidated thin-bedded calcareous 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 underlying 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, alternatively, 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 slumpfolds forming slabs accompanied by isolated limestone blocks embedded in the intra-formational
conglomerate, which is mainly composed of ¯at
pebbles and sparse skeletals (Fig. 5). Layers containing 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-formational 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 breccias is calcarenitic and mainly made up of organic
detritus that contains fossils bearing traces of
mechanical wear reworked from shallower environments. 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
J.C. Kullberg et al. / Sedimentary Geology 139 (2001) 49±70
57
Fig. 4. Flat-pebbles arrangement parallel to bedding (A) and random (B). Coin size in 4A is 3 cm.
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 lamination 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 movements.
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 according with the style and timing of basin structuring.
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J.C. Kullberg et al. / Sedimentary Geology 139 (2001) 49±70
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.
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 deposition on the basin margins. Initial deposition in carbonate-ramp conditions corresponded to the lower part
of the Achada Dolomites and Limestones (Sinemurian±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, foraminifera), algae and dominant molluscs, are characteristic (Type-facies A). Bivalves such as pectinids
(Chlamys), cardiids (Cardium) and oysters indicate
warm and nutrient-rich shallow waters in moderateenergy coastal areas showing local cohesive
substrates. The presence of rare belemnites, echinoderms and brachiopods indicates that lagoonal conditions 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 ferruginized surfaces (minor discontinuities), and common
peloids indicate episodes of decreasing energy
between minor depositional events unaffected by
burrowers. Thus, Type-facies B deposits are interpreted 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 brachiopods, belemnites and common bivalve fragments
(particularly oysters and Lithophaga) as well as serpulids. Gastropods, bryozoans and foraminifera also
occur.
Non-skeletal elements are intraclasts, peloids and
aggregate grains. Cephalopods (belemnites), brachiopods and crinoids (non-reefal elements) probably
derived post-mortem from somewhat deeper or more
open-sea settings. Further components, like gastropods and bivalves, seem to be admixed from a different 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
J.C. Kullberg et al. / Sedimentary Geology 139 (2001) 49±70
59
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.
but relatively high energy in the sub-tidal zone. Bioclastic 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 fragmentation and terrigenous (quartz) input are recorded. The
oolitic bioclastic packstone/grainstone complex shows
low-angle cross lamination and rapid wedging of individual laminae-sets, the better-sorted sediments probably representing bar-systems within this lagoon
(Type-facies C). These features are in accordance with
lateral accretion deposits, and could relate the cortoidoolitic packstones and grainstones with axial channel
deposits within shallow lagoonal environments.
However, outcrop limitations make dif®cult any conclusive interpretation.
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.
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J.C. Kullberg et al. / Sedimentary Geology 139 (2001) 49±70
Fig. 8. Filling of extensional fracture with mud and sparse ¯at pebbles (arrow) showing evidence of synsedimentary faulting.
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 (belemnites 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 Typefacies 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 deformation. Outcrop limitations make dif®cult the precise
interpretation of the lowermost E-1 stratigraphic interval, but similarity with E-3 is envisaged from local
observations. Therefore, the roughly symmetrical
pattern accords with a paroxysmal tectonic disturbance (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 showing 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. 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.
J.C. Kullberg et al. / Sedimentary Geology 139 (2001) 49±70
61
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J.C. Kullberg et al. / Sedimentary Geology 139 (2001) 49±70
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. Sometimes differential deformation within beds can be
observed: centimetre-order displacements in the lower
parts due to brittle deformation, while ductile deformation 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, vertical positions included. In®lling of some fault-related
pockets includes ¯at pebbles (Fig. 8). On the whole,
fault orientation is rather consistent with an approximately E±W extensional regime (Fig. 9B).
2.4.2. Tension gashes
Tension gashes in ¯at pebbles are always subperpendicular to their larger axes, irrespective of
pebble orientations (Fig. 10). Evidently, these fractures already existed before the reworking events
that caused pebble accumulation (conglomerates).
Furthermore, lithi®cation of the thin carbonate horizons 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 sediments 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 movements 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-fractures that show noticeable refraction in more micritic
levels, were more penetrative in argillaceous fractions, resulting in polygonal frames in the whole set
of affected horizons.
Two processes are envisaged to result in microfracturing: (a) the displacement of semi-lithi®ed sediment; 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. Alternatively, carbonate deposits disrupted in place
(process b) and originated two linear and approximately symmetric patterns showing low angles with
respect to bedding (Fig. 11).
As could be expected (Fig. 9D), the general orientation of resulting structures is very similar to that
found in the extensional faults, their genetic association 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-fracturing 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
J.C. Kullberg et al. / Sedimentary Geology 139 (2001) 49±70
63
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.
suggesting a relatively short displacement (a few
meters at the most), favoured by its arched front.
The occurrence of olistoliths can be related to setbeds collapse and then sliding along fault scarps.
Although the majority of faults affected only one
bed, and rarely two or more, with only centimetreto 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 controlling the structure of the Lusitanian Basin (Ribeiro et
al., 1979; Soares et al., 1993), thus favouring the interpretation of an eastwards regional tilting of blocks.
