the exposed passive margin of north america in western...

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 Matan- zas 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 sim- ilarities existing between the Guaniguanico, Pinos and Escambray terranes (Fig. 1) clearly suggest their paleogeographic proximity prior to the Late Creta- ceous 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; Rigassi- Studer, 1963; Meyerhoff, in Khudoley and Mey- erhoff, 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. PSZCZ()LKOWSKI

A N Igl

HAVANA

Ill,, K'~N e \.d, f

Camagoey Province

200 km - ~ . ~ . ,

SGM ~ F L O R I P A .A.Np.BAHA.M.AS!: ~ U ~ / ~ , ~ _ ~ " q l ' q l N N N N N N N ' q I ' q l N '

j K I j ~_ I i ~ , . . / ~) "~11-1~ . _ N IN IR"q "

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. How- ever, 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 (Iturralde- Vinent, 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 95

2 2 0 3 0 '

84000 '

I I ,co

o

LM

APS

0 S

83o30 '

0A

/ Y jv /

/

Mantua (Y LP

0 Pinar del R{o

La Co[oma <

TERRAN 83c,00 '

I

oPinar del R~o

0 L , 25 km

N g ~ / 7 0 20km - I I l

84o00' 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: BV = 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 Neogene- Quaternary deposits. The wells drilled in the Gua- nahacabibes 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 con- tinues to the southwest as a fault-bounded, wedge- shaped 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; Ro- driguez, 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 A. PSZCZOLKOWSKI

and CB represent the Jurassic platform that subsided in the Tithonian and remained submerged in deep- water 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 Mi- nas 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 necessar- ily atop the deep-water Lower Cretaceous deposits of the Northern Rosario belt as suggested by Itur- ralde-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.

Fig. 3. Lithologic column of the Martin Mesa 1 well, located northeast of Cayajabos, in the Havana Province (data from Se- gura 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; Pi- otrowska, 1993; and others). In the present paper, the author accepts the idea, that during the Paleo- gene thrusting the relative positions of the belts and tectonic units were completely reversed (Iturralde- Vinent, 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).


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, AC --

Arroyo Cangre Formation (equivalent to the San Cayetano Formation), ES = E1 Sfibalo Formation, J = Jagua Formation, F = Francisco Formation, SV = 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, PN = 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 maximum thickness (1500 m)

in the subsurface. These rocks are thinner in the

higher tectonic units of the Sierra de los Organos belt (300-600 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 l i thostrat igraphic s cheme

The lithostratigraphic scheme for the Guanigua- nico 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; Co-

biella-Reguera, 1996) accepted this scheme, although

with some minor changes. The results of a recent mi-

cropaleontological 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 fun-

damental for developing the modem stratigraphic

scheme for the Sierra de los Organos belt (see also

Khudoley and Meyerhoff, 1971). Herrera (1961)

98 A. PSZCZOLKOWSKI

P I N A R 1

m 0

1000 "-~,~'~.~'t

~< ,.,,< 2 0 0 0 '< ''< ''<

,~-.~-)

,'- ,>,3 ,,-<

300C - -

N M'-N'-M ~ - ~ - ~

~ - ~ - ~

~ - ~

5 0 0 0 - - ' ~ ~ ~

Lower Cretaceous (Aptian - Albian)

Lower Cretaceous (Valanginian)


Upper Jurassic

Upper Jurassic

Lower Eocene


Lower Cretaceous (Berriasian - ?Barremian)

Upper Jurassic Lower Cretaceous (Berriasian - Valanginian) Lower Cretaceous (Aptian - Albian)

Lower Cretaceous (Berriasian - ?Barremian)

Upper Jurassic

m m - - 5

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 -- Aptian- Albian 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 Fur- razola-Bermddez (1968), Wierzbowski (1976), My- czyfiski (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

EARLY

EOCENE

PALEOCENE

IVlAASTRICHTIAb

CAMPANIAN S~-E~-%5]q]-R~--

CONIACIAN TURONIAN

CENOMANIAI~

APTIAN-ALBIAI~

BARREMIAN

VALANGINIAN

BERRIASIAN

TITHONIAN

KIMMERIDGIAh

- 5

Sierra de los Southern Northem Organos belt Rosario belt Rosario belt

1 MANACAS FM. ;' Vieja Member i.. . . . . . . . . i . . . . . . . . . . . . . . . . . . . . . .

Pica Pica Member I

CACARAJiCARA FM.

