Idea Transcript
Facies, genetic evolution and paleoecological control of mid-Cretaceous carbonate system in the northern Levant Margins: northern Israel
Ran Frank
This work was submitted for the degree "Doctor of Philosophy" to the Senate of Ben-Gurion University of the Negev, Beer Sheva. The study was carried out under the supervision of: Prof. Chaim Benjamini, Ben-Gurion University of the Negev, Beer Sheva. Dr. Binyamin Buchbinder, Geological Survey of Israel, Jerusalem.
Report GSI /03/2010
Jerusalem, February 2010
Acknowledgements First and foremost I want to express my thanks to Y. Rafael of the Geological Survey of Israel. Without his clear judgment and immediate action in the field on 28 July 2005, this work would have never been completed. I wish to thank my mentors and research supervisors, Prof. Chaim Benjamini and Dr. Binyamin Buchbinder, for opening the window to the new world of the Judea Group, for the opportunity to embark on new approaches outside the envelope of conventional wisdom, for the willingness to critically discuss new ideas, taking nothing at face value, in the field and laboratory. Prof. Yehuda Eyal, supervisor of my M.Sc. thesis, continued to cooperate in the course of this research by joining me in the field, teaching me to identify and measure mesostructures, and offering constructive criticism and advice. Thanks are due to research and technical staff at the Geological Survey of Israel, especially Dr. Ahuva Almogi-Labin, whose continuous interest and assistance and facilitation of use of services of the GSI were crucial. Ms. Shulamit Lipson-Benitah assisted in accessing material from the Carmel, and identification of microfauna. Dr. Zeev Lewy assisted in macrofaunal identification and Dr. Ran Calvo provided orthophotos. Thanks are due to Dr. Mike Simmons of NEFTEX, for his immediate assistance in determining orbitolinid faunas. At the Department of Geological and Environmental Sciences at BGU, thanks are due to Eli Shimshishvili for technical support in microscopic facilities, David Kusashvili for thin section preparation, Rivka Eini and Zahala Sharabi for all logistical and bureaucratic support, and to my colleagues Oded Bar-Farago and Zafrir Levy for their interest during the course of this research and in my results. I extend my appreciation for the crucial support of staff and colleagues at BGU and the Geological Survey of Israel, during my stay at Rambam hospital in summer 2005, by visiting, telephoning, and taking interest in my condition, making a major contribution to my rapid recovery. Finally, to the Frank and Wener extended family, who contributed in looking after the children helping to bring this work to conclusion. And finally to my wife Ilana and to Nadav, Gilad and Yoav that accompanied me during the course of this work. This research was supported by research grants 25-17-037, 25-17-042 and 27-17-014 of the Earth Sciences Bureau, Ministry of National Infrastructures, by the Geological Survey of Israel, and by the Department of Geological and Environmental Sciences, Ben Gurion University of the Negev.
Table of contents Section
Contents
Page
Chapter 1: Introduction……………………………………………………………….
1
1.1.
Conceptual background and the aim of this study……………………
1
1.2.
Geological background.…………………………………………………….
1
1.2.1.
Tectonic setting of the Levant margin in the Cretaceous…………………
1
1.2.2.
Cenomanian–Turonian depositional units and lithostratigraphy in northern Israel…………………………………………………………………
2
1.2.3.
Previous concepts on the Judea Group in northern Israel……………….
3
1.3.
Critique and research questions………………………………………….
5
1.4.
Methods and procedures…………………………………………………..
7
1.4.1.
Field and Laboratory procedures……………………………………………
7
1.4.2.
Stratigraphic procedures……………………………………………………..
7
1.4.3.
Methods and concepts………………………………………………………..
7
1.5.
Structure of the thesis………………………………………………………
9
Figures 1.1 – 1.2……………………………………………………………..
10
Chapter 2: Carbonate facies, depositional environments, and cyclic patterns in the Cenomanian-Turonian succession in northern Israel: ……
13
2.1.
Introduction…………………………………………………………………...
13
2.2.
Representative columnar sections, their facies types and construction of subunits……………………………………………………
13
2.2.1.
The Deir El-Assad section, western Galilee………………………………..
13
2.2.1.1.
Subunit DS1 (Sakhnin Fm)…………………………………………………..
13
2.2.1.2.
Subunit DS2 (Bina Fm)……………………………………………………….
14
2.2.2.
Yanuch section, western Galilee…………………………………………….
14
2.2.2.1.
Subunit YN1 (lower Yanuch Fm)…………………………………………….
14
2.2.2.2.
Subunit YN2 (Yanuch Fm)……………………………………………………
14
2.2.2.3.
Subunit YN3 (Yanuch Fm)……………………………………………………
15
2.2.2.4.
Subunit YN4 (top Yanuch, lower Yirka Fm)………………………………...
15
2.2.3.
Bet-Ha’Emek section, western Galilee……..………………………………
16
2.2.3.1.
Subunits BK1 and BK2……………..………………………………………..
16
2.2.3.2.
Subunit BK3 (Kishk Fm)………………………………………………………
16
2.2.4.
Betzet section, northwest Galilee……………………………………………
17
2.2.4.1.
Subunit BZ1 (lower to mid- Deir Hanna Fm)……………………………….
17
2.2.4.2.
Subunit BZ2 (upper Deir-Hanna and Sakhnin Fms)................................
17
2.2.4.3.
Subunit BZ3 (lower Bina Fm)………………………………………………...
18
2.2.4.4.
Subunit BZ4 (Bina Fm)………………………………………………………..
18
2.2.4.5.
Subunit BZ5 (Bina Fm)………………………………………………………..
18
2.2.5.
Dishon section, northeast Galilee……………………………………………
19
2.2.5.1.
Subunit DO1 (Sakhnin Fm)…………………………………………………..
19
2.2.5.2.
Subunit DO2 (Bina Fm)……………………………………………………….
19
2.2.5.3.
Subunit DO3 (Bina Fm)……………………………………………………….
19
2.2.5.4.
Subunit DO4 (Bina Fm)……………………………………………………….
19
2.2.5.5.
Subunit DO5 (upper Bina Fm).................................................................
20
2.2.6.
Manara section, northeast Galilee…………..……………………………...
20
2.2.6.1.
Subunit MN1 (lower Deir Hanna Fm)………………………………………..
20
2.2.6.2.
Subunit MN2 (Deir Hanna – Sakhnin Fms)..............................................
21
2.2.6.3.
Subunit MN3 (Yanuch Fm)…………………………………………………...
21
2.3.
Synthesis of sedimentary facies types and their environments of deposition………………………………………………………....................
21
2.4.
General outlines of cyclicity in northern Israel…………………………
25
2.5.
Discussion…………………………………………………………………….
26
2.5.1.
Low-order cycles and their paleoecological significances………………..
26
2.5.1.1.
Type-1 Low-order Composite Cycle (CC1)…………………………………
26
2.5.1.2.
Type-2 Low-order Composite Cycle (CC2)…………………………………
26
2.5.1.3.
Type-3 Low-order Composite Cycle (CC3)…………………………………
27
2.5.1.4.
Type-4 Low-order Composite Cycle (CC4)…………………………………
27
2.5.1.5.
Type-5 Low-order Composite Cycle (CC5)…………………………………
28
2.5.1.6.
Type-6 Low-order Composite Cycle (CC6)…………………………………
28
2.5.1.7.
Type-7 Low-order Composite Cycle (CC7)…………………………………
28
2.5.1.8.
Type-8 Low-order Composite Cycle (CC8)…………………………………
28
2.6.
Summary and conclusions…………………………………………………
29
Figures 2.1 – 2.11 ……………………………………………………………
30
Chapter 3: Sequence stratigraphy, facies evolution and tectonosedimentary configuration of the Cenomanian-Turonian carbonate system of northern Israel ………………………………………………………………………
41
3.1.
Introduction…………………………………………………………………...
41
3.2.
Sequence stratigraphy and sedimentary facies configuration………
41
3.2.1.
Sequence 1: Early-Mid Cenomanian………………………………………..
42
3.2.1.1.
Albian–Cenomanian sequence boundary (Alb/Ce SB-1)…………………
42
3.2.1.2.
The Early Cenomanian TST (Ce TST-1)……………………………………
42
3.2.1.3.
Early Cenomanian maximum-flooding interval (Ce MFI-1)……………….
43
3.2.1.4.
The Early-Mid Cenomanian RST (Ce RST-1)……………………………...
44
3.2.1.5.
The Mid-Cenomanian sequence boundary (Ce SB-2)……………………
46
3.2.2.
Sequence 2: late Mid - to Late Cenomanian……………………………….
47
3.2.2.1.
The Mid-Cenomanian lowstand system tract (Ce LST-2)………………..
47
3.2.2.2.
The late Mid- to Late Cenomanian transgressive system tract (Ce TST2) and the Late Cenomanian maximum-flooding interval (Ce MFI-2)….
47
3.2.2.3.
Late Cenomanian regressive system-tract (Ce RST-2)………………….
48
3.2.3.
Sequence 3: Late Cenomanian……………………………………………
49
3.2.3.1.
Late Cenomanian sequence boundary (Ce SB-3) and Pelech sequence
49
3.2.3.2.
The latest Cenomanian sequence boundary (Ce SB-4)…………………
50
3.2.4.
Sequence 4: Early –Middle Turonian………………………………………
51
3.2.4.1.
Turonian transgressive system tract (Tu TST)……………………………..
51
3.3.
Discussion…………………………………………………………………….
52
3.3.1.
The mid-Cretaceous tectono-sedimentary framework of northern Israel
3.3.2.
Controlling mechanisms and correlation across the Arabian Plate and
52
beyond………………………………………………………………………..
53
3.3.2.1.
The beginning of Cenomanian sequence-1………………………………
53
3.3.2.2.
End of Cenomanian sequence-1………………........................................
54
3.3.2.3.
The beginning of Cenomanian sequence-2………………………………..
55
3.3.2.4.
Latter part of Cenomanian sequence-2…………………………………….
56
3.3.2.5.
Sedimentary events across the latest Cenomanian sequence boundary
3.4.
and the manifestation of the OAE-2 ……………………………………….
58
Summary……………………………………………………………………….