3. Discussion
During the Mesozoic, the Iberian Peninsula was
located in a particular hinge situation between developing 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 Berlengas 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 structuring. During Mesozoic distension, these faults reactivated as approximately E±W extensional trends
(Ribeiro et al., 1990; Kullberg and Rocha, 1991).
The area studied, near the eastern margin of the Lusitanian 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 indicating 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
64
J.C. Kullberg et al. / Sedimentary Geology 139 (2001) 49±70
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.
movements in¯uencing deposition in the area. Montenat et al. (1988) considered that halokinetic phenomena 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 Sesimbra 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 Cretaceous 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 ¯atpebble production.
Flat-pebble conglomerates are rather particular
Fig. 12. Large olistolith exhibiting type-b micro fracturing and forcing intense deformation on sliding surface.
J.C. Kullberg et al. / Sedimentary Geology 139 (2001) 49±70
Fig. 13. Top surface of a bed showing small olistolith (1), with ªtailº
(2) indicating the direction of movement.
deposits that have been described from various sedimentary 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 subsequent immersions, storm action included (Shinn,
1983). Many authors considered that intra-formational ¯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 cementation 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
65
found in the Sesimbra area were described by
Szulczewski (1968) from the Upper Devonian limestones of the Holy Cross Mts. in Poland. These deposits were originated by sub-aqueous sliding, according
to this author, who considered them, dynamically, as
slide conglomerates that were involved in mass movements 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¯uence 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 conglomerates 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 immersion; (b) storms; (c) tsunami events; and (d) seismicshocks.
(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 Sesimbra 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 deposits. 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 sedimentary structures like hummocky cross-strati®cation
and grain sorting, as well as shelly beds, which are
considered typical components in shallow tempestites
66
J.C. Kullberg et al. / Sedimentary Geology 139 (2001) 49±70
Table 1
Comparison between tsunami and earthquake magnitudes and the corresponding wave run-up elevations
ML ˆ Murty±Loomis tsunami magnitude scale; Iida TM ˆ 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 ˆ 2:66 1 1:66 log ‰…2†=1023 Š:
Values in parenthesis are below the minimum value in Iida's classi®cation, but are considered for discussion purposes; (6) h ˆ 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
ML
2
Tsunami energy (ergs)
3
Earthquake energy (ergs)
4
Earthquake
magnitude
(moment
magn. MW)
5
Iida TM (m)
6
Maximum run-up (m)
10
8
6
4
2
0
24
10 24
10 23
10 22
10 21
10 20
10 19
10 17
10 25 ±10 26
10 24 ±10 25
10 23 ±10 24
10 22 ±10 23
10 21 ±10 22
10 20 ±10 21
10 18 ±10 19
8.8±9.5
8.1±8.8
7.5±8.1
6.8±7.5
6.1±6.8
5.5±6.1
4.1±4.8
4.3
2.7
1.0
20.7
(22.3)
(24.0)
(27.3)
. 20
6.3±8.0
2.0±3.0
0.63±0.75
0.20±0.30
0.06±0.30
0.01±0.30
(Aigner and Reineck, 1982; Duringer, 1984), points to
factors other than storms for ¯at-pebble production in
the Sesimbra area. In addition, the envisaged shallowmarine 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 attributed to a combined origin, since the generating landslides 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
Tidal wave run-ups
events that eventually would produce ¯at-pebble
conglomerates in the Sesimbra area. However, geodynamic and sedimentological data make the application
of this hypothesis dif®cult (see below). The overwhelming majority of historic tsunami, particularly
those with large run-up recorded in the Paci®c
Ocean are due to high-magnitude earthquakes associated 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 MLmagnitude 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) MLmagnitude 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 corresponds to a ªsigni®cant tsunamiº that correlates with
magnitude 1 in Iida's classi®cation, that shows the
J.C. Kullberg et al. / Sedimentary Geology 139 (2001) 49±70
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 unlikely 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 sedimentary record; (2) near-absence of other tectonic disturbances, implying slow and progressive distension; and
(3) facies attributable to shallow lagoonal environments, 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 earthquake epicentre, as Iida (1963) considered: larger
tsunami are generated in deeper waters. Any large
tsunami causing ¯at-pebble conglomerates at Sesimbra would have had a distant origin and a wide regional impact, the record of which is unknown from
contemporaneous deposits elsewhere in the Lusitanian 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). Therefore, 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,
67
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 potentially 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 Norwegian 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 (horizons 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 deposition 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 interpret the studied deposits at Sesimbra as the local
68
J.C. Kullberg et al. / Sedimentary Geology 139 (2001) 49±70
Fig. 14. Proposed mechanism for a three-stage genesis in ¯at-pebble conglomerates studied at Sesimbra.
variant of tectonically induced sedimentary changes
affecting the Lusitanian Basin in different palaeogeographic 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 documented 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-Lusitanian sub-Basins at this time. In such a context, ¯atpebble 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 ¯atpebble 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 conglomerates identi®ed at Sesimbra is proposed (Fig. 14): (1)
differential lithi®cation of thin carbonate and nonbioturbated 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 fragments 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 seismites (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
peculiar deposits in the Mesozoic elsewhere in the world.
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,
J.C. Kullberg et al. / Sedimentary Geology 139 (2001) 49±70
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 critical review made by J. Menzies and an anonymous
reviewer. We appreciate mathematical assistance
from A.F. Mendes.
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