Download The Exposed Passive Margin of North America in Western Cuba

Chapter 4
The Exposed Passive Margin of North America in Western Cuba
ANDRZEJ PSZCZOLKOWSKI
The Mesozoic successions of western Cuba, now exposed in the Guaniguanico terrane, were deposited to the east of the
present NE Yucatan coast. The evolution of these passive margin successions encompasses the syn-rift stage (Early Jurassic?Callovian/early Oxfordian), drift stage (?Callovian/middle Oxfordian-Santonian), and the beginning of the active margin stage
(Campanian-Paleocene). Prior to the middle Oxfordian, the San Cayetano basin was located in an originally narrow rift zone
formed between Yucatan and South America. The onset of shallow-water carbonate sedimentation in the Sierra de los Organos
and Cangre belts occurred in the late Oxfordian or earliest Kimmeridgian. Drowning of a carbonate bank, or platform, in
the early Tithonian resulted in a considerable uniformity of facies in all belts of the Guaniguanico terrane, expressed by
widespread occurrence of ammonite-bearing limestones and radiolarian microfacies, especially in the upper Tithonian deposits.
Pelagic limestones accumulated during the Berriasian and Valanginian, while siliciturbidites occurred in the Northern Rosario,
La Esperanza and Placetas belts of western and central Cuba during the Valanginian-Barremian. These belts belonged to a
deep-water sector of the basin that extended between the Yucatan and Bahamas platforms. During the Aptian-Albian, siliceous
deposition extended across the entire deeper part of the northwestern proto-Caribbean basin. Pelagic carbonate sedimentation
resumed in the Cenomanian. Origin of the regional late Turonian (or Coniacian)-Santonian hiatus in the deep-water, pelagic
sequence of the northwestern proto-Caribbean basin was probably related to paleoceanographic conditions that existed during Late
Cretaceous times. These conditions were associated with paleogeographic changes in the southern part of the proto-Caribbean
basin, when the Nicaraguan Rise-Greater Antilles Arc partially closed the connection with the Pacific.
During the Campanian, abundant volcaniclastic detritus appeared in the upper Moreno Formation of the Northern Rosario belt.
The Bahfa Honda segment of the volcanic arc was located east of the Yucatan block margin and south of the Moreno depocenter.
This arc could be the westernmost part of the Greater Antilles Arc (GAA). Unlike previous interpretations, at the end of the
Cretaceous a more southerly position of the extinct volcanic arc is inferred from the paleotectonic reconstruction and lithology of
the late Maastrichtian deposits. During the Late Paleocene, clastic deposition occurred in a foreland basin setting, in front of a
thrust belt along the southern side of the remnant proto-Caribbean Sea.
INTRODUCTION
The Jurassic to Paleocene sedimentary successions of the passive margins of North America
are exposed in Cuba (Fig. 1). The Mesozoic platform and/or slope deposits crop out in the northern
part of central Cuba. These deposits, traditionally
linked to the Bahamas platform (Meyerhoff and
Hatten, 1974; Pardo, 1975) occur also in the Matanzas Province (Pszcz6~kowski, 1986b) and in eastern Cuba (Iturralde-Vinent, 1996). In western Cuba
(Fig. 2) the Jurassic to Paleocene rocks occur in
the Guaniguanico tectonostratigraphic unit (terrane).
These rocks are considered to belong originally to
the eastern margin and slope of the Yucatan platform
(Iturralde-Vinent, 1994, 1996). Metamorphic rocks
exposed in the Isla de la Juventud (Isle of Pines)
and in the Sierra de Escambray (Fig. 1) are similar to Mesozoic successions of the Guaniguanico
terrane (Khudoley and Meyerhoff, 1971; Mill~n and
Myczyfiski, 1978). Stratigraphic and lithologic similarities existing between the Guaniguanico, Pinos
and Escambray terranes (Fig. 1) clearly suggest their
paleogeographic proximity prior to the Late Cretaceous and Paleogene tectonic events (Pszcz6ikowski,
1981; Iturralde-Vinent, 1994).
Studies of the Pinar del Rio geology were carried
out by oil companies before 1959, which resulted in
many advances in understanding of stratigraphy and
tectonics of this area (Hatten, 1957, 1967; RigassiStuder, 1963; Meyerhoff, in Khudoley and Meyerhoff, 1971; Pardo, 1975). Mapping and research
carried out in western Cuba during the past 26 years
has resulted in publication of many papers, includ-
Caribbean Basins. Sedimentary Basins of the World, 4 edited by E Mann (Series Editor: K.J. Hsti), pp. 93-121.
9 1999 Elsevier Science B.V., Amsterdam. All rights reserved.
94
A
A. PSZCZ()LKOWSKI
N
Igl
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~
jKIj~_ I i ~
U ~ / ~ , ~_
,../
.A.Np.BAHA.M.AS!:
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~) "~11-1~ .
_
NINIR"q"
HAVANA
Ill,, K'~N e \.d,
f
Camagoey
Province
200 km
-~.~.,
Fig. 1. (A) Map of Cuba showing the location of selected geological structures and deep wells. 1 = terranes of passive margin origin
exposed in western and south-central Cuba: GU -- Guaniguanico (stratigraphic terrane), P = Pinos (metamorphic terrane in the Isla
de la Juventud, or Isle of Pines), E = Escambray (metamorphic terrane in the Sierra de Escambray); 2 - the Placetas and Camajuanf
belts in north-central Cuba (Kimmeridgian?/Tithonian to Maastrichtian slope successions) and the Asunci6n metamorphic massif (AN)
in eastern Cuba; 3 - the Remedios belt (Cretaceous platform succession)" 4 = well location sites (shown as encircled numbers: 1 Martfn Mesa 1 (situated in the Martfn Mesa tectonic window), 2 = Pinar 1, 3 = Guanahacabibes, 4 = Los Arroyos 1), BH = Bahfa
Honda terrane (ophiolite and Cretaceous volcanic arc), Gh = Guanahacabibes Peninsula. (B) Schematic map showing the location of the
main terranes and belts in western and central Cuba and adjacent areas (partly after Rosencrantz, 1996 and Case et al., 1984, 1990): a
= Yucat;in platform; b = Florida and Bahamas platforms; c = Camajuanf and Placetas belts (undivided) in north-central Cuba (CA &
PS)" d -- Yucatfin basin; e = Cayman ridge; f = Camagtiey trench; GU = Guaniguanico terrane in western Cuba; BH -- Bahia Honda
terrane (ophiolite and Cretaceous volcanic arc); P -- Pinos terrane; E = Escambray terrane; RS & CC = Remedios and Cayo Coco belts
(undivided) in north-central Cuba (shallow-water and pelagic carbonates); SGM = southeastern Gulf of Mexico. Arrows indicate relative
movement along faults and barbed continuous lines denote major thrusts.
ing overviews of Cuban geology (Lewis and Draper,
1990; Iturralde-Vinent, 1994, 1996), and geological
maps. The present paper focuses on the evolution of
the Jurassic to Early Paleocene passive margin successions now exposed in the Guaniguanico terrane
of western Cuba, before their Paleogene tectonic
deformation.
TECTONIC SETTING
In this paper, a tectonostratigraphic terrane is defined following the criteria of Howell et al. (1985),
adopted in some recent Caribbean geological studies
(for example, Mann et al., 1991). Also in Cuba some
geological structures have been characterized as 'terranes' (Lewis and Draper, 1990; Pszcz6tkowski,
1990; Piotrowska, 1993; Iturralde-Vinent, 1994,
1996; etc.), although the overall, generally accepted
scheme of Cuban tectonostratigraphic terranes is still
to be achieved. It was not the aim of this paper to
propose such a scheme. Rather, the concept of terranes is merely used herein to explain a Mesozoic
evolution of a passive margin successions exposed in
western Cuba. Two types of terranes may be distinguished in western and south-central Cuba, namely
the stratigraphic and metamorphic terranes. However, in western Cuba both types of terranes are not
completely separated, as the Guaniguanico terrane
includes also metamorphic rocks (the Cangre belt).
The term 'Guaniguanico terrane' was introduced
by Iturralde-Vinent (1994). This author (IturraldeVinent, 1994, 1996) proposed a generalized tectonic
scheme of the Pinar del Rio Province and placed the
Guaniguanico terrane among the southwestern Cuban
terranes (with the Pinos and Escambray terranes). In
his opinion, the Guaniguanico terrane is composed
of five juxtaposed belts: Los Organos, Rosario South,
Rosario North, Quifiones and Felicidades.
The Guaniguanico terrane (Figs. 1 and 2) is located mainly in the Pinar del Rio Province but
its eastern extremity reaches the Havana Province.
The Pinar fault forms the southern boundary of
this tectonostratigraphic unit. The eastern part of
the northern boundary of the Guanigianico terrane is defined by the tectonic contact with the
Bahfa Honda terrane and the Guajaibdn-Sierra Azul
unit (Fig. 2B and C). To the west, the northern
THE EXPOSED PASSIVE MARGIN OF NORTH AMERICA IN WESTERN CUBA
83o30 '
84000 '
I
,co
I
95
0A
TERRAN
83c,00 '
I
o
APS
0 S
LM
22030 '
/Y
o P i n a r del R~o
0
Pinar del R{o
0L
,
25 km
jv /
/
Mantua
(Y
LP
La Co[oma <
N
g
~
/7
-
0
I
84o00'
I
20km
l
83030 .
Fig. 2. Location maps (A, C) and tectonic map (B) of the Guaniguanico terrane in the Pinar del Rfo and Havana provinces of western
Cuba. (A) Location of the area shown in (B). (B) Tectonic map of the Guaniguanico terrane (partly simplified, based mainly on data
taken from: Pszczdtkowski et al., 1975" Pszcz6tkowski, 1977, 1978, 1994b; Piotrowska, 1978; Martfnez and V~izquez, 1987); tectonic
units of the Sierra de los Organos belt: VP = Valle de Pons, I = Infierno, G = Sierra de Guane and Paso Real, V = Vifiales, PG = Pico
Grande, SG = Sierra de la Gtiira, A = Anc6n, APS = Alturas de Pizarras del Sur; CB = metamorphosed tectonic units of the Cangre
belt; tectonic units of the Southern Rosario belt: Z = La Zarza, T = Taco Taco, C = Caimito, CP = Cinco Pesos, LT = Los Tumbos, NP
= Niceto P6rez, M = Mameyal, LB = Los Bermejales, PU = Loma del Puerto, LP = La Paloma, LM = Loma del Muerto, J = gabbro
and serpentinite of the Jagua massif in the southwestern part of the Alturas de Pizarras del Norte; tectonic units of the Northern Rosario
belt: B V = Bel6n Vigoa, NO = Naranjo, D = Dolores, LS = La Serafina, CE = Cangre, CH = Sierra Chiquita, QS = Quifiones; GA
= Guajaib6n-Sierra Azul tectonic unit; N - Q = Neogene and Quaternary deposits south of the Cordillera de Guaniguanico; barbed lines
denote thrusts. (C) Location map of the Guaniguanico terrane belts: SO = Sierra de los Organos belt, CB = Cangre belt, SR = Southern
Rosario belt (in the Sierra del Rosario between Soroa and La Palma, and in the Alturas de Pizarras del Norte between La Palma, Mantua
and Guane), NR = Northern Rosario belt, LE = La Esperanza belt, GA = Guajaib6n-Sierra Azul belt; arrows indicate sense of the
movement along the Pinar fault.
boundary of the Guaniguanico terrane is located
in the southeastern Gulf of Mexico, north of the
Pinar del Rio Province (Fig. 1B). To the southwest,
the Guaniguanico terrane is covered by NeogeneQuaternary deposits. The wells drilled in the Guanahacabibes Peninsula revealed the metamorphosed
rocks of the Guaniguanico terrane (Cangre belt) beneath the Oligocene and Miocene rocks, about 55
km to the southwest of Guane (Fig. 1A). According
to Rosencrantz (1996) the Guaniguanico terrane continues to the southwest as a fault-bounded, wedgeshaped block occurring between the Yucatfin basin
and the Yucatfin borderland (see also Fig. 1B). If this
interpretation is correct, the Guaniguanico terrane
may be about 400 km long.
The thrust nappes of the Guaniguanico terrane
consist of Jurassic to Paleogene rocks (Hatten, 1957;
Pszcz6lkowski, 1971) and were formed during the
Early Eocene (Pszcz6tkowski, 1977, 1994b). The
Eocene tectonic deformation affected north-central
Cuba as well. Rigassi-Studer (1963) distinguished
the Sierra de los Organos and Sierra del Rosario
as two distinct stratigraphic successions. The Meso-
zoic successions of the northern Sierra del Rosario
tectonic units and the La Esperanza belt (Fig. 2)
are equivalents (Pszcz6tkowski, 1982, 1994a; Rodriguez, 1987). Consequently, the Esperanza belt is
considered here as a continuation of the Northern
Rosario belt.
In the present paper, the Guaniguanico terrane
is subdivided into the following tectonostratigraphic
belts (from south to north): Cangre (CB), Sierra de
los Organos (SO), Southern Rosario (SR), Northern
Rosario belt (NR) and the Guajaib6n-Sierra Azul
(GA) (Fig. 2C). The stratigraphic successions of the
Cangre and Sierra de los Organos belts are similar (Pszcz61kowski, 1985); however, the Cangre belt
consists of metamorphic (mainly metasedimentary)
rocks. The metamorphic Cangre belt consists of
three tectonic units: Mestanza, Cerro de Cabras and
Pino Solo (Piotrowska, 1978). These units occur in
the southeastern part of the Guaniguanico terrane,
as a narrow tectonic belt along the Pinar fault. The
Sierra de los Organos belt comprises the tectonic
units of (1) the Mogote zone and (2) Alturas de
Pizarras del Sur (APS in Fig. 2B). In general, the SO
96
and CB represent the Jurassic platform that subsided
in the Tithonian and remained submerged in deepwater pelagic conditions during the Cretaceous and
Paleocene.