59
Figures 3.1 – 3.10 ……………………………………………………………
61
Chapter 4: Deterioration of Mid-Cenomanian ramp system in northern Israel: mechanical collapse, subaerial exposure, and submarine omission
71
4.1.
Introduction…………………………………………………………………...
71
4.2.
Mass-transport features of the Mid-Cenomanian FRST-1: ramp termination by mechanical collapse prior to the Ce SB-2 …………...
4.2.1.
71
The Namer Valley outcrops, NW Galilee: shear zones, translational slides and debrites…………………………………………………………….
72
4.2.2.
Adamit region, NW Galilee: A channel levee complex…………………..
73
4.2.3.
Yanuch region, western Galilee: slide scar and turbidites……………….
73
4.2.4.
The Betzet mega-block, NW Galilee……………………………………….
74
4.2.5.
Model for mass-transport across the distally steepened slope………….
4.3. 4.4.
Diagenetic features of the Mid-Cenomanian Sequence Boundary (Ce SB-2) and transgressive surface ……………………………………. Discussion…………………………………………………………………….
4.4.1.
Mechanism of slope collapse………………………………………….…….
76
4.4.2.
Mid-Cenomanian ramp termination: a composite discontinuity…….…..
77
4.5.
Summary……………………………………………………………………….
78
Figures 4.1 – 4.5 ……………………………………………………………..
79
74 75 76
Chapter 5: Fault-controlled calcarenitic shelf-margin apron across the Cenomanian-Turonian boundary, western Galilee, northern Israel ……….
85
5.1.
Introduction…………………………………………………………………...
85
5.2.
Late Cenomanian clinoform body in the western Galilee…………….
86
5.2.1.
Distribution and characteristics of the clinoform body……………………
86
5.2.2.
Genesis of clinoform features………………………………………………..
86
5.3.
Special features of the clinoform body of the western Galilee……...
87
5.3.1.
Significance of the Yirka Fault…………………………………………….....
87
5.3.2.
Special features of the Kishor-Blaya and Hamra clino-successions…...
87
5.3.2.1.
Description of special featutes at Kishor-Blaya and Hamra…………….
87
5.3.2.2.
Late Cenomanian events at Kishor-Blaya and Hamra…………………..
88
5.3.3.
Special features of the Hamra Valley outcrops…………………………..
88
5.3.3.1.
Description of special features at Hamra Valley…………………………
88
5.3.3.2.
Late Cenomanian events at the Hamra Valley……………………………
89
5.3.4.
Special features of the Mt. Gamal outcrops………………………………..
89
5.3.4.1.
Description of special features at Mt. Gamal region………………………
89
5.3.4.2.
Late Cenomanian events at the Mt. Gamal region……………………….
90
5.3.5.
Late Cenomanian turbidites, debrites and shear-zones in the western Carmel region………………………………………………………………….
90
5.3.5.1.
Megadim- and Oren valley outcrops……………………………………….
91
5.3.5.2.
Late Cenomanian events at Megadim and Oren valleys……………….
91
5.3.5.3.
The Sefunim quarry section…………………………………………………
91
5.3.5.4.
Late Cenomanian events at Sefunim………………………………………
92
5.3.5.5.
The Isfiyye section…………………………………………………………….
92
5.3.5.6.
Late Cenomanian events at Isfiyye………………………………………….
92
5.4.
Discussion…………………………………………………………………….
93
5.4.1.
The Late Cenomanian structural-depositional slope configuration of
5.4.2.
northern Israel………………………………………………………………….
93
Evolution of the Late Cenomanian slope ─ toe-of-slope system of N. Israel…………………………………………………………………………….
93
5.4.2.1.
Stages 1 and 2: Formation of open margins……………………………….
5.4.2.2.
Stage 3: Development of slope-apron foresets and base-of-slope
93
turbiditic bottomsets…………………………………………………………..
93
5.4.2.3.
Stage-4: Sea level fall, followed by drowning of the slope-apron……….
94
5.4.2.4.
Stage-5: tectonic uplift and faulting………………………………………..
95
5.4.2.5.
Stage-6: Termination by Drowning…………………………………………
95
5.4.3.
Comparison of the Galilean carbonate shelf-margin apron with siliciclastic deltaic systems……………………………………………………
95
5.4.4.
'Fill and Spill' sedimentary dynamics………………………………………..
97
5.5.
Summary……………………………………………………………………….
97
Figures 5.1 – 5.11 ……………………………………………………………
99
Chapter 6: Sedimentary configuration and tectonic framework of the midCretaceous northern Arabian platform, and their paleoceanographic implications ……………………………………………………………………………
111
6.1.
Introduction…………………………………………………………………...
111
6.2.
Thickness patterns of the Cenomanian succession in the northern Levant ………………………………………………………………………….
112
6.3.
Sedimentary configuration in the Ce TST-1 …………………………….
113
6.3.1.
Characteristics of Ce TST-1 …………………………………………………
113
6.3.2.
SW-NE traverse: Ce TST-1 along the Carmel–Palmyride trend…………
113
6.3.3.
S-N Traverse: Ce TST-1 in the Carmel – Galilee – central Lebanon NW Syria trend………………………………………………………………
114
6.3.4.
Ce TST-1 basin configuration in the northern Levant …………………….
115
6.4.
Sedimentary configuration in the Ce RST-1 ……………………………
115
6.4.1.
Characteristics of Ce RST-1 …………………………………………………
115
6.4.2.
SW-NE traverse: Ce RST-1 along the Carmel –Palmyride trend ……….
115
6.4.3.
S-N Traverse: Ce RST1 along the Carmel – Galilee – central Lebanon – western Syria trend………………………………………………………….
116
6.4.4.
Ce RST1 basin configuration in the northern Levant…………………….
117
6.5.
Sedimentary configuration in the Ce TST-2 …………………………….
117
6.5.1.
Characteristics of Ce TST-2 …………………………………………………
117
6.5.2.
SW-NE traverse: Ce TST-2 along the Carmel –Palmyride trend……….
117
6.5.3.
S-N traverse: Ce TST-2 in the Carmel – Galilee – central Lebanon – western Syria trend …………………………………………………………...
118
6.5.4.
Ce TST-2 basin configuration in the northern Levant……………………..
119
6.6.
Sedimentary configuration in the Ce RST-2 ……………………………
119
6.6.1.
Characteristics of Ce RST-2 …………………………………………………
119
6.6.2.
S-N traverse: Ce RST-2 in the Carmel – Galilee – Central Lebanon trend
120
6.6.3.
SW-NE traverse: Ce TST-2 in the Carmel – Palmyride trend……………
120
6.6.4.
Ce RST-2 basin configuration in the northern Levant…………………….
120
6.7.
Sedimentary configuration in the Turonian TST ………………………
121
6.7.1.
Geometry of Turonian TST in northern Israel………………………………
121
6.7.2.
The Early Turonian succession of northern Israel and northern Levant – earlier model…………………………………………………………………
121
6.7.3.
Basin configuration in the Turonian TST in the northern Levant……….
122
6.8.
Discussion…………………………………………………………………….
122
6.8.1.
Basin configuration: synthesis………………………………………………
122
6.8.1.1.
The continuation of the ‘Sinai-Carmel hinge-belt' north of the southern Carmel ………………………………………………………………………….
6.8.1.2.
122
The elevated margins of the Carmel–SW Palmyrides basin and their trends …………………………………………………………………………..
123
6.8.1.3.
Trend of the elevated margins of the northern basin …………………….
124
6.8.1.4.
Configuration of the Cenomanian-Turonian carbonate system in the northern Levant………………………………………………………………..
124
6.8.2.
Paleoceanographic implications of the tectono-sedimentary framework.
125
6.8.2.1.
Dynamics of the oceanic OMZ in the Cenomanian-Turonian of the northern Levant ……………………………………………………………….
125
6.8.2.2.
OMZ dynamics and Cenomanian drowning events……………….………
127
6.8.3.
Controls on systems tract geometry and facies……………………………
128
6.9.
Summary and conclusions…………………………………………………
130
Figures 6.1 – 6.10 ……………………………………………………………
131
Chapter 7: Synthesis and summary ……………………………………………...
141
References……………………………………………………………………………
143
Appendix 1: Photos……………………………………………………………….. Appendix 2: Database of columnar sections (CD)
159
Abstract This study addresses the genesis of the mid-Cretaceous carbonate system of northern Israel and the northwest Arabian margin, a key component of the Cretaceous Tethyan paleoceanographic system. Mapping units of northern Israel were deconstructed and reassembled into genetic units comprising autochthonous and allochthonous basinal facies, a variety of outer- and mid-ramp facies types, shallow-marine inner ramp facies and features indicating subaerially-exposed environments. Discrete facies units comprise two orders of progradational and retrogradational cycles. Low order progradational cycles reflect contrasting styles of accommodation space filling: An efficient carbonate factory, filling accommodation space to sea-level and forming a distally-steepened ramp; a poorly productive carbonate system that accumulated to wave base only, due to excess nutrients, producing a homoclinal ramp; and a low productivity ramp that aggraded to sea level with minimal in-situ skeletal production but with increased grain supply by transport. The mid-Cretaceous carbonate system forms three Cenomanian sequences and a single Turonian sequence, divided into systems tracts. Eustatic and Palaeoenvironmental imprints are represented by earliest Cenomanian subaerial exposure; Early Cenomanian maximum flooding and oxygenation of hypoxic sea-floor; Mid-Cenomanian highstand progradation followed by forced-regression and masstransport; Mid-Cenomanian subaerial exposure; Late Cenomanian eutrophication during
sea-level
rise;
Late
Cenomanian
subaerial
exposure;
latest
Cenomanian/Turonian eutrophication and gradual development of the OAE-2 (oceanic anoxic event). Late Cenomanian eustatic rise was locally masked by uplift and subaerial exposure. Mid-Cenomanian distally-steepened ramp system was first terminated by mechanical collapse represented by a mass-transport complex, subjected to subaerial exposure, and finally flooded. Mechanical collapse of the ramp took place as a result of forced regression and by overloading caused by the transition from progradational to aggradational regime. The Late Cenomanian shelf-edge was characterized by a transformation from homoclinal ramp to steep slope, involving rotational faulting simultaneous with Late Cenomanian eustatic fall. Resulting sediment instability upon subsidiary blocks triggered mass-movements. Syn-deformational fill on the slope formed coarse-grained clino-successions, thicker and steeper toward the south. This system is represented in the Carmel region by a turbidite sheet, debrites, shear-zones and translational slides.