PONS

FORMATION

CARMITA FM.

SANTA TERESA FM.

! Tumbitas Mb. I

< Z ~ ~u-r~i~aer~' ' u_m.i,:r~ Mb' :1 0 Member ~ ARTEMISA B

< ~- FE~i~,?rie-a-ca-n~ . . . . . . . q u) ~ : , Member < p . . . . . . . . . . . . . . . La Zarza I

~ i ', FORMATION D i San Vicente Member o o , Member ~;:, ',

i .... i

-L_ FM.

FRANCISCO FM.

" I SAN CAYETANO

FORMATION

Fig. 6. Lithostratigraphic scheme of the Sierra de los Organos and Rosario belts in the Guaniguanico terrane: L. = lower (Ox- fordian), U. = upper (Oxfordian), S.V. = San Vicente Member of the Artemisa Formation.

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 obser- vations 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 For- mation. 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 flu- vial 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


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 (Schwage- rina 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 Forma- tion, southeast of La Palma. These ammonites belong to the following taxa: Perisphinctes (?Dichoto- mosphinctes) cayetaensis Myczyfiski, 1976, P. (?Di- chotomosphinctes) cf. anconensis Sfinchez Roig, and P. (Discosphinctes) cf. pichardoi Chudoley et Fur- razola-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 trans- porting 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 A. PSZCZOLKOWSKI

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 Pa- leogene formations. Locally, the E1 S~ibalo Formation is overlain by thin San Cayetano Formation silici- clastics, but often contacts with the ?late Oxfordian- Kimmeridgian limestones of the Artemisa Formation. The E1S~ibalo/Artemisa boundary is tectonically dis- turbed 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 Forma- tion (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 visi- ble 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 (late- middle to late Oxfordian). Consequently, the age of the E1 S~ibalo Formation should be pre-late Oxfor- dian, 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 Forma- tion 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,

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, Os- trea, 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 Zac- arras 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 Fran- cisco Formation is an equivalent of the Jagua Forma- tion. The Francisco Formation consists of shales, mi-


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 Oxfor- dian and also the lower part of the upper Oxfordian (Kutek et al., 1976; Myczyfiski, 1976, 1994a).

Facies interpretation. The deposits of the Fran- cisco 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 Oxfordian- Valanginian)

Subdivision and lithology. The Guasasa Forma- tion 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

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 For- mation (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 ammonite- beating 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 For- mation 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 (Duran- gites, Corongoceras, Kossmatia), those known from

102 A. PSzCZOLKOWSKI

LU

!-- O9

Substage

z Upper <

z 0 -1- !-

~_ Lower

Ammonite zones

Durangites - Himalayites -

Hildoglochiceras (Salinites)

Paralytohoplites carribeanus Pseudolissoceras spp.

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 ant i l leanum (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 Oxfordian- Kimmeridgian 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 Oxfordian- Valanginian)

Subdivision and lithology. The Artemisa Forma- tion 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 Mem- ber (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.


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-Miro- sphinctes 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. Conse- quently, the lower part of the La Zarza Mem- ber 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 Paralytoho- plites, 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 In- oceramus (Parkinson, 1819), although representa- tives of Buchiidae are also present (Myczyfiski and Pszczdtkowski, 1994; Myczyfiski, 1994b).

The Saccocoma-Didemnidae microfacies pre- dominates 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).

The calpionellids document a late Berriasian- early 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 nannofos- sils (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 (Im- lay) 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 docu- mented) 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 (la- gunal?) 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 North- ern 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, APS = 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, BV = 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.