The Rosario belts occur in the eastern and northern parts of the Guaniguanico terrane (Fig. 2). The
Southern Rosario belt (SR) extends from Soroa to
Mantua and Guane. The 1:250,000 scale geological
map of Pszcz6lkowski et al. (1975), published as a
part of the geological map of Cuba (Puscharovsky,
1988), and the 1:50,000 scale map (Martfnez and
V~izquez, 1987) revealed that the southern Sierra del
Rosario tectonic units continue to the west, between
La Palma and Mantua (Fig. 2B and C). In fact, a
distinction between the SR and SO Jurassic lithology
is difficult in this area. Near Mantua, the Jurassic
(pre-upper Oxfordian) formations display features
characteristic for the Sierra de los Organos belt; to
the northeast these formations gradually change their
facies features reaching SR characteristics in the Minas de Matahambre-La Palma area. The Northern
Rosario belt (NR) occurs north of the SR (in the
Sierra del Rosario), and has its counterpart in the
Esperanza belt to the west and southwest.
The Jurassic to Paleogene rocks exposed in
the Martin Mesa tectonic window in the Havana
Province (located around the Martin Mesa 1 well,
but too small to be shown in Fig. 1) are similar to
those known in the Guaniguanico terrane. The limestones, shales and sandstones of Early Cretaceous
age drilled there in the Martin Mesa 1 well (Fig. 3)
are equivalents of rocks occurring in the Northern
Rosario belt (Fig. 4). The Jurassic-Cretaceous and
Paleocene rocks of this belt (NR) differ from those
occurring in the SO and SR to the south. In general,
the Rosario belts are interpreted as the continental
margin slope and (partly) adjacent basin floor.
The Guajaib6n-Sierra Azul belt 1 (GA) occurs as
a single (and narrow) tectonic unit exposed between
the Northern Rosario belt and the Bahia Honda
terrane (Fig. 2). The GA consists mainly of the
Cretaceous shallow-water carbonates (Pardo, 1975),
named the Guajaib6n Formation (Herrera, 1961;
Pszcz6lkowski, 1978, 1982). Contrary to earlier interpretations (Pardo, 1975; Pszcz6lkowski, 1982) the
pre-tectonic location of the Guajaib6n-Sierra Azul
belt was probably to the south, but not necessarily atop the deep-water Lower Cretaceous deposits
of the Northern Rosario belt as suggested by Iturralde-Vinent (1994, 1996). Rather, this belt could
be situated somewhere at the Yucat~in block edge.
1The original name of this belt (Cacarajfcara Belt Pardo,
1975) cannot be maintained because of the existence of the
Cacarajfcara Formation established earlier (Hatten, 1957) in
western Cuba. This formation occurs in the Rosario and La
Esperanza belts only.
A. PSZCZOLKOWSKI
Fig. 3. Lithologic column of the Martin Mesa 1 well, located
northeast of Cayajabos, in the Havana Province (data from Segura Soto et al., 1985, simplified): 1 -- pelagic limestones, with
intercalations of sandstone and shale; 2 -- detrital limestones;
3 = sandstones, shales and marls; 4 --- tectonic contacts. The
Lower Cretaceous rocks and Campanian-Maastrichtian deposits
(the Cacarajfcara Formation?) are equivalents of the formations
exposed in the Northern Rosario belt of the Guaniguanico terrane
(see Fig. 2B).
If so, the GA unit may represent a separate terrane
not related to the Guaniguanico tectonostratigraphic
belts.
The Paleogene thrusting changed the original
relative positions of the belts and tectonic units
of the Guaniguanico terrane. Contrasting opinions
have been expressed on the problem of the pre-tectonic restoration of the Guaniguanico belts and units
(Hatten, 1957; Rigassi-Studer, 1963; Pszcz6tkowski,
1978; Mossakovskiy and de Albear, 1979; Piotrowska, 1993; and others). In the present paper,
the author accepts the idea, that during the Paleogene thrusting the relative positions of the belts and
tectonic units were completely reversed (IturraldeVinent, 1994, 1996). According to this interpretation,
the ophiolites and the Cretaceous volcanic arc rocks
of the Bahia Honda composite terrane are the structurally highest belts of the Pinar del Rio Province
(Hatten, 1957; Pszcz6tkowski and de Albear, 1982;
Pszcz6lkowski, 1990; Iturralde-Vinent, 1994).
THE EXPOSED PASSIVE MARGIN OF NORTH AMERICA IN WESTERN CUBA
97
Fig. 4. Generalized lithostratigraphic scheme of the Guaniguanico terrane in western Cuba (for location of belts see Fig. 2C); lithology:
1 -- sandstones and shales with intercalations of limestone, 2 -- mafic rocks, 3 = fossiliferous limestones and shales, 4 ---=thick-bedded
to massive carbonates (Jurassic in age), 5 = thin- and medium-bedded limestones, 6 = limestones, shales and sandstones (Polier
Formation), 7 = radiolarian cherts and shales, 8 = massive, shallow-water limestones (Cretaceous in age), 9 = shales and sandstones
(Late Cretaceous and Paleogene in age), 10 - detrital limestones and breccia (Cacarajfcara Formation), 11 = limestone and chert
breccias (Anc6n Formation), 12 - Paleogene olistostrome; lithostratigraphic units (circled letters): SC = San Cayetano Formation, A C -Arroyo Cangre Formation (equivalent to the San Cayetano Formation), ES = E1 Sfibalo Formation, J Jagua Formation, F = Francisco
Formation, S V = San Vicente Member of the Guasasa Formation, G = Guasasa Formation, AR -- Artemisa Formation, PL -- Polier
Formation, L = Lucas Formation, ST = Santa Teresa Formation, GB -- Guajaib6n Formation, P N = Pons Formation, CT = Carmita
Formation, PA = Pinalilla Formation, MR = Moreno Formation, PS = Pefias Formation, CA = Cacarajfcara Formation, A N -- Anc6n
Formation, M N -- Manacas Formation.
=
The presence of the Upper Jurassic shallow-water
limestones in the Pinar 1 well (Fig. 5) is a strong
argument in favor of the above-mentioned tectonic
restoration. This well was located 4 km south of
Pons in the Sierra de los Organos (L6pez Rivera
et al., 1987; Pszcz6lkowski, 1994), in a tectonic
window occurring in the central, most uplifted zone
of the Guaniguanico terrane. The Valle de Pons unit
(VP in Fig. 2B) is the lowermost tectonic element
exposed in the Sierra de los Organos (Piotrowska,
1978) and probably in the whole Guaniguanico
terrane. At least three tectonic units were drilled in
the Pinar 1 deep well (Fig. 5). The bottom unit,
reaching below 3300 m, contains very thick Upper
Jurassic shallow-water limestones (1500 m). L6pez
Rivera et al. (1987) suggested that these limestones
belong to the autochthonous unit. However, there
is no conclusive evidence that the Pinar 1 well
really penetrated the whole Guaniguanico nappe
pile and entered into the autochthonous sedimentary
succession.
In any case, the Upper Jurassic shallow-water
carbonates attain their m a x i m u m thickness (1500 m)
in the subsurface. These rocks are thinner in the
higher tectonic units of the Sierra de los Organos
belt ( 3 0 0 - 6 0 0 m), and eventually wedge out in the
overlying tectonic units of the Southern Rosario belt
(Pszcz6tkowski, 1978).
OUTLINE OF STRATIGRAPHY
R e m a r k s on lithostratigraphic s c h e m e
The lithostratigraphic scheme for the Guaniguanico terrane, as used in this paper (Figs. 4 and 6), was
developed during the last 18 years (Pszcz6tkowski
et al., 1975; Pszcz6tkowski, 1978, 1982, 1994a).
Other authors (Iturralde-Vinent, 1994, 1996; Cobiella-Reguera, 1996) accepted this scheme, although
with some minor changes. The results of a recent micropaleontological study on the Paleogene deposits
(Bralower et al., 1993; Bralower and Iturralde-Vinent,
1997) also have been considered.
The work of Hatten (1957, 1967) was fundamental for developing the m o d e m stratigraphic
scheme for the Sierra de los Organos belt (see also
Khudoley and Meyerhoff, 1971). Herrera (1961)
98
A. PSZCZOLKOWSKI
PINAR
Sierra de los
Organos belt
1
EARLY
EOCENE
m 0
Southern
Rosario belt
MANACAS FM. ;'i... . . . . . . .
i
Northem
Rosario belt
1
Vieja Member
......................
Pica Pica Member
I
PALEOCENE
1000
CACARAJiCARA FM.
IVlAASTRICHTIAb
Lower Cretaceous (Aptian - Albian)
CAMPANIAN
Lower Cretaceous (Valanginian)
S~-E~-%5]q]-R~--
Lower Cretaceous (Aptian - Albian)
CONIACIAN
TURONIAN
CENOMANIAI~
APTIAN-ALBIAI~
BARREMIAN
"-~,~'~.~'t
Upper Jurassic
CARMITA FM.
FORMATION
VALANGINIAN
~<,.,,<
BERRIASIAN
Upper Jurassic
2000
'< ''< ''<
,~-.~-)
<Z~
0B
! Tumbitas Mb.
~u-r~i~aer~' '
Member
KIMMERIDGIAh
u) ~ : ,
Member
<
p ...............
D ~ i i San Vicente
o o,
Member
i
Lower Eocene
-
I
u_m.i,:r~Mb' :1
~ ARTEMISA
< ~- FE~i~,?rie-a-ca-n~
.......
TITHONIAN
,'-,>,3 ,,-<
SANTA TERESA FM.
PONS
5
La Zarza
Member
q
I
', FORMATION
~;:,
',
....
i
FRANCISCO FM.
Lower Cretaceous (Aptian - Albian)
-L_
FM.
300C
Lower Cretaceous (Berriasian - ?Barremian)
--
Upper Jurassic
Lower Cretaceous (Berriasian - Valanginian)
Lower Cretaceous (Aptian - Albian)
" I
SAN CAYETANO
FORMATION
Lower Cretaceous (Berriasian - ?Barremian)
N M'-N'-M
~ - ~ - ~
mm--
~ - ~ - ~
5
Fig. 6. Lithostratigraphic scheme of the Sierra de los Organos
and Rosario belts in the Guaniguanico terrane: L. = lower (Oxfordian), U. = upper (Oxfordian), S.V. = San Vicente Member of
the Artemisa Formation.
Upper Jurassic
~ - ~
5000
--'
~ ~ ~
Fig. 5. Lithologic column of the Pinar 1 well, located 4 km south
of Pons in the Sierra de los Organos belt (after L6pez Rivera
et al., 1987, simplified): 1 = massive, shallow-water limestones,
2 - Berriasian-?Barremian pelagic limestones, 3 -- AptianAlbian pelagic limestones, 4 -- Lower Eocene olistostrome, 5 -tectonic contacts.
introduced many lithostratigraphic names for the
Sierra de los Organos area, but only few were
valid and have been incorporated into the modern
scheme (Fig. 6). Imlay (1942), de la Torre (1960,
1988), Furrazola-Bermddez (1965), Judoley and Furrazola-Bermddez (1968), Wierzbowski (1976), Myczyfiski (1976, 1977, 1989, 1994a,b), Myczyfiski
and Pszcz6tkowski (1976, 1990, 1994) and Kutek et
al. (1976) studied fossils and/or biostratigraphy of
the Jurassic and Lower Cretaceous formations in the
Sierra de los Organos belt and Rosario belts.
San Cayetano Formation (?Lower Jurassic to
middle Oxfordian)
Lithology and facies model. This formation
consists of shales, siltstones and sandstones with
some intercalations of conglomerates and limestones. Limestones occur mostly in the upper part
of the San Cayetano Formation. The rocks are
dark-gray to black; they are rhythmically bedded
(Fig. 7). The formation is deeply weathered so that
the exposures favorable for sedimentological observations occur mainly in some streams and rivers. The
sedimentary structures vary among distinct tectonic
units (Meyerhoff and Hatten, 1974). Any reliable
lithostratigraphic subdivision of the monotonous and
often severely tectonized formation, 1000 to 3000(?)
m thick, has not been well established so far.
Haczewski (1976) proposed a general descriptive
model of sedimentation of the San Cayetano Formation. This author distinguished nine facies (A-I)
within the studied sections. In his opinion, the facies A - F occur in the Sierra de los Organos belt.
However, the type sections of the facies A - C and
E are now considered to belong to the Southern
Rosario belt (Pszcz61kowski, 1994b). The facies A D consists of sandstones with some subordinate
pebbly sandstones or conglomerates, fine-grained
sandstones with trough cross-lamination, of bedded siltstones, shales, and rarely very fine-grained
sandstones, siltstones, fine-grained sandstones and
thin-bedded shales. The rocks belonging to these
facies have been interpreted as deposited in a fluvial environment, and partly in a shallow marine
or beach environment. The facies E and F consist
of black shales, sometimes with septarian nodules
and pyrite spheres in facies E deposits. Both facies
THE EXPOSED PASSIVE MARGIN OF NORTH AMERICA IN WESTERN CUBA
99
Fig. 7. Shales, siltstones and fine-grained sandstones of the San Cayetano Formation at Cinco Pesos (Southern Rosario belt). These
deposits are similar to facies G and H of Haczewski (1976).
most likely have originated in extensive lagoons with
restricted circulation.
The facies G-I (rhythmic sandstones and shales
with graded and ripple bedding, alternating graded
sandstones and shales, and thick-bedded sandstones)
occur only in the Southern Rosario belt. The facies
G consists of fine-grained sandstones, siltstones and
shales. Deposits similar to facies G and H are exposed in the Cinco Pesos area (Fig. 7). The thick-bedded, coarse-grained sandstones, sometimes with pebbles up to 7 cm long, are also known in the Cinco
Pesos area; the sandstones belong to the facies I of
Haczewski (1976). Fossiliferous pebbles, containing
late Paleozoic foraminifers and bryozoans, have been
found in these sandstones (Pszcz6tkowski, 1989b).
The foraminifers belong to Fusulinacea (Schwagerina sp. and Parafusulina sp.) of Permian age. One
specimen was identified as Tetrataxis sp.