Expansion and contraction of the OMZ on the northern Arabian platform was traced from the beginning of the Cenomanian. The Late Cenomanian-Early Turonian systems tract reflects deterioration of the aggradational ramp by increasing ecological stress, culminating with eutrophication and hypoxia related to the Tethyan OAE -2. The S-N depositional strike of the mid-Cretaceous Levant margin, the ‘SinaiCarmel hinge belt’, is interrupted north of the southern Carmel. An E-NE trending ‘southern-Carmel_Rutbah-Uplift paleohigh’ in the south was bounded to the north by a parallel trough, the ‘Carmel_SW Palmyride basin’. To the north lies the ‘Galilee–NE Palmyride paleohigh’, descending to a broad subsiding region extending into Lebanon and western Syria. Keywords: Mid-Cretaceous, Carbonate system, Northern Israel, Arabian Platform.
Chapter 1: Introduction
Chapter 1 Introduction 1.1. Conceptual background and the aim of this study Marine carbonate systems are strongly dependent on auto-production by skeletonbearing organisms, and therefore on their ecological requirements and on evolutionary trends. Carbonate factories vary in their nature (e.g., tropical, cool-water/foramol-type, mudmounds; Carannante et al. 1997; Schlager 2000), production capacity, facies-belt geometry and depositional profile (e.g. Burchette & Wright 1992), and in their reactions to climatic, paleoceanographic and tectonic regimes (e.g. greenhouse vs. icehouse conditions; nutrient flux; sea-level change, eustatic or otherwise; basin subsidence). These and other factors determine the capacity of a system to fill accommodation space and ultimately determine carbonate facies and sequence geometries. Therefore, understanding of the variety of factors underlying development of a particular carbonate system is the key to unlocking the geological history of carbonate-dominated continental margins. This study addresses the genesis of the Cenomanian-Turonian carbonate system of northern Israel, in the light of these factors and the particularities of its development on the northwest Arabian margin, an important part of the Cretaceous Tethyan paleoceanographic system.
1.2. Geological background 1.2.1. Tectonic setting of the Levant margin in the Cretaceous The Levantine margin of the Neotethys (Fig. 1.1) was shaped by Late Permian, Middle to Late Triassic, and Early Jurassic extensional tectonics. Rifting led to separation of the Tauride and Eratosthenes blocks from the Arabo-Nubian Platform, forming the Levantine Basin of the eastern Mediterranean (e.g., Bein & Gvirtzman 1977; Garfunkel & Derin 1984; Garfunkel 1998; 2004; Robertson 1998). Normal faulting associated with this rifting determined the depositional strike along the Levant margin, which constitutes the eastern margin of the Levantine Basin. The north-south trend exercised control on facies transitions along a narrow depositional belt, termed the 'Levantine hinge-line’ or –‘belt' (Gvirtzman & Klang 1972; Bein & Gvirtzman 1977). Cessation of faulting in the Mid-Jurassic, lengthy tectonic quiescence, and slow passive-margin type subsidence prevailed from the MidJurassic to the Middle Cretaceous. During this period of quiescence, E-W proximal-to-distal facies transitions prevailed, normal to the north-south striking shelf-margin. Mid-Cretaceous features of this type were described from the subsurface of the coastal plain of Israel as far
1
Chapter 1: Introduction
north as the Carmel region (Bein 1971, 1974; Bein & Weiler 1976; Sass & Bein 1982), and again farther to the north, along the S-N trend of the western Lebanon flexure (Saint-Marc 1974; Walley 1998) (Fig. 1.1). No such trend was shown to be present in the Galilee, northern Israel. 1.2.2. Cenomanian–Turonian depositional units and lithostratigraphy in northern Israel The Judea Group of Israel is a Middle Cretaceous carbonate succession composed of up to 900m of limestones, chalks, dolomites and marls in Central Israel, thinning to 340m in the Eilat region, and thinning again towards the coast and offshore (Arkin et al. 1965; Arkin & Hamaoui 1967). Pioneering studies placed this succession in the Cretaceous and established presence of rudists as major biotic contributors. The Judea Group in the Galilee begins, according to Kafri (1972), with the Kesulot Fm (Formation) (Weiler 1968), overlain by the Yagur Dolomite (Picard & Kashai 1958), the Deir-Hanna Fm (Golani 1961), the Sakhnin Fm (Golani 1961), and the overlying Bina Fm (Shadmon 1959). Westwards, the Judea Group passes from dolomites, limestones, and marls in outcrops, to marls and chalks of the TalmeYafe Fm (Derin & Gerry 1965; Cohen 1971) in the subsurface of the coastal plain and offshore. In northern Israel, the Judea Group was divided in the Carmel region into lithostratigraphic units established by Picard & Kashai (1958) and Kashai (1958), and somewhat differently by Vroman (1960). Detailed descriptions of these units are by Kashai (1966). Figure 1.2 represents the lithostratigraphy currently used for the upper Judea Group carbonates (Cenomanian-Turonian) of northern Israel. It incorporates schemes used by Picard & Kashai (1958), Freund (1959), Kashai (1966), Kafri (1972, 1991), Sneh et al. (1998), Sneh (2002), and Segev & Sass (2006) with some local modifications. There has been a tendency to use different lithostratigraphic concepts for the Galilee and the Carmel, and the schemes are presented here separately. In the Galilee, limestones of the Deir Hanna Formation occur at the base of the upper Judea Group, overlying Late Albian Yagur or Kamon dolomites (Fig. 1.2). The Deir Hanna Formation consists of well bedded or laminar chalks and limestones with chert nodules or bands, sometimes dolomitized near the top. Sakhnin Formation dolomites directly overlie the Deir Hanna Formation in most of the Galilee. The Sakhnin dolomites become well-bedded upwards, or are sometimes highly brecciated (e.g. in the regions of Adamit, Peqi’in, Deir ElAssad, Fig. 1.1). In some localities, breccias of the Sakhnin Formation pass laterally into dolomitized laminites of the upper Deir Hanna Formation (e.g., Adamit, Betzet, Fig. 1.1). The Sakhnin Formation was not mapped in the western Galilee (e.g. Picard & Kashai 1958; Freund 1965) where bioturbated limestones of the Yanuch Formation (Freund 1959)
2
Chapter 1: Introduction
unconformably overlie chalky laminites of the Deir Hanna Formation. Kafri (1972) showed that the bioturbated limestones of the Yanuch Formation in the western Galilee pass into chalks and bedded limestones toward the northwest Galilee (e.g. Hila and Hurfesh regions). Typical lithologies, microfacies, and cyclic patterns of the Yanuch Formation also occur in the northeast Galilee (e.g. Dishon, Manara sections). From the western to central Galilee, the Yanuch Formation is overlain by ammonite-bearing marls of the Early Turonian Yirka Formation. In other parts of the Galilee, a limestone complex mapped as Bina Formation (Shadmon 1959) directly overlies either the Yanuch or Sakhnin Formations. In the Carmel, the Yagur Formation dolomites of Late Albian age (Fig. 1.2) form the lower part of the Judea Group. The Yagur Formation is overlain by the Isfiyye Formation, consisting of well-bedded chalks with chert nodules and bands, and in some places with pyroclastics interbedded at the base (e.g., Sass 1980; Segev & Sass 2006). The Isfiyye Formation is overlain by bioclastic limestones of the Beit-Oren Formation. The Beit-Oren limestones are overlain by chalks with some pyroclastics of the Arkan Formation (Segev et al. 2002; mapped as Khureibe and Junediyye Formations in older studies). The Arkan Formation is overlain by dolomites and limestones of the Muhraqa Formation. This transition is unconformable (Bein 1974; Lipson-Benitah et al. 1997), in places accompanied by pyroclastics (Segev & Sass 2006). Above the Muhraqa Formation are Early Turonian ammonite-bearing marls, chalks and limestones of the Daliyya Formation in the centralnorthern Carmel (Freund & Raab 1969). Limestones of the Bina Formation form the top of the succession. In the region of southern Carmel, the Cenomanian succession is only partly exposed and is mostly dolomitic. At the base a dolomitized chert-bearing equivalent of the Isfiyye Formation, is overlain by bioturbated dolomites of the Zikhron Formation, becoming well bedded toward the top. Dolomites of the Zikhron Formation are considered by Segev & Sass (2006) as the lateral continuation of the Arkan chalk complex to the north. The Zikhron dolomites are overlain by a pyroclastic bed (Segev & Sass 2006) followed by massive dolomites and limestones mapped as Sakhnin and Bina Formations respectively. These lithostratigraphic units for the most part are useful for depiction on geological maps. In order to determine the sedimentary evolution of this region, however, it was necessary to deconstruct them into their basic genetic components, and to reconstruct their cyclic patterns, systems tracts and sequences, using biostratigraphic and other chronostratigraphic controls. 1.2.3. Previous concepts on the Judea Group in northern Israel Some workers proposed sedimentologically-based paleogeographic models for different parts of the Judea Group in northern Israel, and also general models and concepts.