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 to- ward 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 Hauterivian- Barremian)

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 (?Hauterivian- Cenomanian)

Lithology and distribution. The Santa Teresa Formation is composed of green, red or reddish-

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. Radiolari- ans identified by Aiello and Chiari (1995) indicate a Valanginian to middle Aptian age of one sam- ple 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 verti- cally 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.


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 Forma- tion. 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 shallow- water 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

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 Turonian- Campanian 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 For- mation 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, hemi- pelagic 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.


Fig. 14. Pelagic limestones and cherts of the Pefias Formation (Campanian-Maastrichtian); the Las Piedras river, south of Pons, Sierra de los Organos belt. Hammer length 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, subordi- nately 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 Forma- tion 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, respec- tively, 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 Maas- trichtian. 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


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 depres- sions). The volcaniclastic debris locally appeared in the Southern Rosario belt beginning an early stage of the foreland basin development during the (late) Early Paleocene.

Manacas Formation (Paleocene-Lower Eocene)

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

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 char- acterize 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.

JURASSIC TO EARLY PALEOCENE EVOLUTION OF THE PASSIVE MARGIN IN WESTERN CUBA

Plate tectonic reconstructions of the Caribbean region

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 proto- Caribbean 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; K1- itgord 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 in- sertion of the Farallon plate between North and


~ooovv C a l l o v i a n ,oovv

l /t~ /-~",, I v

30ON

- - , '~ . B ==== --~ ,,

I / p," '7 I - - '

b*+*+*+1 % % % " ~ 1 1.+_,.+,+.,_.112 [. =.=,. .. eJ 3 I"' 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 Juras- sic to Paleogene deposits)" R = Remedios belt of north-central and northeastern Cuba (Jurassic?-Late Cretaceous)" SA = South America; TMV-- 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 proto- Caribbean crust.

The position of the Guaniguanico terrane in respect to the Yucatfin block during the Jurassic is presented herein (Figs. 15-17) mainly after Itur- ralde-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

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 Or- ganos) 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 sur- rounding 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 Bore- al/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 magneti- sation 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 Escam-

2 'Information and 1994 Annual Report', Institute of Geological Sciences, Polish Academy of Sciences, Warsaw, 1995, p. 15.


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 ray terrane. These samples y ie lded p a l e o m a g n e t i c

inc l ina t ion indica t ive of a pa leo la t i tude of about 12 ~

Pdrez L a z o et al. (1995) conc luded that their pale-

o m a g n e t i c resul t cor robora tes the ear l ier geo log ica l

in te rpre ta t ions conce rn ing the or iginal loca t ion of

the C u b a n Jurass ic rocks "not too far f rom Hondura s

and Gua temala . . . " (op. cit.), that is, sou thwes t of its

p re sen t -day posi t ion. Moreover , Pdrez L a z o et al.

(1995) inferred f rom their p a l e o m a g n e t i c data con-

ce rn ing the L o w e r Cre taceous 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

Cuba took p lace at pa leo la t i tude values of about

16-17~ Consequen t ly , the en t rance to the proto-

Car ibbean bas in cou ld be b locked by the Cuban or

the Grea te r Ant i l les 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 pa leo la t i tude


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 Bar- rett, 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 (Rosen- crantz, 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 configura- tion, 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 compati- ble with the Caribbean paleotectonic reconstructions proposed by Ross and Scotese (1988) and Pin- dell and Barrett (1990) for the Middle Jurassic- Oxfordian and Late Triassic-Early Jurassic, re- spectively. In general, the composition of the San Cayetano sandstones (Pszcz6tkowski, 1986a; Hut- son 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 Forma- tion. 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.