Fossils and age. The age of the San Cayetano
Formation (?Early Jurassic-middle Oxfordian) is
well established in the uppermost part of this unit
only. In the Sierra del Rosario belt, Myczyfiski and
Pszcz6tkowski (1976) have found some ammonites
in the uppermost part of the San Cayetano Formation, southeast of La Palma. These ammonites belong to the following taxa: Perisphinctes (?Dichotomosphinctes) cayetaensis Myczyfiski, 1976, P. (?Dichotomosphinctes) cf. anconensis Sfinchez Roig, and
P. (Discosphinctes) cf. pichardoi Chudoley et Furrazola-Berm6dez. The ammonites indicate that the
uppermost part of the San Cayetano Formation is of
middle Oxfordian age. In the Sierra de los Organos
belt, the siliciclastic deposits of the San Cayetano
contain very scarce macrofossils, mainly bivalves.
Pugaczewska (1978) identified the following taxa:
Eocallista (Hemicorbula) spp., Vaugonia (Vaugonia)
spp., Gervillaria sp., and Neocrassina spp. In the
Sierra de los Organos belt, bivalves belonging to the
genus Gryphaea Lamarck probably also appear in
the coquinid limestones of the upper part of the San
Cayetano Formation.
Facies interpretation. According to Haczewski
(1976), the San Cayetano rocks were deposited on a
coastal alluvial plain (facies A-C) by a river transporting the material a few hundred kilometers from
the south (or southwest). The sediments debouched
to the sea formed an arcuate delta and some of it
was redistributed by a longshore drift and turbidity
currents. Deposits distinguished as facies G probably
accumulated on the slope of the continental margin.
Facies H and I were characterized as deposits of a
submarine fan accumulating at the base of the slope.
The development of facies H is intermediate between normal and proximal turbidites, while facies I
is of a proximal character (Haczewski, 1976).
The relationship between the diagrams of paleocurrent measurements and the geographical distribution of the localities studied by Haczewski (1976)
changes after the restoration of tectonic units. The
resulting paleocurrent pattern indicates that the directions from south or south-southwest predominate
in the northwestern part of the San Cayetano sedimentary basin (APS in the Sierra de los Organos belt
m Fig. 2B), and the directions from northeast prevail
in its southeastern sector (Southern Rosario belt).
El S~ibalo Formation (Oxfordian)
Lithology and boundaries. The E1 Sfibalo
Formation occurs in the Northern Rosario belt
(Pszcz6tkowski, 1994a). This unit, up to 400 m
thick, consists of basalts and diabases with interbed-
100
ded limestones, marls and sometimes shales. These
horizons of pyroclastic rocks also have been observed. The basalts are massive or pillowed flows
(Pszcz6tkowski and de Albear, 1983). The lower
boundary of the E1 S~ibalo Formation is tectonic;
this unit contacts with various Cretaceous and Paleogene formations. Locally, the E1 S~ibalo Formation
is overlain by thin San Cayetano Formation siliciclastics, but often contacts with the ?late OxfordianKimmeridgian limestones of the Artemisa Formation.
The E1S~ibalo/Artemisa boundary is tectonically disturbed in some sections, but in places a thin bed of
sedimentary breccia separates the two formations.
Age. The Jurassic age of the E1 S~ibalo Formation (Oxfordian-?early Kimmeridgian) was defined
on the basis of infrequent microfossils occurring
in the limestones intercalating with the basalts and
diabases (Pszcz6tkowski, 1989a, 1994a). According
to Cobiella-Reguera (1996), these rocks overlie the
San Cayetano Formation clastics and may be as old
as Callovian in the lowermost part of the E1 S~ibalo
Formation. The alleged position of the E1 S~ibalo
Formation rocks above the San Cayetano Formation
is not supported by the geological relations visible in outcrops described so far from the Sierra
del Rosario. In fact, the San Cayetano-type clastic
deposits (up to 15 m thick) were observed in few
places above the rocks of the E1 S~ibalo Formation in
the Northern Rosario belt (Pszcz6tkowski, 1994b).
Therefore, it seems that the E1 S~ibalo Formation is
(locally?) older than the Francisco Formation (latemiddle to late Oxfordian). Consequently, the age of
the E1 S~ibalo Formation should be pre-late Oxfordian, in some sections at least (Fig. 6). The supposed
Callovian age of the lower part of the E1 S~ibalo
Formation is, however, still not confirmed by any
paleontological or radiometric data.
Facies interpretation, The limestones of the E1
S~ibalo Formation have been deposited in anaerobic
conditions, below the wave-base, in the outer neritic
zone. This origin of this formation was related to the
more advanced rifting stage during ?Middle to early
Late Jurassic times. According to Iturralde-Vinent
(1995a), the magmatic activity of continental passive
margin represented in Cuba was time coincident
with continental rifting until the Oxfordian, and with
oceanic spreading in the Caribbean area since the
Late Jurassic.
Jagua Formation (middle and upper Oxfordian)
Subdivision and lithology. The Jagua Formation occurs in the Sierra de los Organos belt. This
formation consists of limestones and shales subdivided into four subunits, namely the Pan de Azfcar,
Zacarfas, Jagua Vieja and Pimienta members. The
Pan de Azficar Member (Fig. 6) consists of black,
A. PSZCZOLKOWSKI
well-bedded coquinas and bioclastic limestones. The
beds and lenses of silicified limestones are common. The coquinas beds are 0.2 to 1.5 m thick. The
bivalve-echinoderm and bivalve-oolitic microfacies
are very common in the Pan de Azfcar Member. The
coquina beds consist mainly of Gryphaea mexicana
Felix (Pugaczewska, 1978). The bivalve shells are
commonly encrusted by agglutinated foraminifers.
The limestones of this member were deposited at
a depth ranging from a few up to some tens of
meters, in proximity of some oolitic shoals. The
Zacarfas Member, not shown in Fig. 6, consists of
ammonite-beating shales with thin intercalations of
siltstones and bivalve coquinas with Liostrea, Ostrea, Exogyra and Plicatula (Wierzbowski, 1976).
This member attains 40 m in thickness and is of
middle Oxfordian age (op. cit.). The Jagua Vieja
Member comprises black shales and marly limestones, up to 60 m thick. The calcareous concretions
contain numerous ammonites, fish remains, marine
reptile bones (Iturralde-Vinent and Norell, 1996) and
bivalves. The Pimienta Member occurs in the upper
part of the Jagua Formation (Fig. 6). This unit, up
to 60 m thick, consists of well-bedded, dark-gray
to black micritic limestones with some minor shale
intercalations in the lower part of the member.
Age. The ammonites indicate that the Zacarras Member represents the middle Oxfordian
(Wierzbowski, 1976). As judged from the relation
to the Zacarfas and Jagua Vieja members, the Pan
de Azt~car Member is probably middle Oxfordian
in age. The Jagua Vieja Member has been assigned
to the middle Oxfordian on the basis of well preserved ammonites (Wierzbowski, 1976). Myczyfiski
(1976) described from the Pimienta Member several
ammonite species assigned to Mirosphinctes and
Cubaspidoceras. One specimen of Taramelliceras
(Metahaploceras) sp. also has been found. This
ammonite assemblage was assigned to the upper
Oxfordian (Myczyfiski, 1994a).
Facies interpretation, The bioclastic limestones
and coquinas of the Pan de Azfcar Member were
deposited at a paleodepth ranging from a few to
some tens of meters, in the proximity of some
oolitic shoals. The shales of the Zacarfas Member
were accumulated in a deeper part of the shelf
environment adjacent to a delta. The shales and
limestones of the Jagua Vieja Member and the
limestones of the Pimienta Member were deposited
in an outer shelf environment.
Francisco Formation (middle to upper
Oxfordian)
Lithology. In the Sierra del Rosario, the Francisco Formation is an equivalent of the Jagua Formation. The Francisco Formation consists of shales, mi-
THE EXPOSED PASSIVE MARGIN OF NORTH AMERICA IN WESTERN CUBA
Fig. 8. Spicules of Didemnidae (Didemnum carpaticum Migik
et Borza and Didemnum sp.) in a microsparitic limestone of
the Francisco Formation (middle and upper Oxfordian), Southern
Rosario belt; • 220.
critic limestones, and thin sandstone intercalations.
Calcareous concretions occur in shales, sometimes
with bivalves and/or ammonites. At Cinco Pesos,
a volcanic rock (basalt) half a meter thick occurs
within limestones and sandstones. The formation
attains 25 m in thickness.
Age. The Francisco Formation contains some ammonites, rare bivalves, fish and plant remains. Some
limestone beds contain Globochaete alpina Lombard
and spicules of Didemnidae (Fig. 8), while radiolarians were not observed in thin sections. The ammonites indicate the upper part of the middle Oxfordian and also the lower part of the upper Oxfordian
(Kutek et al., 1976; Myczyfiski, 1976, 1994a).
Facies interpretation. The deposits of the Francisco Formation were accumulated in an outer
shelf environment. This environment was probably slightly deeper than in the case of the Jagua
Formation deposits.
Guasasa Formation (?upper OxfordianValanginian)
Subdivision and lithology. The Guasasa Formation has been distinguished in the Sierra de los
Organos belt. This formation is subdivided into four
subunits, namely the San Vicente, E1 Americano,
Tumbadero and Tumbitas members (Fig. 6).
The San Vicente Member, 300-650 m thick, occurs in the lower part of the Guasasa Formation. This
member consists of massive or thick-bedded, gray
to black limestones, often dolomitized, sometimes
with chert nodules and lenses. Micritic limestones
with Favreina predominate in the lower part of the
member, while oncolitic and algal calcarenites occur in the upper part. At the top, there are also
well-bedded limestones up to a dozen meters thick.
The San Vicente Member includes a sedimentary
101
limestone breccia (sharpstone) separating the massive limestones of the Guasasa Formation from the
Jagua Formation (Hatten, 1957). A description of
the microfacies of San Vicente Member was given
by Pszcz6lkowski (1978).
The E1 Americano Member comprises well-bedded, dark-gray to black limestones, up to 45 m thick.
The Saccocoma-Didemnidae microfacies is characteristic for the limestones of the lower part of this
member, while biomicrites with calpionellids and
radiolarians occur in its upper part (Myczyfiski and
Pszcz6lkowski, 1990). The E1 Americano Member
terminates the Jurassic part of the Guasasa Formation (Fig. 6). The Tumbadero Member consists
of well-bedded, often laminated, radiolarian biomicrites and calcilutites with black chert intercalations.
The thickness of this unit ranges from 20 to 50
m. The Tumbitas Member consists of thick-bedded,
light-gray calpionellid, calpionellid-radiolarian and
nannoconid biomicrites with infrequent thin intercalations of dark-gray or reddish limestones. The
thickness of these deposits is about 40 m, but in
some sections attains 80 m.
Age. The ?late Oxfordian-Kimmeridgian age of
the San Vicente Member is defined mainly on the
basis of the lithostratigraphic position of this unit
(Fig. 6).
The late Oxfordian age of the basal San Vicente
Member limestones is suggested by the ammonitebeating Jagua Formation occurring below the massive limestones of this unit (Wierzbowski, 1976;
Myczyfiski, 1976). However, the limestone breccia
locally separating the San Vicente Member from the
Pimienta Member of the Jagua Formation indicates
that upper Oxfordian limestones had been partly
eroded before the deposition of the shallow-water
San Vicente carbonates. A stratigraphic hiatus of
unknown duration could be related with this (latest
Oxfordian?) erosive event. The Kimmeridgian age of
the San Vicente Member carbonates was confirmed
by some microfossils (Fermindez Carmona, 1989).
The E1 Americano Member of the Guasasa Formation is Tithonian in age. Ammonites collected at
the base of the E1 Americano Member are early
Tithonian in age (Houga, 1974; Myczyfiski, 1989).
Recently, few specimens belonging to the genus
Hybonoticeras have been identified in the Sierra
de los Organos belt (Myczyfiski, 1996a). These
ammonites indicate the Hybonoticeras-Mazapilites
Zone in the Sierra de los Organos belt (Fig. 9).
Their presence suggests that the age of the upper
boundary of the San Vicente Member is close to
the Kimmeridgian-Tithonian boundary. The upper
Tithonian ammonite assemblage (Myczyfiski, 1989,
1994a, 1996a,b; Myczyfiski and Pszcz6~kowski,
1990) contains some cosmopolitan taxa (Durangites, Corongoceras, Kossmatia), those known from
102
A. PSzCZOLKOWSKI
LU
!-O9
z
<
Substage
Upper
z
0
-1!-
~_
Ammonite zones
Durangites - Himalayites Hildoglochiceras (Salinites)
Paralytohoplites carribeanus
Pseudolissoceras spp.
Lower
Hybonoticeras - Mazapilites
Fig. 9. Ammonite zones established in the Tithonian limestones
of the E1 Americano Member (the Guasasa Formation), eastern
part of the Sierra de los Organos belt (Myczyfiski, 1996a,b and
pers. commun., 1996).
Mexico (Hildoglochiceras (Salinites), Proniceras),
as well as endemic taxa (Vinalesites rosariensis
(Imlay), Protancyloceras hondense Imlay, Butticeras
butti (Imlay) and Butticeras antilleanum (Imlay)).
Also bivalves Buchia aft. B. okensis (Pavlov) and
Buchia aft. B. piochii (Gabb) have been reported
from the upper Tithonian limestones of the Sierra de
los Organos belt (Myczyfiski, 1989).
The ammonites are rare and poorly preserved in
the deposits of the Tumbadero and Tumbitas members of the Guasasa Formation. Therefore, stratigraphy of these members is based on calpionellids.
The Tumbadero Member is Berriasian in age, while
the Tumbitas Member is of late Berriasian-early
Valanginian age (Pszcz61kowski, 1978).
Facies interpretation. The ?late OxfordianKimmeridgian rocks of the San Vicente Member
were deposited on the shallow-water carbonate bank
(Pszcz6ikowski, 1978, 1981) or platform. The Lower
Tithonian S a c c o c o m a - D i d e m n i d a e biocalcisiltites of
the E1 Americano Member accumulated in a deeper
(outer neritic) environment. The upper Tithonian
biomicrites of this member were deposited in an
outer neritic to bathyal environment. The radiolarian
and/or calpionellid biomicrites of the Tumbadero
and Tumbitas members are bathyal deposits and
were laid down below the aragonite compensation
depth.