3
Chapter 1: Introduction
Sass & Bein (1982) summarized the prevailing concept for the paleogeography of the Judea Group. They posited a wide, shallow, relatively uniform “shelf – lagoon” represented by dolomitic lithofacies of the Judea Group, with local low-relief basins, filled by limy, chalky or marly lithofacies (e.g., Freund 1978). The “shelf – lagoon” platform was laterally and abruptly transitional to basinal Talme-Yafe lithofacies in the west (Bein & Weiler 1976) via an N-S striking rimmed shelf-edge, dominated by rudist “reefs” and patch reefs which are exposed in the Carmel and Galilee (Bein 1976; Freund 1978; Sass & Bein 1982). Braun & Hirsch (1994) similarly suggested that the upper Albian of Israel was dominated by a wide and shallow shelf-lagoon platform, intra- and supratidal in the Negev and Jordan but subtidal in central Israel. Rudist reefs and buildups separated this platform from the basinal facies of the Talme-Yafe Fm. Vroman (1958) suggested that the appearance of the ammonite Leoniceras can be used to mark the Cenomanian-Turonian boundary in two northern basins of the western Galilee and Carmel. Freund (1962) elaborated on this approach, showing that the distribution of ammonites within marly facies in the Early Turonian, indicate a series of parallel SW-NE elongate basins extending into Egypt, Transjordan, Syria and Lebanon. In the Yirka – Peqi’in region of the western Galilee, he described a narrow SSW-NNE trending trough (Yirka basin) in the Late Cenomanian – Turonian (Freund 1958, 1965). This was considered to be synclinal, extending from the Carmel region (Daliya Fm), via western Galilee (Yanuch and Yirka Fms) into Lebanon. Rudist reefs of the Yanuch and Kishk Fms were developed on the flanks of this trough. In the Cenomanian of the northwestern Galilee, Kafri (1972) interpreted lithofacies changes also as forming alternating basins and swells, however following significantly different trends. Limy/chalky lithofacies of the Deir-Hanna and Yanuch Fms filled basins and are laterally transitional to the dolomitic Sakhnin Fm formed on swells. Levitte & Sneh (2002) and Sneh (2002) failed to document these lateral transitions between the Deir-Hanna and Sakhnin Fms in the Galilee. In the Carmel region, Bein (1974 1976) showed lateral variations from shallow-water dolomites, via “reef” complex, into deeper slope and basin lithofacies, the latter with calciturbidites. Sass & Bein (1982) divided the Judea Group in the Carmel into three depositional cycles, each showing similar facies variations and terminating with regression, viz: Albian – Early Cenomanian; Late Cenomanian; and latest Cenomanian – Turonian. Bein (1977) showed a narrow basin complex trending E-W in the Turonian Daliyya Fm of the Carmel. On the flanks, shallow-water deposits edged with rudists “reefs”, with pelagic muds interfingering. In the channel, allochthonous calcarenitic grainstones were deposited. This depression was considered an erosional channel, cut into the shallow platform during regression. The channel was compared with the narrow NE-SW Yirka basin of the western
4
Chapter 1: Introduction
Galilee (Freund 1962, 1965). Freund (1962) suggested that the two depressions were actually separated basins because they show different orientations. Honigstein et al. (1989) subdivided the Early Turonian Daliyya marls of the Carmel into the Galame sub-basin to the south, affected by anoxic conditions, and the Daliyat-el-Carmel sub-basin to the north, somewhat shallower and more oxygenated. The higher content of organic matter in the more anoxic sub-basin was attributed to high primary production, enhanced by coastal upwelling and local volcanism, which led to expansion of the subsurface oceanic oxygen-depleted layer into the relatively deeper water column of the Jalame sub-basin. They considered the anoxic conditions to be related to the second global Oceanic Anoxic Event (OAE-2) of Late Cenomanian – Early Turonian times. Buchbinder et al. (2000), however, suggested that upwelling alone sufficiently explains organic-rich intervals in the Daliyya Fm. Bein & Weiler (1976) considered the basinal Talme-Yafe Fm west of the ‘hinge-line’ (Gvirtzman & Klang 1972), and outcropping in the Carmel and Galilee, to be a huge prism of allochthonous platformal deposits accumulated on the continental rise and slope during a continuous transgressive phase beginning in the Early Albian. The deposits were transported mainly by bottom currents and by turbidity currents, and their final geometry was shaped by contour-parallel currents. Lipson-Benitah et al. (1997) showed that the top of the Yagur Fm in the Carmel represents transformation of shallow and warm inner-platform, to a new cycle of outer-shelf facies. This cycle shows variations in total and relative abundance of planktonic foraminifera representing variations in water-depths. Buchbinder et al. (1997) first mentioned the concept of drowning unconformities in the Judean group for the Albian-Cenomanian boundary and for the Late Coniacian. Buchbinder et al. (2000) proposed a sequence-stratigraphic framework for the Late Cenomanian– Turonian succession in Israel and northern Israel as well. The regional unconformity of the Late Cenomanian Neolobites vibrayeanus Zone of Israel was related to drowning of the Cenomanian carbonate platform due to eutrophication and anoxia.
1.3 Critique and research questions The uniformitarian approach (‘the present is the key for the past’) to carbonate systems was stimulated in the mid-20th century by principles and models derived from studies of modern-day coral-reefs, including Holocene Bahamian-type and Florida-shelf case studies. Such concepts were behind pioneering interpretations of the mid-Cretaceous carbonate succession of Israel at that time. More recent paleogeographic and paleoecologic studies of the mid-Cretaceous carbonate system of Israel (e.g. Bein 1974, 1976; Sass & Bein 1982; Braun & Hirsch 1994) continued to apply these concepts.
5
Chapter 1: Introduction
However, the extreme variability of carbonate systems in time requires caution in implementation of the principle of uniformitarianism, as argued by Insalaco et al. (2000). Modern carbonate systems are not at all like their Cretaceous counterparts. They never were analogous to the extremely- thick and widespread accumulations of greenhouse foramol-type rudist-dominated carbonate platforms of the mid-Cretaceous passive-margins of the Levant basin margins. Moreover, modern approaches to carbonate systems over the last 40 years have incorporated concepts such as development of facies belts (Wilson 1975); profiles and ramp configurations (Read 1985; Burchette & Wright 1992; Pomar 2001); the role of sea level (e.g. Vail et al. 1977a; Haq et al. 1987); sequence stratigraphy of carbonate systems (Handford & Loucks 1993; Hunt & Tucker 1993); global sequence stratigraphic correlation across the Arabian Platform (Sharland et al. 2001); paleoenvironmental controls on carbonate systems and accommodation space (e.g. Schlager 1991, 1993, 1999) and the role of oceanic anoxic events (Schlanger & Jenkyns 1976). Most significantly, the ecological role of rudist bivalves, a major biotic element controlling geometry and facies in Cretaceous carbonate systems, was re-evaluated (e.g. Ross & Skelton 1993; Gili et al. 1995; Steuber 2000; Steuber & Löser 2000). In this light, the time is ripe for raising new questions regarding the CenomanianTuronian (Ce-Tu) carbonate system in northern Israel which are the focus of this research: 1. What are the basic and large-scale genetic stratigraphic units constructing the Cenomanian-Turonian sedimentary system in northern Israel? Does the local sequence framework correspond to that of the Arabian plate and beyond? 2. What is the tectono-sedimentary configuration of the mid-Cretaceous carbonate system in northern Israel and broadly in the northern Levant? How does it corresponds to previous concepts of the mid-Cretaceous carbonate system of Israel (presented above) such as the 'hinge-belt' configuration of the Levant margin, the widespread 'shelf-lagoon'-type carbonate platform, and presence of intrashelf basins. 3. In what way was the carbonate system of northern Israel and the northern Levant affected by the global paleoceanographic events of the mid-Cretaceous? What was the effect of the Oceanic Anoxic Event (OAE) - 2 and eustatic changes on facies type and distribution, geometry of large-scale genetic units and platform configuration? Can local tectonic, paleoceanographic or sedimentary events be isolated?
6
Chapter 1: Introduction
1.4. Methods and procedures 1.4.1. Field and Laboratory procedures •
This study is based on high-resolution sequence stratigraphy at the level of detailed bed to bed sampling of outcrops, including macroscopic observation of sedimentary textures, structures, and paleontological content, at outcrop level.
•
Mesoscopic examination of textures, sedimentary structures and diagenetic features at outcrop and hand sample scales.
•
Microscopic examination of thin-sections using polarized light microscopy and occasionally, cathodoluminescence. Microfacies analysis (following Dunham 1962; Embry & Klovan 1971 and Flügel 2004) included the recognition of non-skeletal and skeletal grains, microfacies definition, and special emphasis on diagenetic stratigraphy, especially of bounding surfaces.
•
Micro- and macro-fossils were submitted for expert determination.
1.4.2. Stratigraphic procedures •
Lithofacies units based on grain composition, carbonate facies classification, and fossil content and condition, were divided into interpretative facies-associations reflecting static or dynamic paleoenvironmental suites.
•
Vertical logs were prepared highlighting vertical facies changes, cyclic trends at various hierarchies, and relevance to defined and mapped lithostratigraphic units.
•
Large-scale facies units and facies transitions were traced in outcrops, panoramic mosaics and aerial photos.
•
Criteria for definition of bounding-surfaces were established in outcrops and by petrographic procedures.
1.4.3. Methods and concepts The approach implemented in this study is based on elaboration of the early, three system-tracts sequence-stratigraphic model, proposed by Van Wagoner et al. (1988) and Posamentier et al. (1988), into a four system-tract model by Hunt & Tucker (1992) and Helland-Hansen & Gjelberg (1994), and on modifications to carbonate systems by Hunt & Tucker (1993). Typically, a sedimentary sequence is composed of four system tracts: lowstand (LST), transgressive (TST), highstand (HST) and occasionally forced regressive system tracts (FRST). Stratigraphic bounding surfaces used here include: (a) the sequenceboundary unconformity at the top of the FRST and its basinward correlative conformity; (b)
7
Chapter 1: Introduction
the basal surface of forced-regression at the base of the FRST; (c) the regressive surface of marine erosion (RSME; Plint 1988) at the base of forced-regressive shoreface deposits; (d) the maximum-flooding surface at the top of the TST; (e) a transgressive surface and shoreface ravinement surface at the start of the TST. The concept of type-3 sequence boundary and drowning unconformity introduced by Schlager (1999) was used to explain short-lived subaerial exposures on HSTs, prior to the development of drowning successions. For the purpose of regional correlation, the Transgressive-Regressive sequence stratigraphy of Embry (2002) was implemented, using the sequence boundaries and maximum-flooding surfaces for broadly defining transgressive and regressive system-tracts. Using this approach, the resolution of the sequence stratigraphic correlation is lowered, but the reliability of the model is increased. The four system-tract method of Hunt & Tucker (1992) could not always be implemented successfully on sections described in the literature in Israel, Lebanon or Syria. That limitation may result from insufficiently detailed or focused field descriptions, lower sampling resolution, consequent lack of biostratigraphic data, and implementation of a lithostratigraphic rather than allostratigraphic (sequence stratigraphic) approach for the description and analyses of sections. As a result, considerable literature data (e.g. Kafri 1972, 1991) could not be integrated into this thesis. Descriptions and analysis of synsedimentary gravity-collapse features related to mechanical failure of carbonate slopes is a major theme presented in chapters 4 and 5. This subject can be addressed quantitatively by means of rock mechanics analysis, but the approach taken in this thesis is purely qualitative aimed at reconstructing the mid-Cretaceous shelf-edge in northern Israel, where slope-failure processes are at times a defining feature. The mechanics of failure of carbonate ramps may at times be treated by means of rock mechanics analysis, but has rarely been done quantitatively in subsequently lithified material from the geological record. Reconstructing the original state of cohesion of the sediment can be highly subjective, in the absence of key parameters, such as opaline silica and organic material subsequently removed, pre-diagenesis grain parameters, and even the angle of repose on the paleoslope, whose inclination was determined here by paleobathymetry of benthic organisms.