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 Kim- meridgian carbonate rocks of the Sierra de los Organos and Sierra del Rosario belts. In the Titho- nian 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 Is- lands (Myczyfiski, 1999). In the Sierra del Rosario, the Tithonian bivalves belong mainly to the genus Inoceramus (Parkinson, 1819), although represen- tatives 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 at- tempt to locate the Guaniguanico terrane at a higher paleolatitudinal position (30~ during the Titho- nian, 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, 800- 1000 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 with- out deep-water successions of their proto-Caribbean continental slope and adjacent basin floor. Consider- ing 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 left- lateral transform motion probably occurred between these two continental blocks (Marton and Buffler, 1994). Such a situation could exist also during the


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 con- sisted 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. Ac- cording 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 explana- tion 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 Mey- erhoff, 1971; Meyerhoff and Hatten, 1974; Iturralde- Vinent, 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 Yu- cat~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

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 si- multaneous 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 obvi- ous 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 domi- nant 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 Oxfordian- Santonian)

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 Oxfor- dian 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. Infre- quent ammonites have been found only in the uppermost San Cayetano clastics in the Southern Rosario


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 (South- ern 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 Kim- meridgian. 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 Oxfor- dian or earliest Kimmeridgian time. These local up- lifts (rotated fault blocks?) could be a barrier inhibit- ing 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 Kim- meridgian, 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 Vi- cente 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-

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 North- ern 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 Titho- nian 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 predom- inated 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 proto- Caribbean 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 Ca- majuanf 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, oxy- genated, 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 expla- nation 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. Basin- ward, these thick-bedded, pure pelagic limestones passed gradually into black, thin-bedded radiolar-


ian limestones occurring in the Northern Rosario belt. The ammonites are uncommon in the pelagic limestones of Berriasian and Valanginian age (My- czyfiski, 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 Valanginian- Barremian. A similar petrographic composition of Hauterivian-Barremian siliciturbidites occurs in the Northern Rosario belt, La Esperanza belt and Plac- etas 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 differ- ences 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 hypo- thetical 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 disper- sal. 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.


level rise, paleogeographic and paleoceanographic conditions).

During the Cenomanian, the pelagic carbonate sedimentation was restored in the Rosario and Plac- etas 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 Creta- ceous 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; Pin- dell, 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 proto- Caribbean ocean basin has been progressively sub- ducted beneath the Caribbean plate during the Late Cretaceous-Eocene. An arc terrane collided with the southern Yucatan margin during the Campa- nian 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

Fig. 19. Generalized paleogeographic reconstruction of the proto- Caribbean 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-water limestones 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 Pin- dell 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 con- nected 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 Maas- trichtian (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 North- ern 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


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 At- lantic 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 calciru- dites 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 Jurassic- Cretaceous 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 syn- rift stage (Lower Jurassic-?Callovian/early Oxfor- dian), 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 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 ammonite- bearing 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 shallow- water 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


Campanian , the volcanic arc was located east of

the Yucatfin b lock margin and south of the Moreno

depocenter . This arc could be the wes te rnmos t part

of the Grea ter Ant i l les Arc (GAA), as proposed by

Pindel l and Dewey (1982) and Pindel l et al. (1988).

The late Maas t r icht ian deposits of the passive mar-

gin of Yucatfin are represented by the Cacarajfcara

Format ion in the Rosar io belts. A more souther ly

posi t ion of the ext inct volcanic arc at the end of

the Cre taceous is inferred f rom the paleotec tonic re-

const ruct ion and l i thology of the late Maas t r icht ian

deposits.

Dur ing the Paleocene, the sed imentary succes-

sions of the Guan iguan ico terrane, or iginal ly de-

posi ted on (and along) the passive margin of Yu-

catfin, fo rmed the foreland basin substrate. This fore-

land basin was located in front of a thrust belt a long

the southern side of the remnant pro to-Car ibbean

Sea.

ACKNOWLEDGEMENTS

The author is grateful to Richard T. Buffler,

Thomas W. Donnel ly , John E Lewis and Gy6rgy

Mar ton for their review of the manuscr ipt , and to

Paul M a n n for his useful c o m m e n t s on the text and

figures. The discussions with Ryszard Myczyfiski

and Manue l I turra lde-Vinent on some problems con-

cerning the geo logy of Cuba are appreciated.

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