Artemisa Formation (upper OxfordianValanginian)
Subdivision and lithology. The Artemisa Formation is subdivided into three members: San Vicente,
La Zarza and Sumidero (Fig. 6). The San Vicente
Member was recognized but in a few sections of the
Southern Rosario belt. In the Northern Rosario belt,
the thick-bedded dolomitized limestones occurring
in the lower part of the Artemisa Formation are
lithologic equivalent to the San Vicente Member.
The La Zarza Member consists of bedded (0.1 to
0.8 m), gray to black micritic limestones (Figs. 10
and 11) with some intercalations of shale, siltstone
and fine-grained sandstone in the lower part of this
unit. The limestone beds contain rare aptychi and
fish remains. In the upper part of the member,
there are calcilutites interbedded with dark-gray to
black bioclastic limestones and coquinas composed
of ammonite shells and aptychi. The thickness of
the La Zarza Member attains 200 m. The limestones
of this member are overlaid by light-brown, rose
and gray to black biomicrites with intercalations of
radiolarian chert, assigned to the Sumidero Member (Pszcz6tkowski, 1978). The limestones contain
Fig. 10. Kimmeridgian limestones of the La Zarza Member (Artemisa Formation) exposed in a quarry west of La Palma, Southern
Rosario belt.
THE EXPOSED PASSIVE MARGIN OF NORTH AMERICA IN WESTERN CUBA
Fig. 11. Black Tithonian limestones of the La Zarza Member,
Artemisa Formation, exposed in a road-cut west of Cinco Pesos,
Southern Rosario belt.
abundant radiolarians. Calpionellids are common in
the lower part of the member, while numerous aptychi occur in its upper part. The Sumidero Member
attains 200 m in thickness in some sections.
Fossils and age. The Cubaspidoceras-Mirosphinctes assemblage has been identified in the
basal limestone beds of the Artemisa Formation
(Kutek et al., 1976). These ammonites indicate
the late Oxfordian age of these deposits. Consequently, the lower part of the La Zarza Member is late Oxfordian-Kimmeridgian in age. The
Tithonian ammonite assemblage occurs in the upper part of the La Zarza Member. These ammonites were studied by Imlay (1942), Judoley
and Furrazola-Bermfidez (1968), Houga (1974), and
Myczyfiski (1989, 1994b). The taxa Paralytohoplites, Butticeras, Vinalesites, Protancyloceras and
Hildoglochiceras (Salinites) are characteristic for
the Tithonian of the Rosario belts (Myczyfiski and
Pszczdtkowski, 1994; Myczyfiski, 1996a,b). The bivalves occurring in the Tithonian limestones of
the Rosario belts belong mainly to the genus Inoceramus (Parkinson, 1819), although representatives of Buchiidae are also present (Myczyfiski and
Pszczdtkowski, 1994; Myczyfiski, 1994b).
The Saccocoma-Didemnidae microfacies predominates in the lower Tithonian limestones of the
Southern Rosario belt, being replaced gradually by
the radiolarian microfacies in the Chitinoidella spp.
Zone and lower part of the Crassicollaria Zone
(Myczyfiski and Pszcz6tkowski, 1994).
103
The calpionellids document a late Berriasianearly Valanginian age of the lower and middle parts
of the Sumidero Member. Ammonites attributed to
Thurmanniceras cf. novihispanicus (Imlay) indicate
a Valanginian age (Myczyfiski, 1977) and nannofossils (Nannoconus spp.) suggest a late Valanginian
age for the upper part of the Sumidero Member.
The mentioning of the presence of Southern
Boreal/Northern Tethyan faunas (30~
in the
Valanginian deposits of the Sierra de los Organos
and Sierra del Rosario (Pessagno et al. in Chapter
5) needs a comment. Two specimens of Buchia sp.
figured by Myczyfiski (1977, pl. 8: 6-7) were collected from the Artemisa Formation at La Catalina.
The Tithonian limestones are exposed at this locality
(Myczyfiski, 1989; Myczyfiski and Pszcz6tkowski,
1994), but no Cretaceous rocks with macrofauna are
known there. Therefore, these specimens of Buchia
sp. are late Tithonian (or earliest Berriasian?) in
age. The specimens of Vinalesites rosariensis (Imlay) figured by Myczyfiski (1977, pl. 3: 1), were
also found in the Artemisa Formation at La Catalina
and are late Tithonian in age. In fact, specimens
of Buchia from the (stratigraphically well documented) Lower Cretaceous deposits of the Sierra del
Rosario and Sierra de los Organos, were not figured in any paper. In this situation, the presence of
a Southern Boreal/Northern Tethyan macrofauna in
the Valanginian deposits of the Guaniguanico terrane
needs confirmation or should be rejected.
Facies interpretation. The deposits of the lower
part of the La Zarza Member accumulated in relatively quiet conditions of a partly restricted (lagunal?) environment. The limestones of the upper
part of the La Zarza Member were deposited in the
outer neritic environment, as indicated by its faunal
content. Pelagic deposits of the Sumidero Member
accumulated in bathyal zones below the aragonite
compensation depth, as evidenced by the relative
abundance of aptychi and scarcity of ammonites in
the radiolarian limestones.
Polier Formation (upper Berriasian-?Albian)
Lithology. This formation consists of gray
pelagic limestones with intercalations of turbiditic
sandstones and shales (Fig. 12). Radiolarian or
radiolarian-spicule microfacies are typical of the
limestones. Sometimes the sandstones and shales
may be distinguished as a distinct subunit in the
topmost part of the formation (the Roble Member).
The Polier Formation attains its maximum thickness
(about 300 m) in the Cangre tectonic unit (CE in
Fig. 13). The Polier Formation occurs in the Northern Rosario belt. Similar rocks of Early Cretaceous
age are also known in the Esperanza belt (Fig. 2).
Age. Calpionellids indicate the late Berriasian age
104
A. P S Z C Z O L K O W S K I
Fig. 12. Limestones, sandstones and shales of the Polier Formation (Lower Cretaceous), exposed west of Cayajabos, Northern Rosario
belt; hammer 28 cm long.
Fig. 13. Palinspastic restoration of relative position and width of the major tectonic units occurring in the Guaniguanico terrane (Fig.
2B). The highest units of the Northern Rosario belt were originally located far to the southeast (Iturralde-Vinent, 1994; Pszcz6tkowski,
1994b), while the structurally lowermost units exposed in the Sierra de los Organos belt (VP, /) are shown at the opposite, northwest
extreme of this pre-tectonic reconstruction. Tectonic units: VP = Valle de Pons, I = Infierno, V = Vifiales, A = Anc6n, PG & SG =
Pico Grande and Sierra de la Gtiira, A P S = Alturas de Pizarras del Sur, CB -- Cangre belt (Cerro de Cabras, Mestanza and Pino Solo
metamorphosed units), C + T = Caimito and Taco Taco, Z = La Zarza, A P N = Alturas de Pizarras del Norte (Loma del Muerto and
La Paloma units), CP -- Cinco Pesos, MA = Mameyal, B V = Bel6n Vigoa, NO -- Naranjo, CE = Cangre, CH = Sierra Chiquita, QS =
Quifiones. Lithology: 1 = sandstones and shales, 2 -- mafic rocks, 3 = fossiliferous limestones and shales, 4 -- massive shallow-water
limestones, 5 = thin- to medium-bedded limestones, 6 = limestones, shales and sandstones (Polier Formation), 7 = radiolarian cherts and
shales, 8 = sandstones and shales (Late Cretaceous and Paleogene in age), 9 -- detrital limestones and breccia (Cacarajfcara Formation),
10 = limestone and chert breccias (Anc6n Formation), 11 = Paleogene olistostrome; lithostratigraphic units (encircled letters) as in
Fig. 4.
THE EXPOSED PASSIVE MARGIN OF NORTH AMERICA IN WESTERN CUBA
of the basal beds of the Polier Formation. Ammonite
imprints are common in the middle part of this
formation. Valanginian to ?Albian taxa have been
identified by Myczyfiski (1977). However, the bulk
of the deposits is Hauterivian-Barremian in age.
Rhyncholites (Hou~a, 1969) and aptychi are common in the Polier Formation. The lower boundary of
the Polier Formation is diachronous.
Facies interpretation. The turbidites of siliciclastic and mixed (siliciclastic and calcarenitic) composition are the most characteristic constituent of the
Polier Formation. The sandstone beds are 0.02-1.0
m thick and most of them are graded, with horizontal
and cross-lamination in their upper part (the Bouma
sequence). Flute, groove and prod casts are very frequent on the soles of sandstones. Except some thick
sandstone beds of the topmost part of the formation,
these deposits display features of distal turbidites.
Paleocurrent measurements indicate paleoflow toward the south and southeast (Pszcz6tkowski, 1982).
The pelagic limestones and turbidites accumulated
in the deep-water part of the basin, above the calcite
compensation depth.
Lucas Formation (upper HauterivianBarremian)
Lithology. The Lucas Formation consists of
thin-bedded black limestones with occasional intercalations of fine-grained sandstones and shales. The
limestones are radiolarian biomicrites. The thickness
of this formation attains 300 m.
Age. The Lucas Formation is the stratigraphic
equivalent of the Polier Formation (Fig. 6) in the
Quifiones tectonic unit (Fig. 13). The aptychi and
ammonite imprints are the only macrofossils collected from the Lucas Formation. The ammonites
indicate the late Hauterivian to early Barremian age
of this formation (Myczyfiski, 1977). The top of
the Polier and Lucas formations marks the upper
boundary of the Upper Jurassic-Lower Cretaceous
limestone sequence in the Northern Rosario belt.
Both formations occur directly below the radiolarian
cherts and shales of the Santa Teresa Formation.
In the NR this important facies change occurred
significantly later than in the SR (Figs. 4 and 6).
Facies interpretation. The ammonite-bearing radiolarian limestones of the Lucas Formation are interpreted as pelagic deposit accumulated in bathyal
environment, less influenced by influx of terrigenous
sediment.
Santa Teresa Formation (?HauterivianCenomanian)
Lithology and distribution. The Santa Teresa
Formation is composed of green, red or reddish-
105
brown and sometimes black radiolarian cherts and
radiolarites with thin interbeds of silicified shales.
Nannoconus-radiolarian biomicrites and thin turbidites occasionally occur in some sections. The
thickness of the formation attains 40 m.
This formation is exposed in the La Esperanza
belt, the Northern and Southern Rosario belts, in
the Matanzas Province, in central Cuba and the
Camagtiey Province. The author has not seen the
Santa Teresa Formation deposits in the Sierra de los
Organos belt. Nevertheless, Iturralde-Vinent (1996)
suggests the presence of a radiolarian chert unit
within the pelagic limestones and cherts of the Pons
Formation.
Age. Infrequent planktonic foraminifers of
Albian-Cenomanian age occur in the Santa Teresa
Formation. The lithostratigraphic position suggests
that deposition of this formation began during the
Hauterivian in the Southern Rosario belt and during
the Aptian in the Northern Rosario belt. Radiolarians identified by Aiello and Chiari (1995) indicate
a Valanginian to middle Aptian age of one sample taken in the transitional beds from the Polier
Formation to Santa Teresa Formation.
Facies interpretation. The siliceous rocks of this
formation originated in the deep-water environment,
near (and sometimes below) the calcite compensation depth. Minor terrigenous input was still active
during deposition of the lower part of the Santa
Teresa Formation in the Northern Rosario belt.
Pons Formation (?upper Valanginian-Turonian)
Lithology. The Pons Formation consists of gray
to black micritic limestones interbedded with cherts.
The amount of chert intercalations is variable vertically and laterally. The nannofossil-radiolarian limestones and radiolarian cherts are common in the
Pons Formation, although other microfacies types
are also present in this pelagic sequence. The Pons
Formation is developed in the lower tectonic units
of the Sierra de los Organos belt only (Fig. 13).
The total thickness of Pons Formation is 120 to
150 m.
Age. The ?late Valanginian age of the lowermost
part of Pons Formation may be assumed on the basis
of (1) the lithostratigraphic position of this unit,
and (2) the presence of calpionellids (Tintinnopsella
carpathica Murgeanu et Filipescu) in the limestones
exposed at the base of the type section south of Pons
(Pszcz6tkowski, 1978; de la Torre, 1988). Hatten
(1957) considered the upper boundary of the Pons
Formation to be of Turonian age.
Facies interpretation. The pelagic limestones
and radiolarian cherts were deposited in a deep
bathyal environment, between the aragonite compensation depth and calcite compensation depth.
106
Terrigenous influx was negligible during deposition
of the Pons Formation limestones.
Carmita Formation (Cenomanian-Turonian)
Lithology. The Carmita Formation is composed
of green, red and gray limestones (biomicrites and
calcarenites), radiolarian cherts and shales. The calcarenites (calciturbidites) are up to 10 m thick.
These graded limestones contain minor amounts of
sharp-edged quartz, wacky sandstone fragments, and
plagioclase detrital grains. In thin sections, ooids,
shallow-water and pelagic limestone fragments are
also seen.
Maximum thickness of this formation is 70 m.
This unit is better developed in the Northern Rosario
belt, mainly because of Maastrichtian submarine
erosion in the Southern Rosario belt (Figs. 3 and 6).
Age. The planktonic foraminifers identified in
thin sections by de la Torre (1988) indicate a
Cenomanian-Turonian age for the Carmita Formation. The younger age of the marls and calcareous
shales of the topmost part of this formation, although
possible, has not been confirmed so far.
Facies interpretation. The deep-water deposits
of the Carmita Formation were accumulated in
bathyal environment, below the aragonite compensation depth. The calciturbidites contain a shallowwater debris transported from a carbonate platform.
Pinalilla Formation (Cenomanian-Turonian)
Lithology. This formation is built of thick-bedded to massive, gray-green limestones, up to 170 m
thick. These carbonate rocks are mainly biomicrites
containing planktonic foraminifers. The Pinalilla
Formation is developed in the Quifiones tectonic
unit of the Northern Rosario belt only (Fig. 13).