8
Chapter 1: Introduction
1.5. Structure of the thesis 1. Chapter 1, introduction (this chapter):
To present the aims of this research, the
concepts, methods and procedures applied, and the scientific background relevant to all the following chapters. 2. Chapter 2, database: Presentation of database, and interpretation of depositional environments, facies units and cyclic patterns of the Cenomanian-Turonian succession of northern Israel, and introducing the paleoecological influence on highand low-order cycles. 3. Chapter 3, sequence stratigraphy and correlation: (a) A sequence stratigraphic framework for northern Israel, based on the recognition of cyclic stacking patterns, identification of bounding surfaces, and chrono- and biostratigraphic constraints; (b Cenomanian-Turonian system tract evolution in northern Israel, and the emergent sedimentary-structural configuration and; (c) correlation of bounding surfaces to the global Tethyan framework. 4. Chapter 4, Mid-Cenomanian ramp termination: The way in which the MidCenomanian distally steepened ramp was terminated by mechanical ramp collapse, sub-aerial exposure and sub-marine omission. 5. Chapter 5, system tracts around the Cenomanian/Turonian boundary: The way that local tectonics, synsedimentary block rotations and eustatic changes interacted to effect a dramatic transformation of a homoclinal ramp into a steep slope 'shelf-margin apron' system. In this study, the Cenomanian “rudist reef” concept, deeply rooted in earlier models (Freund 1965; Bein 1974) is objectively re-examined in northern Israel 6. Chapter 6, regional and global implications: Extension of the local structuralsedimentary framework of northern Israel (Chapter 3) to the northern Arabian platform in Lebanon and Syria, and the regional expression of the global oceanic anoxic event (OAE-2) in the northern Levant.
9
10
11
Chapter 2: Facies and cyclic patterns
Chapter 2 Carbonate facies, depositional environments, and cyclic patterns in the Cenomanian-Turonian succession in northern Israel 2.1. Introduction This
chapter
is
mostly
descriptive,
presenting
the
database
and
its
paleoenvironmental interpretation that form the basis of the sequence stratigraphic framework presented in Chapter 3 and of the various syntheses presented in the remaining chapters. 22 colunmar sections were measured, sampled and analysed. In this chapter the database of six representative columnar sections is presented and then synthesized into facies types presented in Table 2.1. Facies types, cyclic patterns, cyclic hierarchies, and finally, the paleoecological significances of the low-order cycles are then discussed. The complete database of 22 columnar sections, their facies description, division into subunits, and the detailed descriptions of thin sections, are all given in Appendix 2.
2.2. Representative columnar sections, their facies types and construction of subunits The six sections described here are the Deir El-Assad, Yanuch and Beit-Ha’Emek of the western Galilee, Betzet of the NW Galilee, Dishon of the eastern Galilee, and Manara of the NE Galilee (see Fig. 1.1 for locations). Most of the 20 facies types presented in Table 2.1 were synthesized from these sections. 2.2.1. The Deir El-Assad section, western Galilee (Fig. 2.1) 2.2.1.1. Subunit DS1 (Sakhnin Fm) The lower part of DS1 is massively bedded, completely dolomitized, with some unbroken radiolitid rudists observed. The upper part is well-stratified, with rare fragments of mollusks or shallow-water benthic foraminifera (mostly nezazzatids) in thin-section. Bedding is in alternations of crinkled micrite laminites with massive dolomitized mudstones. Lagconglomerates, flat-pebble conglomerates, ferruginous cements, peloidal fabrics and cement-filled fenestrae and vadose silts occur mostly in the stratified upper part. DS1 forms a single shallowing-upward cycle. The bioturbated lower part was deposited on the mid-ramp below fairweather wave-base. Homogenization by bioturbation and occurrences of unbroken rudists suggest that wave or current energy has a minor affect on sediment dispersal. The upper well-stratified part of DS1 is of meter-thick peritidal cycles (e.g. Pls. 1a, 2e, 5b,c) starting with lag-conglomerates reflecting omission and erosion prior to ensuing sub-tidal deposition. Above, shallow-subtidal mudstones or wackestones were
13
Chapter 2: Facies and cyclic patterns
deposited under low-energy conditions. The cycles end by peritidal microbial laminites with vadose-type diagenetic features and flat-pebble conglomerates, indicating short-term subaerial exposures and storm erosion on tidal flats. This stratified peritidal facies of Deir ElAssad region passes into meter-sized breccias toward the west. 2.2.1.2. Subunit DS2 (Bina Fm) This is a highly bioturbated, fossil-rich succession (Pl. 1b) with large chondrodontid bivalves, Acteonella sp. (Opisthobranch gastropods), small gastropods, rare radiolitid rudists, fragmented echinoderms, and shallow-water benthic foraminifera such as miliolids and nezazzatids. Intense bioturbation and abundant macro- and micro-fauna point toward deposition below fairweather wave-base. Compared to the underlying mud-dominated peritidal facies (DS1), subunit DS2 reflects significant increase in skeletal production during relative sealevel rise. Cycle DS2 also represents the termination of the peritidal inner-ramp of subunit DS1 by deepening, and the termination of cyclic sedimentation. 2.2.2. Yanuch section, western Galilee (Fig. 2.2) 2.2.2.1. Subunit YN1 (lower Yanuch Fm) YN1 makes up the lowermost part of the Yanuch section. This succession is characterized by monotonous, thin-bedded, well-laminated mudstones or calcisiltites (Pl. 1c) with few planktonic foraminifera, pithonellid calcispheres and some echinoderm fragments. The transition to YN2 above is sharp and erosive. This unit was deposited on a stagnant hypoxic bottom of open-marine basinal slope. Most of the sediment is a bioclastic-calcisiltitic suspended load. 2.2.2.2. Subunit YN2 (Yanuch Fm) YN2 begins directly above the erosional discontinuity surface of YN1 with 8 meters of massive packstones and floatstones (Pl. 1c). The sediments are completely bioturbated containing coarse-grained fragments of oysters, some rudists and echinoderms, as well as pithonellid calcispheres and planktonic foraminifera. YN2 ends with a distinctive interval composed of well-laminated fine-grained ferruginous pelagic packstones with abundant pithonellid calcispheres, planktonic foraminifera, pelagic echinoderms, sponge spicules and phosphatic fragments. The bioturbated, calcisphere-rich skeletal deposits with chert and phosphatic grains indicate nutrient-enriched surface waters, increased primary productivity, and consequently, increased biogenic activity above an oxic sea-floor (c.f. Banner 1972; Dias-Brito 2000; Föllmi
14
Chapter 2: Facies and cyclic patterns
et al. 1994). The well-laminated pelagic termination of YN2 suggests a return to a stagnant sea floor, still under nutrient-rich surface waters. 2.2.2.3. Subunit YN3 (Yanuch Fm) YN3 is composed of three similar genetic cycles; only the first (YN3.1) is here described. The base of YN3.1 is a bioturbated interval composed of packstones or floatstones with a few phosphatic grains. There is some macro-benthic fauna of unbroken, solitary radiolitid rudists, crinoid and echinid plates, gastropods (including Nerinea sp.), as well as pithonellid calcispheres, planktonic foraminifera and few benthic foraminifera (e.g. miliolids, Pseudorhapydionina sp. and rotaliids). Early calcite cements occur towards the top of the cycle. This bioturbated interval is overlain by a grainstone bed (Pl. 1d) that is occasionally graded, with fragments of rudists and red- and green calcareous algae. The grainstones are overlain by a shell-rich bed marking the beginning of the following YN3.2, consisting mainly of oysters with fragments of echinoderms and phosphate grains. This shell bed is replaced upwards by fine-grained packstones with deep-water elements, mainly pithonellid calcispheres, planktonic foraminifera, sponge spicules, fragments of pelagic echinoderms and small thin-shelled bivalves. The bioturbated lower part of the YN3.1 represents resumption of minor skeletal production after the maximum-flooding stage that ended YN2. The overlying grainstones reflect higher hydraulic energy and are considered as wave/current-controlled shoreface facies. They represent a peak in an upward-shallowing trend. The following shell-bed is an event-concentration (sensu Kidwell 1986) formed by a reduction in background carbonate supply. This event is associated with termination of progradation and drowning of the underlying bed due to weakening of wave- and storm-induced winnowing that promoted rudist and oyster occupation of available hard substrate. 2.2.2.4. Subunit YN4 (top Yanuch, lower Yirka Fm) Subunit YN4 is composed of three similar small-scale cycles with common depositional features; only the first (YN4.1) is here described. Cycle YN4.1 overlies YN3 with a sharp erosive contact. Above this surface, lie cross-bedded fining-upwards peloidal grainstones of molluscan debris and calcareous green algae (Pl. 1e). Reworked grains, mainly mud- and mold-peloids (sensu Flügel 2004) with some shallow-water lagoonal benthic foraminifera such as miliolids, Rhipidionina sp. and Cuneolina sp. are common in these grainstones (Pl. 1f). The grainstones pass upwards into fine-grained pelagic packstones and marls interbedded with well-sorted, graded or massive peloidal-bioclastic grainstones (Pl. 1, g,h). YN4.1 terminates with a massive bed of coarse bioclastic packstone of intensely bioeroded bioclasts and small phosphatic concretions, fish debris, and oyster debris. YN4.2
15
Chapter 2: Facies and cyclic patterns
and YN4.3 also commence with autochthonous deep-water marls and chalks and terminate with similar bioclastic concentrations. The sharp-based basal surface of YN.4.1 and the presence of shallow-water mudpeloids in the cross-bedded grainstones above it indicate reworking of earlier lagoonal deposits by shoreface waves. The stratigraphic position of the shoreface grainstones is therefore above an erosional discontinuity. The shoreface facies deepens upward into the Early Turonian Yirka Fm consisting of thin-bedded pelagic chalks-marls intercalated with graded grainstones representing turbidites. The massive limestone bed at the end of YN4.1 represents reduced basinal sedimentation, as indicated by the authigenic mineralization and the intense biodegradation of bioclasts. This condensed bed terminates this deepeningupward cycle. Above YN4.1 two additional genetic cycles YN4.2 and YN4.3 were recognized. 2.2.3. Bet-Ha’Emek section, western Galilee (Fig. 2.3) 2.2.3.1. Subunits BK1 and BK2 Subunits BK1 and BK2, mapped as Yanuch and Yirka Fms, are similar to previouslydescribed subunits YN3 and YN4. 2.2.3.2. Subunit BK3 (Kishk Fm) BK3 begins with thin-bedded poorly-washed peloidal grainstones alternating with hummocky cross-stratified peloidal grainstones (Pl. 2a; Buchbinder et al. 2000). Overlying this part is a thick interval of stratified peloidal grainstone forming clinoforms dipping 10°-15° to the SE. Some clino-beds are colonized by decimeter-scale rudist concentrations, embedded in a well-cemented peloidal matrix (Pl. 2b). They are loosely-packed and disarticulated, unaffected by post-mortem biogenic activity (e.g. borers, encrusters) or secondary mineralization. The terminal part of BK3 is marked by well-stratified alternating marl and dolomite with crinckly laminites and mud-cracks. BK3 consists of six meter-scale small-scale cycles (BK3.1-3.6) dominated by mudand mold-peloids and some quartz grains. Small-scale cycles BK3.1 and BK3.2 are meterscale shallowing-upward cycles, each commencing near storm-wave-base, with thin-bedded peloidal grainstones, and terminating in hummocky cross-stratification in the offshore transition zone. Under fairweather conditions, relatively weak wave-controlled motion of the bottom-water caused gentle reworking of peloids.
Under storm conditions, interaction
between the grainy bottom and oscillatory waves led to hummocky cross-stratification. All these peloidal grainstones were subsequently cemented by characteristic granular or blocky intragranular cements. BK3.3 to 3.6 are meter-scale shallowing-upward cycles intercalated within the clinobeds of BK3, indicating migration of offshore-bars within the fairweather wave zone. Each
16
Chapter 2: Facies and cyclic patterns
cycle begins with a simple rudist-bearing event-bed (sensu Kidwell 1986) of homogeneous internal composition. These event beds would have formed under decreased storm activity. Physical abrasion of shells and the highly-cemented mud-free grainy matrix indicate high hydraulic energy and active inter-granular water flushing characterizing events of rudist colonization. The limestone-marl upper part of BK3 is composed of numerous, meter-scale peritidal cycles. Massive mudstones and wackestones with sporadic shallow-water benthic foraminifera are shallow sub-tidal indicators. They are capped by supratidal marls or by microbial mats with fenestrae and desiccation cracks. The entire BK3 can be viewed as a shoaling-upward cycle. BK3.1 to BK3.3 represent the offshore-transition zone terminating above storm wave-base. BK3.4 to BK3.6 were deposited primarily in the upper shoreface zone. The upper part represents shoaling to the peritidal zone. 2.2.4. Betzet section, northwest Galilee (Fig. 2.4) The Betzet section demonstrates facies evolution of the interval from the lower Deir Hanna Fm to the top of the Bina Fm. 2.2.4.1. Subunit BZ1 (lower to mid- Deir Hanna Fm) BZ1 begins with monotonous, thin-bedded or laminated, partly silicified white chalk with fine-grained bioclasts, sponge remains, echinoderm debris and micro-pelagic fossils such as planktonic foraminifera and pithonellid calcispheres. Upwards, the fabric changes to laminar wackestone or packstone enriched in pithonellid calcispheres (e.g., Pithonella sphaerica). The cycle terminates with a hard limestone with coarse-grained fragments of echinoderms, bioeroded mollusk debris, rare planktonic foraminifera, some quartz grains and siliceous cementation of borings (Pl. 2c); there is no evidence for current action or grain mass-movement. BZ1 is a deepening-upwards interval. The laminated mudstones at the base reflect sedimentation by suspension into hypoxic bottom waters. Occasional mass occurrence of calcispheres points toward nutrient rich surface-waters (Banner 1972; Jarvis et al. 1988; Wendler et al. 2002). The massive bioclastic limestone at the top of BZ1 corresponds to a hiatal concentration sensu Kidwell (1993) forming under a low rate of accumulation. 2.2.4.2. Subunit BZ2 (upper Deir-Hanna and Sakhnin Fms) BZ2 commences with laminated, chert-bearing basinal mudstones of the upper DeirHanna Fm above the hiatal limestone of top BZ1. The laminites pass sharply upwards into bioturbated dolomites of the Sakhnin Fm. Pithonellid calcispheres and sponge spicules in the
17
Chapter 2: Facies and cyclic patterns
lower part are replaced upwards by dolomitized macrobenthos; shallow-water foraminifera, especially miliolids, are present. The upper part of BZ2 is well-stratified, with microbial laminites, fenestrae, flat-pebble conglomerates, and vadose silts. This interval is locally associated with shear-zones and breccias. BZ2 is a shallowing-upwards interval, with hypoxic laminites at the base, passing upwards into progradational bioturbated mollusk-rich mid-ramp facies. Further shallowing to peritidal depths is reflected by the upper stratified part. 2.2.4.3. Subunit BZ3 (lower Bina Fm) BZ3 is composed of two parts. The lower part is composed of alternating laminated sterile mudstone and massive bioturbated wackestones (Pl. 2d), with bioturbated beds bearing miliolids and other shallow-water benthic foraminifera. The upper part is composed of well-stratified dolomites with microbial laminites and ferruginous crusts (Pl. 2e). The lower part of BZ3 reflects fluctuations of oxygen or nutrient input on the innerramp below fairweather wave-base. The upper part, with supratidal features, reflects the peritidal environment. 2.2.4.4. Subunit BZ4 (Bina Fm) BZ4 is composed of three beds: (a) Massive poorly fossiliferous mudstones with foraminiferal wackestone and few rudists, possibly hippuritids (Pl. 3a). This bed passes upwards to the following unit by a sharp erosive surface (Pl. 3b); (b) Nodular mudstones with scattered dolomite rhombs and few phosphatic grains (Pl. 3a); (c) The topmost bed, only 10cm in thickness, is well-laminated mudstone with sub-horizontal ferruginous films and bioclasts showing preferred orientation (Pl. 3c). The massive basal bed of subunit BZ4 was deposited in a low-energy sub-tidal midramp setting. The subsequent nodular limestones presumably formed at a slower deposition rate in a deeper setting after a period of non-deposition and transgressive erosion. The topmost bed reflects bottom stagnation and hypoxia, corrosion of bedding planes and sediment starvation during flooding. 2.2.4.5. Subunit BZ5 (Bina Fm) The lower part of BZ5 commences with pelletal-intraclastic miliolid packstone, recrystallized and neomorphic wackestones or sterile mudstones, topped by crinkly microbial laminites. Toward the upper part, beds become thicker (3-5m) and bioturbated with large gastropods, ostreids (chondrodont bivalves), abundant sponge spicules, echinoderm debris and unbroken rudists. These upper beds are also topped by crinkly microbial laminites.