Age. The planktonic foraminifers identified in
thin sections demonstrate a Cenomanian-Turonian
age. Limestones of the upper part of the formation
are middle to late Turonian in age.
Facies interpretation. The thick-bedded pelagic
limestones of the Pinalilla Formation do not contain
any chert intercalations. These limestones were deposited in a bathyal environment, probably less deep
than in the case of the Santa Teresa and Carmita
formations.
Pefias Formation (Campanian-Maastrichtian)
Lithology. This formation consists of dark-gray
to black, thin-bedded limestones with abundant
black chert intercalations (Fig. 14). The limestones
are biomicrites containing profuse calcified radiolarians and less numerous planktonic foraminifers. The
thickness of the limestones and cherts attains 80 m
A. PSZCZOLKOWSKI
in the type section situated south of Pons, in the
Sierra de los Organos belt. This section occurs in
the Valle de Pons tectonic unit (Fig. 13), while the
rocks of the Pefias Formation have been eroded in
the higher tectonic units of the Sierra de los Organos
belt.
Age. The planktonic foraminifers, studied in thin
sections, indicate a Campanian-Maastrichtian age
for this formation (de la Torre, 1988). According
to Hatten (1957) and Meyerhoff (in Khudoley and
Meyerhoff, 1971) the Pefias Formation is TuronianCampanian in age. Recently, Iturralde-Vinent (1994,
1996) reported the presence in the Sierra de los
Organos belt of a hiatus that spans the Coniacian,
Santonian and Campanian. Unfortunately, the results
of the paleontological study mentioned by this author are still not published. In the present paper, a
hiatus that comprises the late Turonian-Santonian is
accepted (Fig. 6).
Facies interpretation, Pelagic deposits of the
Pefias Formation were accumulated in a bathyal environment, below the aragonite compensation depth.
This part of the basin was effectively protected from
the terrigenous influx.
Moreno Formation (Campanian)
Lithology. The Moreno Formation is composed
of marly limestones, detrital limestones, shales, siltstones and sandstones. The marly limestones prevail
in the lower part of this formation. The detrital
limestones are calciturbidites, up to 5 m thick.
These deposits contain common shallow-water bioclasts. Volcanic lithoclasts and quartz are minor
constituents of calciturbidites. Shales and sandstones
(siliciturbidites) constitute the upper part of the formation. The sandstones contain abundant angular
volcanic lithoclasts and plagioclase. This formation
in known mainly in the Northern Rosario belt. The
maximum thickness of the Moreno Formation rocks
is 240 m, but in places this unit was eroded partly
or completely at the base of the late Maastrichtian
Cacarajfcara Formation (Figs. 4 and 6).
Age. The Campanian age of the Moreno Formation is based on identification of planktonic
foraminifers in thin sections (Pszcz6tkowski, 1994a).
There is a hiatus between the Pinalilla Formation and
the overlying Moreno Formation, which spans the
Coniacian and Santonian (Fig. 6). A similar hiatus
exists between the Carmita and Moreno formations,
although the age of the uppermost deposits of the
former unit is still not too precise.
Facies interpretation. The deep-water, hemipelagic deposits of the lower part of the Moreno
Formation originated in a bathyal environment, in
conditions of increasing influx of shallow-water and
volcaniclastic debris from a volcanic arc terrane.
THE EXPOSED PASSIVE MARGIN OF NORTH AMERICA IN WESTERN CUBA
107
Fig. 14. Pelagic limestones and cherts of the Pefias Formation (Campanian-Maastrichtian); the Las Piedras river, south of Pons, Sierra de
los Organosbelt. Hammerlength 28 cm.
The shales and sandstones of the upper part of the
formation are interpreted as an evidence for convergence between the Upper Cretaceous passive margin
and a volcanic arc terrane approaching from the
southwest. These terrigenous deposits were accumulated in front of this arc terrane. They are preserved
mainly in the southernmost tectonic units of the
Guaniguanico terrane (Fig. 13).
Cacarajicara Formation (upper Maastrichtian)
Lithology. This formation is developed as detrital
limestones composed of breccia, calcarenite and calcisiltite or calcilutite (Hatten, 1957; Pszcz6tkowski,
1978, 1986c, 1994a). The breccia consists of limestone and chert clasts 0.1 to 5 m across, subordinately of shale fragments; upward, it passes gradually into fine calcirudite and coarse calcarenite.
Locally, this breccia is up 180 m thick (the Los
Cayos Member). The fine calcarenite and calcisiltite
constitute the upper part of the formation. Abundant
redeposited shallow-water bioclasts of Cretaceous
age and relatively infrequent grains of volcanic rocks
occur throughout the Cacarajfcara Formation. The
planktonic foraminifers are frequent in the upper,
fine-grained part of the formation. The thickness of
the Cacarajfcara Formation rises from 5-30 m in the
SR up to 450 m in the NR.
Age. Planktonic foraminifers indicate the upper
Maastrichtian Abathomphalus mayaroensis Zone.
The list of the identified taxa was given by
Pszcz6tkowski (1994a).
Facies interpretation. The Cacarajfcara Formation was interpreted as a clastic unit that originated as a result of a single, unusual depositional
event close to the Cretaceous/Tertiary boundary
(Pszcz6tkowski, 1986b). The Pefialver and Amaro
formations, in western and central Cuba, respectively, are stratigraphic and facies equivalents of the
Cacarajfcara Formation. The origin of these massive
megabeds was probably related to an extraordinary
earthquake and tsunami wave at the end of the Maastrichtian. That unusual event could be associated
with the Chicxulub structure of YucaUin (Hildebrand
et al., 1991), although an accurate stratigraphic correlation of the Cuban megabeds with the suggested
impact site is still not known.
Anc6n Formation (Paleocene)
Subdivision and lithology. In the Sierra de los
Organos belt, the Anc6n Formation is subdivided
into the La Gtiira and La Legua Members (Fig. 6),
and the informal member of micritic and marly limestones. In the Southern Rosario belt, the formation
is undivided, although the terrigenous rocks occurring locally in the Cinco Pesos tectonic unit may be
considered as an informal unit (member).
The Anc6n Formation consists of biomicrites,
marly limestones, breccias and locally also of terrigenous deposits. Gray, green, and reddish biomicrites contain planktonic foraminifers, calcareous
nannoplankton and, sometimes, radiolarians. In the
Sierra de los Organos belt, there are some chert
nodules and/or lenses in the green biomicrites of the
lower part of the formation. The breccias composed
of limestone and chert clasts are a characteristic
lithology of the formation in the Sierra de los
Organos belt. The terrigenous rocks occurring in
the Southern Rosario belt contain polymictic sandstones with abundant volcaniclastic fragments. The
total thickness of the Anc6n Formation attains 50
m in the Sierra de los Organos and 120(?) m in
the Southern Rosario belt. Only few outcrops of this
108
A. PSZCZOLKOWSKI
formation are known in the Northern Rosario belt
and the yellowish shales of the Manacas Formation
are the main component of the Paleocene deposits
there.
Age. Thin section analyses of planktonic
foraminifers indicate that in the Sierra de los
Organos belt the Anc6n Formation is Paleocene in
age (Pszcz6~kowski, 1978; de la Torre, 1988). In the
Southern Rosario belt, this formation is of late Early
Paleocene to Late Paleocene age (Pszcz6~kowski,
1994a). The earliest Paleocene deposits are missing
in the Southern Rosario belt.
Facies interpretation. The breccia units originated in proximity of synsedimentary faults affecting
the Jurassic-Cretaceous limestones (Pszcz61kowski,
1978). Some uplifted limestone blocks supplied
a carbonate debris during the Paleocene. Pelagic
biomicrites and marly limestones accumulated in
deeper areas of the sea bottom (submarine depressions). The volcaniclastic debris locally appeared in
the Southern Rosario belt beginning an early stage
of the foreland basin development during the (late)
Early Paleocene.
belt (Pszcz6~kowski, 1994b). The Vieja Member is
Early Eocene in age (Pszcz6~kowski, 1994a,b), but
in some sections its lower part may be as old as
Late Paleocene (Bralower et al., 1993; Bralower and
Iturralde-Vinent, 1997).
Facies interpretation. Large amounts of volcanic
rock fragments, serpentinite and other types of exotic
material in the Manacas Formation define this unit
as a foreland basin sedimentary fill (Iturralde-Vinent,
1995b). The deep-water stage with four phases characterize evolution of this basin: (1) deposition of
fine-grained sediments (mostly shales), (2) turbiditic
sedimentation, (3) development of multi-component
olistostrome and (4) serpentinite slide. Phase (4)
of the Guaniguanico foreland basin evolution was
proposed by Iturralde-Vinent (1996). The deposits
of a shallow-water stage, typical of many foreland
basins (Covey, 1986), have not been reported from
the Manacas Formation.
Manacas Formation (Paleocene-Lower Eocene)
Plate tectonic reconstructions of the Caribbean
region
Subdivision and lithology. This formation is subdivided into the Pica Pica Member (lower) and Vieja
Member (upper) (Fig. 6). The Pica Pica Member
is composed of layered terrigenous deposits: shales,
sandstones and conglomerates. Usually, yellow- or
red-weathering shales are developed in the lower part
of this member. In these shales, there are some breccia interbeds in the southwestern part of the Sierra
de los Organos belt. The lithoclastic sandstones and
conglomerates, sometimes with infrequent intercalations of marly and detrital limestones, predominate
in the upper part of the Pica Pica Member.
The Vieja Member consists of cobbles, boulders
and larger olistoliths of various rocks enclosed in
an argillaceous and/or silty matrix. These olistoliths
were derived from the igneous, metamorphic and
sedimentary rocks. Fragments of rocks derived from
the Guaniguanico stratigraphic successions occur
throughout the whole Vieja Member. Olistoliths of
the serpentinite and other rocks of the ophiolitic
complex are very frequent in this unit. Also volcanic
rocks, derived from the Cretaceous volcanic arc, are
common, especially in the Rosario belts. The total
thickness of the Manacas Formation attains about
500 m in the most complete sections.
Age. Planktonic foraminifers occurring in the
Manacas Formation are Paleocene-Early Eocene
in age (Pszcz6ikowski, 1978, 1994a, 1988). The
lower boundary of the formation is diachronous;
the oldest deposits (late Early Paleocene) of the
Manacas Formation occur in the Northern Rosario
JURASSIC TO EARLY PALEOCENE EVOLUTIONOF
THE PASSIVE MARGIN IN WESTERN CUBA
The Jurassic to Early Paleocene evolution of the
North American margins exposed in Cuba are considered within the tectonic history of the Yucat~in
and Bahamas platform margins, as well as the protoCaribbean basin development. The plate tectonic
reconstructions of the Caribbean region (or the Gulf
of Mexico-Caribbean system) have been outlined in
several papers during the last 15 years (e.g. Pindell
and Dewey, 1982; Anderson and Schmidt, 1983; K1itgord et al., 1984; Ross and Scotese, 1988; Pindell
et al., 1988; Pindell and Barrett, 1990; Marton and
Buffler, 1994).
In the present paper, the plate tectonic reconstruction by Marton and Buffler (1994) is accepted as a
general paleotectonic basis for the Jurassic. These
authors presented the six-stage (Early Jurassic(?) to
Berriasian) rift and drift evolution of the Gulf of
Mexico and adjacent regions using the plate reconstruction software Plates 2.0. A rotation pole for the
Yucat~in block in the southeastern Gulf of Mexico
is proposed for the Callovian to Berriasian drifting
stage (Marton and Buffler, 1994).
The plate tectonic models proposed by Ross and
Scotese (1988) and Pindell and Barrett (1990) are
adopted for the Cretaceous. These models assume
(1) the formation of a proto-Caribbean Sea due to
separation of North and South America, (2) the
development of the Greater Antilles Arc along the
(south)western margin of this sea, and (3) the insertion of the Farallon plate between North and
THE EXPOSED PASSIVE MARGIN OF NORTH AMERICA IN WESTERN CUBA
~ooovv
Callovian
l
/t~
,oovv
/-~",,
I
v
30ON
--
====
'7
~ 1
,'~
.
--~
I
/
I
b*+*+*+1 % % % "
1.+_,.+,+.,_.112 [. =.=,...eJ 3 I"'
--
B
,,
p,"
'
J4 ~ 5
Fig. 15. Reconstruction of the continental blocks and rift basins
between North America an South America during the Callovian
(after Marton and Buffler, 1994, simplified and partly modified by the author): BFZ = Bahamas fracture zone" BP =
Blake-Bahamas plateau; C = Camajuanf belt of north-central
Cuba (Upper Jurassic-Paleogene); E = Escambray terrane; G =
Guaniguanico terrane; MSM = Mojave-Sonora megashear; NA
= North America; NWB = northwestern Bahamas; P = Pinos
terrane; PS -- Placetas belt of north-central Cuba (Proterozoic
marbles, Early to Middle Jurassic granitoids and Late Jurassic to Paleogene deposits)" R = Remedios belt of north-central
and northeastern Cuba (Jurassic?-Late Cretaceous)" SA = South
America; T M V - - Trans-Mexican Volcanic belt; WMT = Western
Main Transform; 1 -- continental blocks, 2 -- Proterozoic marbles and Early to Middle Jurassic granitoids of the Placetas belt,
3 = San Cayetano Formation and its metamorphic equivalents
now exposed in western and south-central Cuba, 4 - rift basins
in the Gulf of Mexico area, 5 = Central Atlantic.
South America resulting in the subduction of protoCaribbean crust.
The position of the Guaniguanico terrane in respect to the Yucatfin block during the Jurassic is
presented herein (Figs. 15-17) mainly after Iturralde-Vinent (1994, 1996). This author located the
Guaniguanico, Pinos and Escambray terranes on the
eastern margin of the Yucatfin platform. In fact, the
exact original position of these Cuban terranes during the Jurassic and Cretaceous is still uncertain.
The location of the Guaniguanico terrane very close
to the present northeastern coast of the Yucatan
Peninsula, as suggested by Iturralde-Vinent (1994,
1996), may be difficult to sustain, mainly because
the remnants of the Mesozoic successions occurring
in western Cuba were not reported from the eastern
margin of the Yucat~.n block (Viniegra, 1971; Lopez
Ramos, 1975).
Probably, the Guaniguanico Mesozoic successions were deposited about 100 km to the east and
northeast of the present northeast Yucatan coast.
Such a possibility is not contradicted by the existing data on the Yucatan borderland topography and
structure (Rosencrantz, 1990, 1996). This province
extends from western Cuba to Honduras and has an
average width of 100 km. In general, the Yucatfin
109
borderland represents the eastern extension of the
Yucatfin platform. However, the northern part of this
borderland and western Cuba may be structurally
continuous (Fig. 1B). The Guaniguanico terrane
rocks and the northern part of the Yucatfin borderland together form a large wedge-shaped block now
isolated within the transform domain (Rosencrantz,
1996, figs. 2 and 4). Deformed metaterrigenous
rocks dredged from the northern Yucatan borderland
(Pyle et al., 1973) may represent equivalents of the
Cangre belt located in the Guaniguanico terrane and
in the subsurface of the Guanahacabibes Peninsula.
According to Pessagno et al. (Chapter 5) the
Jurassic and Early Cretaceous successions in western Cuba (Sierra del Rosario and Sierra de los Organos) show lithostratigraphic, paleobathymetric, and
paleolatudinal signatures which are nearly identical
to those of San Pedro terrane remnants in central
Mexico. Pessagno et al. conclude: "These Cuban
remnants are allochthonous when compared to surrounding Central Tethyan successions in the Blake
Bahama basin and elsewhere in Cuba. They contain
high latitude bivalves such as species of B u c h i a that
can only be derived (exclusive of Greenland) from
a Pacific source. The presence of Southern Boreal/Northern Tethyan faunas (30~ in the Sierra de los
Organos and Sierra del Rosario remnants as late as the
Early Cretaceous (Valanginian) suggests much later
tectonic transport by northwest to southeast movement along the Walper Megashear and by subsequent
southwest to northeast movement as the Caribbean
plate plowed its way through the gap between the
North American and South American plates."
Also Pszcz6tkowski and Myczyfiski 2 suggested
that the occurrence of the bivalves B u c h i a and the
radiolarians P a r v i c i n g u l a sp. and Pantanellidae in
the Tithonian of western Cuba may indicate (according to the criteria of Pessagno et al., 1993) the
Southern Boreal Province or the Northern Tethyan
Province, and that the Jurassic sequences of western
Cuba could have been located in the Pacific during
Tithonian time. They considered, however, that this
interpretation requires further investigations.
A paleomagnetic study of hand-samples collected
from the Jurassic rocks of the Sierra de los Organos
and Sierra del Rosario revealed postfolding magnetisation in mafic rocks of the E1 Sfibalo Formation, but
no meaningful results were derived from samples
of the San Cayetano and Artemisa formations due
to the weak magnetization and/or large scatter of
paleomagnetic data (Bazhenov et al., 1996). P6rez
Lazo et al. (1995) studied their samples from the San
Cayetano Formation of the Sierra del Rosario, and
from the Collantes Formation marbles of the Escam2 'Information and 1994 Annual Report', Institute of Geological
Sciences, Polish Academy of Sciences, Warsaw, 1995, p. 15.
110
A. PSZCZOLKOWSKI
Fig. 16. (A) Reconstruction of the continental blocks around the proto-Caribbean Sea during the middle Oxfordian (partly adapted from
Marton and Buffler, 1994 and Iturralde-Vinent, 1994): NR = Northern Rosario belt, SR = Southern Rosario belt, SO = Sierra de los
Organos belt, E = Escambray terrane, P -- Pinos terrane, BFZ = Bahamas fracture zone. (B) Enlarged rectangle shown in Fig. 1A; 1
= oceanic crust, 2 -- E1 Sfibalo Formation (mainly mafic rocks), 3 = clastic deposits of the San Cayetano Formation and equivalent
metamorphosed units (Arroyo Cangre Formation, and others), 4 -- ammonite-bearing clastic deposits of the uppermost San Cayetano
Formation in the Southern Rosario and La Esperanza belts, 5 = bivalve-bearing deposits of the uppermost part of the San Cayetano
Formation in the Sierra de los Organos and southwestern part of the Southern Rosario belt (mainly the Loma del Muerto tectonic unit), 6
= land areas.
Fig. 17. (A) Reconstruction of the proto-Caribbean basin for the Tithonian, partly adapted after Marton and Buffler, 1994. (B) Enlarged
area indicated by rectangle in (A). Yucatfin passive margin: SO = Sierra de los Organos belt, SR = Northern and Southern Rosario
belts; Florida-Bahamas passive margin: P = Placetas belt, C -- Camajuanf belt, R = Remedios belt; PC = proto-Caribbean Sea, BFZ
--- Bahamas fracture zone; 1 = oceanic crust, 2 -- pelagic ammonite-bearing biomicrites and calcarenites, 3 = terrigenous deposits, 4 =
shallow-water carbonates, 5 - land areas.
b r a y terrane. T h e s e s a m p l e s y i e l d e d p a l e o m a g n e t i c
i n c l i n a t i o n indicative of a p a l e o l a t i t u d e of about 12 ~
Pdrez L a z o et al. (1995) c o n c l u d e d that their paleo m a g n e t i c result c o r r o b o r a t e s the earlier g e o l o g i c a l
i n t e r p r e t a t i o n s c o n c e r n i n g the original l o c a t i o n of
the C u b a n Jurassic rocks " n o t too far f r o m H o n d u r a s
and G u a t e m a l a . . . " (op. cit.), that is, s o u t h w e s t of its
p r e s e n t - d a y position. M o r e o v e r , Pdrez L a z o et al.
(1995) inferred f r o m their p a l e o m a g n e t i c data conc e r n i n g the L o w e r C r e t a c e o u s and A p t i a n - T u r o n i a n
rocks, that the d e v e l o p m e n t of the volcanic arc in
C u b a t o o k p l a c e at p a l e o l a t i t u d e values o f a b o u t
16-17~
C o n s e q u e n t l y , the e n t r a n c e to the protoC a r i b b e a n b a s i n c o u l d be b l o c k e d by the C u b a n or
the G r e a t e r A n t i l l e s volcanic arc (see also R e n n e et
al., 1991 for the A p t i a n - C e n o m a n i a n p a l e o l a t i t u d e
THE EXPOSED PASSIVE MARGIN OF NORTH AMERICA IN WESTERN CUBA
of their Zaza terrane). This result constrains (or even
contradicts?) a possibility of tectonic transport of
the 'Sierra de los Organos and Sierra del Rosario
remnants' by southwest (from Pacific) to northeast
movement after the Barremian?-Aptian.
The location of the San Cayetano basin along
the Yucat~in margin during Middle and Late Jurassic
(Oxfordian) times cannot be defined exactly by the
available geological data. The Mesozoic successions
of western and south-central Cuba suffered severe
tectonic deformations, including large-scale thrusting and metamorphism. During the Paleogene, these
successions were thrust to the north, and probably
sheared off from the Yucatan block, as the inactive
volcanic arc passed along the eastern margin of this
platform (Ross and Scotese, 1988; Pindell and Barrett, 1990; Hutson et al., 1998). The presence of
a transform boundary parallel to the Yucatan borderland shows that the arc moved northward from
a location south of the Yucat~in Peninsula (Rosencrantz, 1996). Nevertheless, various indirect geological evidences constrain the inferred position of the
Guaniguanico terrane during the Jurassic. Only a
part of these arguments may be expressed herein,
because of limited space.
The San Cayetano Formation facies configuration, reconstructed from the present-day internal
tectonostratigraphic pattern of the Guaniguanico terrane, results as roughly parallel to the Yucat~in
eastern margin. The directions of sediment transport from south or south-southwest measured in the
San Cayetano Formation sandstones of the Sierra de
los Organos belt (Haczewski, 1976), are compatible with the Caribbean paleotectonic reconstructions
proposed by Ross and Scotese (1988) and Pindell and Barrett (1990) for the Middle JurassicOxfordian and Late Triassic-Early Jurassic, respectively. In general, the composition of the San
Cayetano sandstones (Pszcz6tkowski, 1986a; Hutson et al., 1998) can be explained on the basis of the
geological structure and Jurassic history of Yucat~in
(Lopez Ramos, 1975).
The shallow-water to neritic San Cayetano, Jagua
and Francisco formations gradually disappear from
northwest to southeast. These units are laterally
replaced by the deep neritic E1 S~ibalo Formation. The carbonate bank of the Sierra de los
Organos belt (Pszcz6lkowski, 1978, 1981) is a
clear evidence of shallowing phase close to the
Oxfordian/Kimmeridgian boundary; this phase is
characteristic for the Late Jurassic successions of
western Cuba. A deepening phase occurred at the
Kimmeridgian/Tithonian boundary or during the
earliest Tithonian. Both events seems to be unknown
in Mexico, especially in the sections described by
Pessagno et al. (Chapter 5) as belonging to the San
Pedro del Gallo terrane.
111
The Oxfordian faunal assemblages of the San
Cayetano and Jagua formations do not contain
any taxa characteristic for high latitudes (see also
Wierzbowski, 1976 and Myczyfiski, 1976, 1994a).
These taxa were also not reported from the Kimmeridgian carbonate rocks of the Sierra de los
Organos and Sierra del Rosario belts. In the Tithonian limestones of the Sierra de los Organos belt
there are frequent specimens of bivalves belonging to the genera Anopaea and Buchia (Myczyfiski,
1999). Some specimens of Anopaea sp. resemble
taxa described earlier from Tithonian deposits of
Antarctica, Himalayas, New Zealand, and Sula Islands (Myczyfiski, 1999). In the Sierra del Rosario,
the Tithonian bivalves belong mainly to the genus
Inoceramus (Parkinson, 1819), although representatives of Buchiidae are also present (Myczyfiski,
1994b). The occurrence of Tithonian bivalves characteristic for northern and southern high latitudes
may be explained by upwelling of deeper waters in
the northwestern part of the proto-Caribbean basin
(Coleman et al., 1995).
The Tithonian ammonites of western and central
Cuba (Imlay, 1942, 1980; Millfin and Myczyfiski,
1978; Myczyfiski, 1989, 1994a; Myczyfiski and
Pszcz6tkowski, 1990, 1994) are so similar, that it
would be unrealistic to place them in two different,
widely separated, oceanic basins. Therefore, any attempt to locate the Guaniguanico terrane at a higher
paleolatitudinal position (30~
during the Tithonian, far from the proto-Caribbean basin, should be
accompanied by a similar shift of the Escambray
terrane, as well as of the Placetas and Camajuanf
belts of central Cuba. In this case, the Guaniguanico,
Escambray (and Pinos) terranes and the Placetas and
Camajuanf belts must be considered as parts (fragments) of a composite terrane, or superterrane, 8001000 km long and 150-200 km wide, transported
(1) from NW to SE, and later (2) from SW to NE
(or NNE). However, this tectonic shift would leave
the Bahamas platform and the Yucatan block without deep-water successions of their proto-Caribbean
continental slope and adjacent basin floor. Considering the aforementioned problems, it seems that the
hypothesis of tectonic transport of the Sierra de los
Organos and Sierra del Rosario successions along
the Walper Megashear still needs a lot of evidences
and probably additional studies.
Syn-rift stage (Lower Jurassic-?Callovian/early
Oxfordian)
South America was close to the Yucatan block
during the late Middle Jurassic (Fig. 15). Some leftlateral transform motion probably occurred between
these two continental blocks (Marton and Buffler,
1994). Such a situation could exist also during the
112
early Oxfordian with a rift zone (Iturralde-Vinent,
1994, 1996) or rift/spreading center (Marton and
Buffler, 1994) between the two continental blocks.
The basement of the Mesozoic sedimentary successions, that originally belonged to the Yucat~in
eastern margin, is unknown in the Guaniguanico
terrane. The thrust units of this terrane do not
contain any basement rocks at their base, because
the initial tectonic detachment occurred above the
basement/sedimentary cover boundary.
During the sedimentation of the San Cayetano
deltaic deposits, their source areas probably consisted of metasedimentary and terrigenous rocks, and
also granitoids (Pszcz6tkowski, 1978). A prolonged
transport, recycling of the pre-Middle Jurassic terrigenous rocks, and weathering resulted in the high
content of quartz in the San Cayetano sandstones
(Pszcz6tkowski, 1986a). In the Southern Rosario
belt, rare pebbles of the heavily silicified limestones
occur in the thick-bedded pebbly sandstones. Some
of those pebbles contain the late Paleozoic fossils
(Pszcz6tkowski, 1989b). Evidently, these pebbles
were derived from the limestone succession situated
some hundreds kilometers from the San Cayetano
basin during the early to middle Oxfordian. According to Donnelly et al. (1990) a correlation of
fossiliferous pebbles found in the San Cayetano
Formation with Permian beds of eastern Guatemala
is not apparent. However, any conclusive explanation of the provenance of these pebbles is difficult,
mainly because of small amount and size of fossiliferous clasts.
At present it is not possible to demonstrate any
direct connection between the San Cayetano basin
and the northwestern edge of South America as
a source of the clastic material for the Jurassic
deposits in western Cuba. The hypothesis for a South
American provenance for the San Cayetano clastics
expressed by some authors (Anderson and Schmidt,
1983; Klitgord et al., 1984; Ryabukhin et al., 1984)
is still to be proved. Apparently, the fossils and
microfacies of the silicified limestone pebbles from
the San Cayetano Formation are not similar to the
Permian limestones occurring in the Palmarito and
Tucutunemo formations known in northern South
America (Benjamini et al., 1987).