18
Chapter 2: Facies and cyclic patterns
Subunit BZ5 is composed of peritidal shallowing-upward cycles. The abundant rudists, gastropods and oysters of the upper part may reflect more open circulation on the inner ramp and increased skeletal production of the subtidal units. 2.2.5. Dishon section, northeast Galilee (Fig. 2.5) 2.2.5.1. Subunit DO1 (Sakhnin Fm) DO1 at Dishon consists of well-stratified peritidal cycles, similar to those described from Deir El-Assad and Betzet. 2.2.5.2. Subunit DO2 (Bina Fm) DO2 commences with a moderately- to poorly-bedded succession dominated by fecal pellets, peloids, aggregate-grains, microbial lumps, heavily degraded micritized bioclasts and cortoids and some shallow-water benthic foraminifera (Pl. 3d), with rare open-marine components. It passes upwards into coarsening-upward grainstones composed mainly of mud-peloids (lithoclasts, rip-ups), rounded aggregate-grains and benthic foraminifera. Two generations of calcite cementation were precipitated within the grainstones: early, very thin isopachous fringes, and later blocky or poikilotopic crystals. The grainstones pass upward into thin-bedded pelagic limestones with planktonic foraminifera, pithonellid calcispheres, echinoderms (including pelagic forms) and phosphatic grains (Pl. 3e). DO2 represents deepening upward conditions, commencing with relatively stressed semi-restricted lagoonal facies. This facies passes upwards through an erosional surface into high-energy shoreface grainstones composed of grains derived from the underlying lagoonal facies. DO2 ends in pelagic outer ramp facies reflecting open-marine nutrient-rich water column. 2.2.5.3. Subunit DO3 (Bina Fm) DO3 is composed of two small-scale cycles DO3.1 and DO3.2, of the same type as those composing YN3 in the Yanuch Fm. 2.2.5.4. Subunit DO4 (Bina Fm) DO4 is a 50m-thick succession composed of well-bedded and well-laminated mudstones and wackestones (Pl. 3f), divisible by discontinuities into four small-scale cyclic subunits (DO4.1-DO4.4). DO4.1 begins with a bioclastic shell-rich bed with solitary corals, echinoderms, coarse fragments of ostreids, gastropods and few benthic foraminifera, and some unbroken radiolitid rudists. Shells and bioclasts are severely bioeroded, borings are filled by silica cements, and the grains are coated by ferruginous oxides. This bed is overlain by
well-stratified,
well-laminated
mudstones
19
with
sparse
gavelinellids,
planktonic
Chapter 2: Facies and cyclic patterns
foraminifera, thin-shelled ostracodes, rare fragments of echinoderms (including pelagic forms) and fecal pellets (Pl. 3g-i). Ooid-pelletoid grainstones overlie these laminites, with first-generation isopachous marine pore-filling cements, followed by granular mosaics. DO4.2 is similarly composed of laminites topped by ooid-pelletoid grainstones, but DO4.3 is topped by a massive bed of pelletal-lithoclastic packstones with Cuneolina sp., Dicyclina sp. (foraminifera), Thaumatoporella sp. (algae) and fragments of echinoderms. It is overlain by DO4.4 beginning with a shell-rich bed bearing bioeroded and fragmented radiolitid rudists, ostreids and echinoderms, a few benthic foraminifera, and cavities filled by silica cements or phosphatic concentrations (apatite). This shell bed passes upwards to subtidal laminites. The energetic conditions terminating DO3 were rapidly replaced by condensation and colonization by open marine macrofauna. This system was not sustained, and shells were slowly bioeroded, partly silicified and ultimately buried by distal ramp laminites of DO4.1. This well-laminated facies bears low density gavelinellid-dominated foraminiferal faunas, interpreted as indicating somewhat hypoxic bottom waters that shallow upwards to the oolitepellet facies of the shoreface zone. These shallowing-upwards conditions are repeated in DO4.2 and DO4.3. The shell bed at the base of DO4.4, however, represents drowning of the top of DO4.3, indicated by development of condensation features. The well laminated gavelinellid mudstones of this cycle represent return to oxygen-depleted open marine conditions and do not shallow upwards. DO4 reflects gradual deepening of the DO3 facies to below storm wave base, indicating significant flooding. 2.2.5.5. Subunit DO5 (upper Bina Fm) An oyster-rich shell bed overlies cycle DO4. The stressed hypoxic conditions represented by the mostly well laminated mudstones of cycle DO4 were replaced by oxic conditions. 2.2.6. Manara section, northeast Galilee (Fig. 2.6) 2.2.6.1. Subunit MN1 (lower Deir Hanna Fm) MN1 is well-laminated pelagic mudstones 15m thick, overlying coarse bioclastic limestones (floatstones) with oysters and bivalves (Ceratostreon flabellata and Liopistha pervinquierei) and the rudist Eoradiolites lyratus (Ident. Z. Lewy 2006 pers. Comm.) (Pl. 4a). Fine-grained bioclastic debris forms the major constituents of MN1. Microfaunal elements includes small and poorly preserved heterohelicids and pithonellid calcispheres, fragmented echinoderms (including pelagic forms), fragments of bivalves (some highly bored), rare nezazzatids and some small biserial benthic foraminifera. The upper MN1 is a massive bioclastic packstone, 80cm in thickness, with abundant flat orbitolinids including O. sefini (Ident. M. Simmons 2007, pers. comm.) (Pl. 4b). Also present are biserial benthic
20
Chapter 2: Facies and cyclic patterns
foraminifera (probably Praechrysalidina cretacea), rare Pseudolituonella sp., rotaliids (including gavelinellids), trochospiral planktonic foraminifera with incipient keels, a few phosphatic grains and bone fragments of fish, serpulids and gastropods. The matrix is of fine grained bioclastic debris with small spar-filled pithonellids and calcitized sponge spicules. The lower laminated mudstones with heterohelicids and small pithonellids suggest hypoxic conditions in the water column. The orbitolinid bed at the top was deposited in an outer-ramp setting under higher oxygen levels. O. sefini, indicating the lowermost Cenomanian, places the cycle at just above the Albian-Cenomanian boundary. The AlbianCenomanian transition in (the neighboring) Lebanon was reported at this level within a similar succession. 2.2.6.2. Subunit MN2 (Deir Hanna – Sakhnin Fms) MN2 is composed of well laminated, fine grained pelagic packstones with abundant echinoids at the base, passing upwards into laminated or thin-bedded, often graded calcarenites and calcisiltites (Pl. 4c). At the top, 20m of massive dolomites occur, with unbroken rudists and bivalves. A well-cemented, silicified lithoclastic hardground is developed on the surface, with diagenetic features such as vadose pisolites, floatstones and empty dissolution vugs. MN2 is a shallowing-upward cycle commencing with basinal laminites, passing upwards to calcarenitic turbidites, shallowing further upwards into a rudist-bivalve mid-ramp facies, and ends by subaerial exposure. 2.2.6.3. Subunit MN3 (Yanuch Fm) Interval MN3 is constituted of well bedded, laminated chert-bearing calcisiltite mudstones with fine bioclastic debris and planktonic foraminifera, pithonellid calcispheres, abundant echinoderms, and the ammonite Calycoceras sp. (Id. Z. Lewy 2006 pers. comm.). This interval was deposited in open marine environment on distal hypoxic slope. The ammonite Calycoceras sp. denotes late Mid-Cenomanian to Late Cenomanian age for this interval.
2.3. Synthesis of sedimentary facies types and their environments of deposition Table 2.1 summarizes the carbonate facies types occurring in the columnar sections described above, and in the detailed sections given in Appendix 2. In all, twenty Facies Types (FT) were synthesized, given as FT-1 – FT-20 (column 1). Column 2 in Table 2.1 gives the details of sedimentology, bedding and faunal components. These sedimentary facies reflect a variety of depositional environments ranging from autochthonous basinal
21
Chapter 2: Facies and cyclic patterns
deposits/surfaces to supratidal and subaerial exposure facies, as well as environments characterized by mass-transport. The paleoenvironments indicated by these features are shown in column 3, and schematically in Figure 2.7. They correspond to the carbonate ramp subdivision into basin, outer-ramp, mid-ramp and inner ramp of Burchette & Wright (1992). Shell beds and fragmental concentrations were analyzed according to the criteria established by (Kidwell 1986, 1993), and subaerially-exposed environments were analyzed following Dunham (1969), Longman (1980) and Flügel (2004). Column 4 lists the lithostratigraphic units in which these facies types are present, while their position within the sequence stratigraphic scheme is given in column 5. Note that although dolomite is a rather common lithology, only FT 19 has a facies consistent with formation of syngenetic or penecontemporaneous dolomite. For this reason, dolomites present in other facies are considered to be the result of later dolomitization processes originating in hydrological continuity with suitable brines. This assumption is borne out by field and petrographic evidence, where original carbonate components are usually readily identifiable, and not associable with hypersaline paleoenvironments or others in which early crystallization of dolomite takes place in the sedimentary environment.
22
Chapter 2: Facies and cyclic patterns
Table 2.1: Facies types of the Cenomanian-Turonian succession of northern Israel, with environmental interpretation, lithostratigraphy and sequence-stratigraphic position.
23
Chapter 2: Facies and cyclic patterns
24
Chapter 2: Facies and cyclic patterns
2.4. General outlines of cyclicity in northern Israel Sedimentary cycles in northern Israel were classified into two orders: High-order cycles that cannot be further divided into smaller cycles, termed Undividable Cycles (abbreviated as UC); and low-order cycles that are composites of high-order cycles, termed Composite Cycles (abbreviated as CC). These high- and low-order cycles are objectively described, free of the time-related constraints originally defined by Vail et al. (1977b). High-order cycles are groupings of facies types FT1 to FT20 (Table 2.1, Fig. 2.7). They are arbitrarily numbered UC1-12. They are graphically presented and explained in Figures 2.8 and 2.9. UC’s are mostly decimeters to a few meters in thickness but may in some cases be more than ten meters thick. UC-1 to UC-6, are shallowing-upward cycles (Fig. 2.8) and UC-7 to 12 are deepening-upward cycles (Fig. 2.9). Some of these high-order cycles (UC-1 to UC-6, and UC-7, 11 and 12) are combined into low-order Composite Cycles (CC), but deepening-upward cycles UC-8 to UC-10 (Fig. 7) are independent. The low-order composite cycles are graphically simplified in Figure 2.10. Their thicknesses range between few meters to a few tens of meters.
25
Chapter 2: Facies and cyclic patterns
2.5. Discussion 2.5.1. Low-order cycles and their paleoecological significances Stacking of high-order cycles (UC’s) into different types of low-order Composite Cycles (CC’s) are as follows: 2.5.1.1. Type-1 Low-order Composite Cycle (CC1) Type-1 low-order Composite Cycle (Figs. 2.10, 3.1) begins with basinal laminites, passes upwards into massive fossiliferous mid-ramp deposits, and is topped by stacked UC1 peritidal cycles. This is a shallowing-upward, progradational cycle. The productive mid-ramp facies of this cycle resulted in filling of accommodation space to near sea level. Thus, the vertical facies succession of CC1 reflects a progradation pattern by which carbonate was produced efficiently enough in order to fill the entire accommodation space beginning in the deep basin to sea-level (Fig. 2.11a). A schematic ramp model for CC1 (Fig. 2.10a) highlights details of the paleoecological control on this type of progradation. The middle part of CC1 represents a bioturbated carbonate factory providing micrite flux both into the slope/basin and to the peritidal environment. A stagnant well-laminated basin facies occurs beyond the slope. The sharp boundary separating the bioturbated from the laminated basin facies is attributed to oxygen constraints, modeled as a transition from a mixed layer to an oxygen minimum zone (OMZ). Progradation of the carbonate factory was inhibited by a raised OMZ. Surface phytoplankton blooms are evident from the lower basinal part of CC1 (at Betzet) where intervals rich in pithonellid calcispheres occur. Cyclic appearances of layered chert nodules, most probably of biogenic opaline origin, similarly indicate periodic blooms of siliceous phytoplankton. Progradation of the carbonate factory was therefore cyclically interrupted by the fluctuating OMZ. Fluctuation of oxygenation in the water column resulted in an ecological gradient from the carbonate factory to the basin. Rapid fall-off of carbonate production where the OMZ impinged on the slope caused distally-steepened ramp morphology to develop (see model in Fig. 2.10a). 2.5.1.2. Type-2 Low-order Composite Cycle (CC2) Type-2 Low-order Composite Cycle (Figs. 2.10b, 3.1) includes stacked UC-3 highorder cycles, each with a basal hiatal shell concentration, passing upwards into bioturbated or laminated pelagic basinal or outer-ramp facies and topped by cross bedded or bioturbated shoreface grainstones. The pelagic mid-part tends to predominate in CC2. It is rich in calcispheres and planktonic foraminifera, phosphate grains, silica as grain replacement and pore cements. Fragments of rudists occasionally appear among bioeroded (bored) molluscan skeletal debris.