Instead, some data are consistent with a possible
connection of the San Cayetano depocenter with
Central America (Meyerhof, in Khudoley and Meyerhoff, 1971; Meyerhoff and Hatten, 1974; IturraldeVinent, 1975; Haczewski, 1976; Pszcz6ikowski,
1987). Prior to the Oxfordian, the San Cayetano
clastics were located in the originally narrow, but
steadily widening, rift zone formed between Yucat~in and South America (Fig. 15). According to
Hutson et al. (1998), the presence of grains with
Taconic and Grenvillian ages supports a Yucat~in
A. PSZCZOLKOWSKI
source for the San Cayetano Formation. The extent
of the San Cayetano depocenter to the southwest,
along the Yucat~in margin, is still uncertain. The
model of propagating westward rift/spreading center between Yucat~in and South America, with simultaneous counterclockwise rotation of the former
block (Marton and Buffler, 1994), requires facies
shift from the northeast to the southwest in the
San Cayetano Formation. Indeed, some existing data
agree with this paleotectonic model. The most obvious observation pertains to the facies changes within
the San Cayetano deposits in the Southern Rosario
belt between La Palma and Mantua (Pszcz6ikowski,
1994b). From NE to SW, the facies G-I disappear
and facies A-C and E (Haczewski (1976) are dominant southwest of Minas de Matahambre, mainly in
the Loma del Muerto tectonic unit (Fig. 2). These
facies changes are parallel to the regional strike of
the present-day tectonic structures (NE-SW). The
San Cayetano Formation deposits, about 1400 m
thick, also occur in the subsurface of the southwestern part of the La Esperanza belt (Los Arroyos
1 w e l l - Fern~indez et al., 1987; for location see
Fig. 1). Deposition of the San Cayetano clastics was
accompanied by syn-sedimentary magmatic activity, mainly of a mafic character (Piotrowski, 1977;
Pszcz6tkowski, 1978; Iturralde-Vinent, 1995a).
Drift stage (?Callovian/middle OxfordianSantonian)
During the middle Oxfordian, the gap between
Yucat~in and northwestern South America widened,
and the narrow proto-Caribbean seaway was formed.
A paleogeographic location of the Guaniguanico,
Escambray and Pinos terranes during middle Oxfordian time is shown in Fig. 16A. The presence of
the mafic rocks in the Northern Rosario belt (the
E1 S~ibalo Formation) suggests that the oceanic crust
formation in this gap commenced during the middle
Oxfordian or earlier, south of the Northern Rosario
and La Esperanza belts.
Recently, the Callovian sea-floor spreading was
interpreted to have started simultaneously in the Gulf
of Mexico and in the Caribbean (Marton and Buffler,
1994).
Marine conditions developed before the middle
Oxfordian in all belts of the Guaniguanico terrane.
In the Sierra de los Organos succession, the first limestone intercalations with marine fauna occur about
400 m below the top of the San Cayetano Formation.
Nevertheless, exact timing of the onset of marine deposition in all belts of the Guaniguanico terrane is
still to be established. Fig. 16B shows a reconstruction of facies during the middle Oxfordian. Infrequent ammonites have been found only in the uppermost San Cayetano clastics in the Southern Rosario
THE EXPOSED PASSIVE MARGIN OF NORTH AMERICA IN WESTERN CUBA
belt. These ammonite-beating clastic deposits accumulated between the E1 S~balo limestones and basalts
to the southeast and the shallow-water bivalve-bearing sandstones and shales to the northwest.
The advance of the middle Oxfordian transgression resulted in a major facies change, when
the San Cayetano deltaic sediments were replaced
by limestones with bivalves and shales with ammonites (Sierra de los Organos and Cangre belts)
and ammonite-bearing limestones and shales (Southern Rosario belt). Facies differentiation in the W - E
(or NW-SE) direction existed during deposition of
the Jagua and Francisco formations. Surprisingly,
marine macrofossils were not found in the sedimentary rocks of the E1 S~balo Formation. The
limestones occurring between the mafic rocks of
this formation contain an impoverished microfossil assemblage dominated by Globochaete alpina
Lombard indicating a deeper (and partly restricted?)
depositional environment. The contact between the
E1 S~balo and Artemisa formations is erosional and,
in some sections, tectonic. Locally, thin breccia with
volcanic clasts occurs at the E1 S~balo/Artemisa
formations boundary. This boundary may be interpreted as the unconformity below the basal Artemisa
Formation of ?late Oxfordian-early Kimmeridgian
age roughly correlatable with the limestone breccia
at the base of the Guasasa Formation in the Sierra de
los Organos belt.
The onset of the carbonate shallow-water sedimentation in the Sierra de los Organos and Cangre
belts occurred in the late Oxfordian or earliest Kimmeridgian. Shallow-water carbonates are known also
from the Southern Rosario and Northern Rosario
belts. Development of the shallow-water bank above
the E1 S~balo Formation in the Northern Rosario belt
(in the Bel6n Vigoa and Naranjo tectonic units
Fig. 13), with manifestations of erosion of basalts
and diabases, indicates that this volcano-sedimentary
sequence has been locally uplifted during late Oxfordian or earliest Kimmeridgian time. These local uplifts (rotated fault blocks?) could be a barrier inhibiting free communication of the Sierra de los Organos
and Southern Rosario belts with the open, but still
narrow, proto-Caribbean Sea. The fine-grained limestones (Fig. 10) and clastics (mainly shales and
siltstones) accumulated in the inner, semirestricted,
part of the Southern Rosario belt. During the Kimmeridgian, subsidence kept pace with the relatively
high rate of sedimentation. About 400 to 650 m of
shallow-water limestones and dolomitic limestones
formed in the Sierra de los Organos belt (San Vicente Member in Figs. 6 and 13).
The transition from shallow-water deposition to
pelagic conditions of sedimentation occurred close
to the Kimmeridgian/Tithonian boundary. In the
Sierra de los Organos, this change was rather grad-
113
ual, with appearance of some pelagic microfossils
(Saccocoma sp., Colomisphaera spp. ) and deposition of a few thin-bedded limestone units within the
thick-bedded to massive calcarenites of the upper
part of the San Vicente Member. In the Northern Rosario belt, the lower Tithonian limestones
overlay the Kimmeridgian shallow-water dolomitic
limestones.
Drowning of the shallow-water carbonates resulted in a considerable uniformity of facies in all
belts of the Guaniguanico terrane (Fig. 17). At Tithonian time, the sedimentation rate (8-10 m/m.y.) of
the black pelagic limestones (Fig. 11) was some
ten times lower than that of the Kimmeridgian
shallow-water carbonates (80-100 m/m.y.). The
Saccocoma-Didemnidae microfacies that predominated in the lower Tithonian ammonite-bearing
limestones, was replaced gradually by a radiolarian microfacies (Myczyfiski and Pszcz6~kowski,
1994). The latter microfacies is typical for the upper
Tithonian and Lower Cretaceous limestones in all
Guaniguanico belts. In the Rosario belts, favorable
conditions for radiolarians existed since the early
Tithonian. Moderate fertility of the northern protoCaribbean surface waters is also suggested by an
elevated content of phosphatic grains, mainly abundant fish debris (bones, scales and teeth) in the upper
Tithonian limestones of the Sierra de los Organos.
Radiolarian limestones are also common in the Camajuanf succession of central Cuba (Fig. 17) and in
the Camagfiey Province.
In some sections, there are frequent specimens of
inoceramids belonging to the genera Anopaea and
Buchia (identified by Dr. R. Myczyfiski). The occurrence of Tithonian bivalves characteristic for high
latitudes is consistent with upwelling of cold, oxygenated, and nutrient-rich deeper waters in the northern (or northwestern) part of the proto-Caribbean
basin. According to Baumgartner (1987), off-shore
winds created upwelling and high fertility of the surface waters in the Late Jurassic 'Caribbean Tethys'.
A Jurassic/Cretaceous boundary event, marked
by positive shift in the both carbon and oxygen
curves seems to be specific for the proto-Caribbean
basin, or even for the Caribbean-Gulf of Mexico
region (Coleman et al., 1995). The possible explanation for this proto-Caribbean event is the invasion
of geochemically different water masses across the
Jurassic-Cretaceous boundary (op cit.).
During the Berriasian and Valanginian, pelagic
limestones with chert interbeds accumulated in the
northwestern part of the proto-Caribbean basin. By
late Berriasian-early Valanginian, gray nannoconid
limestones with abundant calpionellids were deposited in the Sierra de los Organos belt. Basinward, these thick-bedded, pure pelagic limestones
passed gradually into black, thin-bedded radiolar-
114
A. PSZCZOLKOWSKI
ian limestones occurring in the Northern Rosario
belt. The ammonites are uncommon in the pelagic
limestones of Berriasian and Valanginian age (Myczyfiski, 1977). At the Tithonian/Berriasian boundary, the basin floor probably descended below the
aragonite compensation depth.
In the Northern Rosario belt (farthest south sections: BV, NO, CE, CH and QS, in palinspastic reconstruction
Fig. 13), the first turbidite
sandstones appeared in the late Berriasian deposits.
Nevertheless, the main influx of siliciclastic material in this belt occurred during the ValanginianBarremian. A similar petrographic composition of
Hauterivian-Barremian siliciturbidites occurs in the
Northern Rosario belt, La Esperanza belt and Placetas belt in the Matanzas Province and north-central
Cuba indicating a common source for the clastic
material (Fig. 18). The Northern Rosario and La
Esperanza belts were situated nearer to the source
area. This conclusion results from established differences in abundance and thickness of siliciturbidites
between the Northern Rosario and La Esperanza
belts and Placetas belt (Pszcz6{kowski, 1982, 1987).
The source area probably was located at the northeastern end of the Yucatfin block (Fig. 18). However,
terrigenous deposits do not appear in the Sierra de
los Organos and Southern Rosario belts, although
they were situated relatively close to this hypothetical source area. Accepting the idea that the
Paleogene thrusting completely reversed the relative
positions of the belts and tectonic units (Fig. 13),
one should account for the lack of terrigenous ma-
terial in the above-mentioned belts. Probably, the
Northern Rosario and Placetas belts belonged to the
deep-water sector of the basin which extended between the Yucatfin and Bahamas passive margins and
the speculative spreading zone (ridge?) responsible
for the generation of the proto-Caribbean oceanic
crust. This deep-water part of the basin was much
narrower at the Yucatfin-Florida Straits, creating a
fan-like mode of the sediment transport and dispersal. Nevertheless, turbidity currents could not reach
more marginally (and upslope?) located sedimentary successions now occurring in the Sierra de los
Organos and Southern Rosario belts.
In some sections, thin turbidites occur also among
the radiolarian cherts in the lower part of the Santa
Teresa Formation. The influx of the terrigenous material ceased during the Aptian-Albian(?). At the
end of the Early Cretaceous, the siliceous deposition
extended across the entire deeper part of the northwestern proto-Caribbean basin. The Santa Teresa
Formation appears in all belts related to the Yucat~in
passive margin, except the Sierra de los Organos
and the Guajaib6n-Sierra Azul belts. This formation
occurs also in the Placetas belt, originally located
south of the Bahamas platform, represented in central Cuba mainly by the Remedios belt (Fig. 18).
These radiolarian cherts of Early Cretaceous to early
Cenomanian age are not known in the southeastern Gulf of Mexico. Probably, their accumulation in
the northwestern part of the proto-Caribbean basin
was a net result of several different reasons (basin
margins subsidence since the Tithonian, eustatic sea-
Hauterivian - Barremian
~Gulf of Mexico):~ ~
"~o#/^'~,/'"-,. I
FB
,
"~-~
'
~
/
'
""".
"
I- V-.-2..:.:.'.
:...." ~
,
500 km
!
Fig. 18. (A) Lower Cretaceous (Hauterivian-Barremian) paleogeography of the proto-Caribbean and adjacent areas. (B) Enlarged area
indicated by rectangle in (A), with suggested provenance and distribution of terrigenous sediments (siliciturbidites). NA - North America,
SA = South America, FB -- Florida-Bahamas block, SO = Sierra de los Organos belt, SR = Southern Rosario belt, NR = Northern
Rosario belt, LE = La Esperanza belt, P -- Placetas belt, C = Camajuanf belt, R - Remedios belt; 1 - oceanic crust, 2 = radiolarian
cherts and shales, 3 -- pelagic limestones, 4 - shallow-water carbonates, 5 = suggested distribution of siliciturbidites, 6 = land area.
Black arrow indicates probable provenance of the terrigenous sediments.
THE EXPOSED PASSIVE MARGIN OF NORTH AMERICA IN WESTERN CUBA
level rise, paleogeographic and paleoceanographic
conditions).
During the Cenomanian, the pelagic carbonate
sedimentation was restored in the Rosario and Placetas belts. The Turonian deposits occur in the
Pinalilla Formation and in the upper part of the
Carmita Formation, and were also reported from
the Pons Formation of the Sierra de los Organos
belt (Hatten, 1957). The Coniacian-Santonian, or
even late Turonian-Santonian, deposits are very
scarce, or even entirely missing (?) in western Cuba,
due to non-deposition and, sometimes, Late Cretaceous and/or Paleocene erosion (Figs. 4 and 6). The
Carmita Formation (Cenomanian-Turonian) pelagic
limestones are more clayey in their uppermost part
and are overlain by the marls and limestones of the
Campanian Moreno Formation. The thick-bedded
pelagic limestones of the Pinalilla Formation occur
immediately below the thin-bedded limestones and
shales of the Moreno Formation, and no traces of
erosion are discernible along the Pinalilla/Moreno
formations boundary.
The origin of a regionally extensive late
Turonian-Santonian hiatus in the deep-water,
pelagic sequence is probably related to paleoceanographic conditions existing during Late Cretaceous
times in the northwestern part of the proto-Caribbean
basin. For example, this hiatus is coincident with the
maximum flooding of the South American continent in northwestern Venezuela during the Late
Cretaceous highstand of sea level (Lugo and Mann,
1995). In addition, during the Turonian-Santonian,
the Nicaraguan Rise-Greater Antilles Arc partially
closed the connection of the proto-Caribbean basin
with the Pacific (Pindell and Dewey, 1982; Pindell, 1991). However, the influence of this paleogeographic change on the sedimentation in the
proto-Caribbean basin is still to be evaluated.
The beginning of the active margin stage
(Campanian-Paleocene)
According to Pindell and Dewey (1982), Ross
and Scotese (1988) and Pindell (1991), the protoCaribbean ocean basin has been progressively subducted beneath the Caribbean plate during the Late
Cretaceous-Eocene. An arc terrane collided with
the southern Yucatan margin during the Campanian to early Maastrichtian (Donnelly, 1989; Pindell
and Barrett, 1990). The Campanian Sepur clastics
filled the foredeep in front of the arc terrane (Sp in
Fig. 19).
The northeastern passive (proto-Caribbean) margin of Yucatan was also affected by the approaching arc terrane in the Campanian. In the Northern
Rosario belt of western Cuba, the Moreno Formation
(MR in Fig. 19) contains abundant volcaniclastic
115
Fig. 19. Generalized paleogeographic reconstruction of the protoCaribbean basin for the late Campanian (tectonic framework
adapted from Ross and Scotese, 1988 and Pindell et al., 1988).
SP = Sepur foredeep, MR = Moreno depocenter, VB = Vfa
Blanca deep-water basin (western Cuba), SJM
San Juan y
Martfnez shallow-water basin (western Cuba); 1 = Caribbean
plate oceanic crust, 2 = proto-Caribbean and Atlantic oceanic
crust, 3 = pelagic limestones and cherts (Pefias Formation in the
Sierra de los Organos belt), 4 = shales, volcaniclastic sandstones
(turbidites) and marly limestones of the Moreno Formation, 5 =
lack of the Campanian deposits, 6 = shallow-waterlimestones of
the Remedios belt in north-central Cuba, 7 -- deep-water clastics
of the Vfa Blanca Formation, 8 = shallow-water limestones and
conglomerates of the San Juan y Martfnez Formation. Heavy line
with triangles denotes subduction of the proto-Caribbean crust
beneath the Greater Antilles Arc.
=
material, mainly in its upper part (Pszcz6~kowski,
1994a,b). The calciturbidites with volcanic lithoclasts occur in the lower part of this formation. The
petrographic composition of the detrital limestones
and sandstones clearly indicates the volcanic arc
source for the Moreno Formation turbidites.
During the Campanian, the volcanic arc was located east of the Yucatan block margin and south
of the Moreno depocenter. This arc could be the
westernmost part of the Greater Antilles Arc (GAA),
as proposed by Pindell and Dewey (1982) and Pindell et al. (1988). The position of the GAA shown
in Fig. 19 is, in general, in accordance with the
tectonic reconstructions by Ross and Scotese (1988)
and Pindell and Barrett (1990). The results of the paleomagnetic investigations published so far (Renne
et al., 1991; Chauvin et al., 1994; P6rez Lazo et al.,
1995; Bazhenov et al., 1996) indicate between 550
and 1600-+-600 km northward displacement of the
Zaza volcanic arc (or Zaza terrane) during the Late
Cretaceous, Paleocene and Early Eocene.
Although the Moreno depocenter could be connected with the Sepur foredeep, their Campanian
deposits were accumulated in different tectonic and
paleobathymetric settings (Fig. 19). The Sepur For-
116
A. PSzCZOLKOWSKI
Fig. 20. Paleogeographic reconstruction of the proto-Caribbean
basin for the latest Maastrichtian; modified tectonic framework
adapted from Ross and Scotese (1988, fig. 9): Ca -- Cacarajfcara
Formation (western Cuba), Am -- Amaro Formation (north-central Cuba), Pr -- Pefialver Formation (western Cuba), SJM
San Juan y Martfnez shallow-water basin (western Cuba), Cf
= Cienfuegos basin (south-central Cuba); 1 = Caribbean plate
oceanic crust, 2 -- proto-Caribbean and Atlantic oceanic crust,
3 = pelagic limestones and cherts of the Pefias Formation, 4
-- detrital limestones of the Cacarajfcara, Amaro and Pefialver
megabeds (maximum thickness of 200-450 m), 5 -- pelagic
limestones and cherts of the Camajuanf belt in north-central
Cuba, 6 = shallow-water limestones of the Remedios belt in
north-central Cuba, 7 = shallow-water limestones of the San
Juan y Martfnez Formation in western Cuba and marls of the
lowermost part of the Vaquerfa Formation in the Cienfuegos
basin. Heavy line with black triangles denotes subduction of the
proto-Caribbean crust, while heavy line with open triangle marks
underthrusting.
=
mation was deposited on a carbonate shelf in southern Yucatan. This shelf was depressed and buried
by the Sepur serpentinite-bearing flysch (Pindell
and Barrett, 1990). The Moreno marly limestones
and siliciturbidites were laid down in a deep-water
basin on the thinned continental crust and partly
on oceanic crust. During the Campanian, the Sierra
de los Organos belt was not affected by the influx
of volcaniclastic debris, as this material was not
reported from the pelagic limestones and cherts of
the Pefias Formation. Campanian sediments were not
preserved (or deposited?) across the vast area situated between the Moreno depocenter, Sierra de los
Organos belt and the Bahamas platform and slope
(Fig. 19).
The Maastrichtian deposits of the passive margin of Yucatan are represented by the Cacarajfcara
Formation in the Rosario belts, and by the pelagic
limestones and cherts of the Pefias Formation in
the Sierra de los Organos belt (Figs. 4, 6 and
14). The location of these deposits is shown in
paleogeographic reconstruction for the latest Maastrichtian (Fig. 20). South of the Bahamas plat-
form, the limestones and cherts of the Lutgarda
Formation are known in the Camajuanf belt, while
the Amaro Formation is a characteristic unit occurring in the deep-water Placetas belt (Am in
Fig. 20). The Amaro Formation is an equivalent
of the Cacarajfcara and Pefialver formations in western Cuba. The peculiar character and origin of the
late Maastrichtian Cacarajfcara, Amaro and Pefialver
megabeds was studied by Pszczdlkowski (1986b)
and was also discussed by Iturralde-Vinent (1992).
The latest Maastrichtian paleogeography, shown in
Fig. 20, is partly based on conclusions presented
in these papers. However, the position of GAA in
respect to the Bahamas platform is more southerly in
Fig. 20 than that assumed earlier by Pszcz6tkowski
(1986b) and Iturralde-Vinent (1992). Such a position of the GAA at the end of the Cretaceous is
inferred from the paleotectonic reconstruction and
lithology of the late Maastrichtian deposits, as the
Cacarajfcara Formation does not contain a significant
amount of a coarse-grained volcaniclastic material
(> 1 cm) derived from the extinct volcanic arc, even
in the Cangre and Sierra Chiquita tectonic units.
The thickest sections of the Cacarajfcara megabed
clastic deposits (450 m) were measured in these
two tectonic units (Fig. 13), and this fact shows
that their position was well to the south (Fig. 20),
not at the entrance of the Gulf of Mexico between
Florida and Yucatfin, as proposed by Iturralde-Vinent
(1992).
During the Paleocene, the convergence of the
extinct GAA segment with the Bahamas platform
margin continued (Fig. 21). The position of the extinct, westernmost GAA segment (occurring now in
the Bahia Honda terrane) at the Yucatfin margin as
indicated in Fig. 21, results from a tectonic and
lithostratigraphic analysis of the Lower Paleocene
deposits in the Guaniguanico terrane. In the Northern Rosario belt, the Manacas Formation shales
overlying the Cacarajfcara Formation are evidence
of a major change from pelagic conditions during
the Cretaceous to the Paleogene foreland basin environment. The narrow northwestern sector of the
proto-Caribbean basin was either a peripheral or
a retroarc foreland basin, located in front of the
thrust belt along the southern side of the remnant
proto-Caribbean Sea. The discrimination of ancient
peripheral and retroarc foreland basins is difficult
(Ingersoll, 1988). Within the plate tectonic model
accepted herein (Figs. 19-21), the Paleocene basin
of western and central Cuba was a peripheral foreland basin. However, according to an alternative
plate-tectonic model (Iturralde-Vinent, 1994, 1996)
an retroarc foreland basin formed in western and
central Cuba during the Paleocene.
The material derived from the volcanic suites and
ophiolite contributed to the foreland basin deposits
THE EXPOSED PASSIVE MARGIN OF NORTH AMERICA IN WESTERN CUBA
117
CONCLUSIONS
Fig. 21. Simplified paleogeographic reconstruction of the northern Caribbean region for the Early Paleocene: Vb -- Vfbora basin
(western Cuba), Cf = Cienfuegos basin (south-central Cuba),
T.f = La Trocha fault in central Cuba (see Hatten, 1967); 1
-- Caribbean plate oceanic crust, 2 = proto-Caribbean and Atlantic oceanic crust, 3 = pelagic biomicrites and breccias of the
Anc6n Formation (western Cuba), 4 = shales and claystones
(Manacas and Vega Alta formations), 5 - breccias and calcirudites of the Vega Formation (Camajuanf belt of the Bahamas
platform margin), 6 = Remedios belt in north-central Cuba, 7 =
syn-sedimentary normal faults. Other symbols as in Fig. 20.
in western Cuba. Pelagic limestones prevailed in the
Southern Rosario belt (Anc6n Formation), although
with a clear influence of the arc-originated detritus
in the Cinco Pesos tectonic unit. Pelagic limestones
and breccias are widespread in the Sierra de los
Organos belt. The limestone and chert breccias were
formed as a result of a considerable erosion of the
underlying Cretaceous limestones (in places also
Tithonian), along syn-sedimentary fault escarpments
(Pszczdtkowski, 1978). These faults, schematically
shown in Fig. 21, could originate in a zone of
extension induced by bending of the underthrusting
plate during arc-passive margin collision (Bradley
and Kidd, 1991). The Paleocene-Middle Eocene
limestone breccias with a considerable thickness are
also known in the Camajuanf belt of west-central and
central Cuba (Pszcz6tkowski, 1983). The JurassicCretaceous sedimentary successions deposited on
(and along) the passive margin of Yucatfin, now
exposed in western Cuba, formed the foreland basin
substrate during the Paleocene-Early Eocene. In
the Sierra de los Organos belt, a change from the
passive margin to a foreland basin occurred during
the Early to Late Paleocene. The terminal collision
of the extinct volcanic arc with the passive margin
occurred in the Late Paleocene to Early Eocene in
western Cuba (Bralower et al., 1993; Bralower and
Iturralde-Vinent, 1997; Gordon et al., 1997).
The Mesozoic successions of western Cuba, now
exposed in the Guaniguanico terrane, were deposited
more than 100 km to the east of the present northeast Yucat~in coast. The evolution of these Yucat~in
passive margin successions encompasses the synrift stage (Lower Jurassic-?Callovian/early Oxfordian), drift stage (?Callovian/middle OxfordianSantonian), and the beginning of the active margin
stage (Campanian-Paleocene). Prior to the middle
Oxfordian, the San Cayetano basin was located in
a narrow, but steadily widening, rift zone formed
between Yucatfin and South America. The advance
of the middle Oxfordian transgression resulted in a
major facies change, when the San Cayetano deltaic
sediments were replaced by shallow-water limestones with bivalves and/or by deeper ammonitebearing deposits. The restoration of the carbonate
shallow-water sedimentation in the Sierra de los
Organos and Cangre belts and its onset in the
Rosario belts occurred in the late Oxfordian or earliest Kimmeridgian. Subsidence kept pace with the
relatively high rate of sedimentation in the Sierra de
los Organos belt; about 400 to 650 m of shallowwater limestones and dolomitic limestones formed
during the Kimmeridgian.
Drowning of the shallow-water carbonates in the
early Tithonian resulted in a considerable uniformity
of pelagic facies in all belts of the Guaniguanico
terrane. The Hauterivian-Barremian siliciturbidites
that occur in the Northern Rosario, La Esperanza
and Placetas belts of western and central Cuba are
interpreted to have a common source for the clastic material. These belts probably belonged to the
deep-water sector of the basin, which extended between the Yucat~in and Bahamas passive margins.
During the Aptian-Albian, the siliceous deposition
extended across the entire northwestern, deeper part
of the northwestern proto-Caribbean basin. In the
Cenomanian, pelagic carbonate sedimentation was
restored in the Rosario and Placetas belts. The late
Turonian (or Coniacian)-Santonian deposits are very
scarce, or even entirely missing in western Cuba, due
to non-deposition and the Late Cretaceous and/or
Paleocene erosion. During the Turonian-Santonian,
the Nicaraguan Rise-Greater Antilles Arc partially
closed the connection of the proto-Caribbean basin
with the Pacific. Among other factors, this paleogeographic change could create specific paleoceanographic conditions in the northwestern part of the
proto-Caribbean basin.
The eastern passive margin of Yucatfin was affected by approaching arc terrane in the Campanian.
In the Northern Rosario belt of western Cuba, the
Moreno Formation contains abundant volcaniclastic material, mainly in its upper part. During the
118
C a m p a n i a n , the volcanic arc was located east of
the Yucatfin b l o c k m a r g i n and south of the M o r e n o
depocenter. This arc could be the w e s t e r n m o s t part
of the G r e a t e r Antilles Arc ( G A A ) , as p r o p o s e d by
Pindell and D e w e y (1982) and Pindell et al. (1988).
T h e late M a a s t r i c h t i a n deposits of the passive margin of Yucatfin are r e p r e s e n t e d by the Cacarajfcara
F o r m a t i o n in the Rosario belts. A m o r e southerly
position of the extinct volcanic arc at the end of
the C r e t a c e o u s is inferred from the p a l e o t e c t o n i c rec o n s t r u c t i o n and lithology of the late M a a s t r i c h t i a n
deposits.
D u r i n g the P a l e o c e n e , the s e d i m e n t a r y successions of the G u a n i g u a n i c o terrane, originally deposited on (and along) the passive m a r g i n of Yucatfin, f o r m e d the foreland basin substrate. This foreland basin was located in front of a thrust belt along
the southern side of the r e m n a n t p r o t o - C a r i b b e a n
Sea.
ACKNOWLEDGEMENTS
T h e author is grateful to Richard T. Buffler,
T h o m a s W. Donnelly, John E L e w i s and G y 6 r g y
M a r t o n for their review of the manuscript, and to
Paul M a n n for his useful c o m m e n t s on the text and
figures. T h e discussions with R y s z a r d Myczyfiski
and M a n u e l Iturralde-Vinent on s o m e p r o b l e m s concerning the g e o l o g y of Cuba are appreciated.
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