26
Chapter 2: Facies and cyclic patterns
Rudist bivalves, which usually form the progradational systems of the Cenomanian (e.g. Steuber & Lösser 2000) are rare in CC2. The presence of phosphatic grains and the activity of bioeroders in these cycles suggest that the rate of accumulation was relatively slow. Accommodation space was filled only to fair-weather wave base indicating that carbonate production rate could not keep-up with relative sea-level rise as in UC-1 cycles. The grainstones forming the UC-3 tops reflect therefore inefficient progradational episodes in which shoreface deposits were formed at fairweather wave-base, above which they were dispersed. This type of progradation is reflected in Figure 2.11b. The failure of CC2 to fill accommodation-space as a result of slow carbonate production rate needs explanation, as growth potential of healthy platforms is usually capable of exceeding long term subsidence or sea-level change (Schlager 1981; Schlager & Philip 1990; Bosscher & Schlager 1992). Some of the features characterizing CC2 point towards a stressed nutrient-enriched environment. Banner (1972), Jarvis et al. (1988) and Wendler et al. (2002) connected midCretaceous pithonellid-rich deposits to nutrient-rich conditions in surface waters; Hallock (1988) found a positive connection between bioerosional phenomena and eutrophic conditions; Föllmi et al. (1994) considered phosphogenesis as a link in a chain of feedback mechanisms in association with nutrient mobilization and increased primary productivity. The failure of the CC2 cycle to completely fill accommodation space is therefore attributed to relatively eutrophic conditions that prevented development of a prograding carbonate factory by inhibiting rudists and other skeletal macrobenthos. The bioturbated intervals at the base of CC2 indicate incipient development of a benthic community on the ramp but later in the cycle they were replaced by laminated or highly-pelagic sediments reflecting rise of OMZ before a productive carbonate system could be established. 2.5.1.3. Type-3 Low-order Composite Cycle (CC3) Type-3 Low-order Composite Cycle (Figs. 2.10c, 3.1) begins with a stack of lower shoreface UC-4 high-order cycles, passing upwards into UC-5 upper shoreface cycles, and ending in peri-tidal UC-1 cycles. This cycle is mostly composed of peloidal grains and an autochthonous carbonate-factory such as in CC1 is absent. In-situ skeletal growth is limited to episodic events of rudist shell beds found at the bases of UC-5 cycles. Nevertheless, accommodation space was filled to near sea-level. Accommodation and progradation were maintained by supply of peloidal grains from the inner-ramp to the shoreface. 2.5.1.4. Type-4 Low-order Composite Cycle (CC4) Type-4 low-order composite cycle CC4 (Figs. 2.10d, 3.1) begins with stacked peritidal UC-1 cycles with rare macrobenthos and passes upwards into open-marine inner-ramp UC-2 cycles in which the sub-tidal mud substrate is colonized by rudists, large chondrodont
27
Chapter 2: Facies and cyclic patterns
bivalves, opisthobranch gastropods (Acteonella) and sponges. CC4 is a peritidal shallowingupward progradational cycle. 2.5.1.5. Type-5 Low-order Composite Cycle (CC5) Type-5 low-order composite cycle CC5 (Figs. 2.10e, 3.1) is constructed of stacked UC-6 cycles, each beginning with well-laminated mudstones and topped by massive shallow subtidal wackestones. The laminated mudstones of CC5 were deposited on a stagnant hypoxic sea-floor and the massive beds reflect oxic bioturbated mid-ramp facies. CC5 represents fluctuations in oxygen level on the mid-ramp and not necessarily expansion or contraction of accommodation space. 2.5.1.6. Type-6 Low-order Composite Cycle (CC6) Type-6 low-order composite cycle CC6 (Figs. 2.10f, 3.1) is composed of stacked UC7 cycles with platy or laminated marl or chalk at the base (turbidites, ammonites) and skeletal hiatal concentrations at the top. CC6 is a deepening-upwards basinal cycle. 2.5.1.7. Type-7 Low-order Composite Cycle (CC7) Type-7 low-order composite cycle CC7 (Figs. 2.10g, 3.1) is composed of stacked UC-11 and/or UC-12 deepening-upward cycles. CC7 reflects retrogradation of hypoxic laminites. 2.5.1.8. Type-8 Low-order Composite Cycle (CC8) Type-8 low-order composite cycle CC8 (Figs. 2.10h, 3.1) is composed of UC-8 cycles, each commencing with shallow-lagoonal facies, deepens upwards into shoreface grainstones and ends with pelagic outer ramp facies. CC8 is a low-order deepening-upward cycle. A common feature for some of the high-order (UCs) and low-order (CCs) cycles is deposition under stressed paleoecological conditions. Paleoecological stress is reflected in the low-order CC2 cycle (discussed above) but is especially expressed in deepening upward cycles.
CC6 and CC7 (Fig. 2.10f,g) are deepening-upward cycles composed of well
laminated basinal marls and chalks, with phosphatic grains, some pyrites, grain/cement silicification and abundant calcispheres and heterohelicid foraminifera. Coeval lower Turonian facies in the pelagic Daliyya marls of the Carmel region contains aside from heterohelicid foraminifera, dinoflagellates, high concentrations of total organic carbon (1.06%–2.02%), and pyrite, and were considered an ecologically-stressed hypoxic facies with high vegetative productivity (Honigstein et al. 1989). CC8 deepening-upward cycle (Fig. 2.10h) bear intervals enriched with phosphatic micro-grains and pithonellid calcispheres
28
Chapter 2: Facies and cyclic patterns
(dinofllagelate cysts). All these low-order cycles reflect nutrient enrichment, stagnation, and slow production rate of skeletal carbonate. Ecological stress is also indicated in few of the high-order cycles (UCs). In UC8 deepening-upward cycle (Fig. 2.9) thin-bedded or laminated mudstones bear microbial lumps, micritized grains, cortoids (Pl. 3d) and abundand pithonellis calcispheres, suggesting lengthy submarine exposure, slow deposition and rare injection of skeletal carbonate grains. In UC9 deepening-upward cycle (Fig. 2.9) sterile mudstone and laminated mudstone with some phosphatic grains are dominant, indicatin bottom stagnation of a non-productive facies. The UC10 deepening-upward cycle reflects a similar paleoecological trend (Fig. 2.9), bearing pithonellid calcispheres, phosphate, laminites, and bioeroded bioclasts.
2.6. Summary and conclusions 1. Carbonate facies types of the Cenomanian and Turonian rocks of the Galilee are diverse: •
Basinal and outer-ramp facies include oxic bioturbated packstones and wackestones, hypoxic pelagic laminites, and allochthonous deposits such as turbidites and conglomeratic and sandy carbonate debrites.
•
The mid-ramp encompasses high energy, wave-controlled environments of the shoreface, or low energy bioturbated settings, occasionally with rudists and gastropods.
•
The inner-ramp includes subtidal mudstones, and hypersaline intertidal to supratidal features including stressed inter- to supratidal microbialites.
•
Other environments include different types of starved sea floors, proximal and distal macrofaunal colonization surfaces, and semi-restricted lagoons.
2. Vertical evolution of the facies units shows that the basic building blocks of major parts of the Cenomanian and Turonian succession are individual high-order Undividable Cycles (UC) that often construct low-order Composite Cycles (CC), up to tens of meters in thickness. Twelve types of high-order cycles were documented, categorized into shallowing and deepening upward cycles. 3. Features of the high- and low-order cycles, especially of the deepening-upward cycles, are indicative of a paleoecological stress. These cycles reflect slow production rate of skeletal carbonate, high nutrient conditions, bottom stagnation and hypoxia.
29
30
31
32
33
34
35
36
37
38
39
40
Chapter 3: Sequence stratigraphy, N. Israel
Chapter 3 Sequence stratigraphy, facies evolution and tectono-sedimentary configuration of the Cenomanian-Turonian carbonate system of northern Israel 3.1. Introduction The Galilee and Carmel of northern Israel are located in the central part of the Levantine passive margin of the Arabian Plate. This chapter deals with mid-Cretaceous carbonates deposited during a major formative episode in the tectono-sedimentary development of this region. Previous studies of the Cenomanian-Turonian of northern Israel focused on localized sedimentological, palaeogeographical or palaeoecological issues. Broadly, Freund (1965) and Kafri (1972, 1991) considered the Galilee a shallow carbonate platform transected by intrashelf basins. Bein (1976, 1977), Bein & Weiler (1976) and Sass & Bein (1982) considered the Carmel a transitional region, in which shallow marine carbonate platform of Israel passes to the west to a rudist reef-dominated rim, and further westwards into the deeper water laminites, turbidites and contourites of the Talme Yaffe Fm (Cohen 1971). This chapter addresses some open questions for the Cenomanian-Turonian: (1) How do the sedimentary cycles of the carbonate system of northern Israel (Chapter 2) organize into systems tracts and sedimentary sequences? (2) How do proximal-to-distal facies changes, and system-tract geometries, reflect the tectonic configuration of this part of the Arabian plate? (3) Can sequences found in northern Israel be correlated with major depositionalenvironmental events (palaeoceanographic/palaeoecological factors; global change) of the Tethyan realm? This chapter will demonstrate how depositional strikes in northern Israel, trending broadly W-SW ─ E-NE, reflect deviation from the S-N depositional strike prevailing southward along the Levant coast. These trends are shown to correspond to other large scale Mesozoic structural and depositional features across the northern Arabian plate.
3.2. Sequence stratigraphy and sedimentary facies configuration Figures 3.1 and 3.2 show the studied sections in northern Israel, their subdivision into systems tracts and sequences, bounding-surfaces, and chronostratigraphic assignation. Biostratigraphic and radiometric control are from the literature, mostly from Freund & Raab 1969; Lewy & Raab 1978; Lipson-Benitah et al. 1997; Segev et al. 2002. Three Cenomanian
41
Chapter 3: Sequence stratigraphy, N. Israel
sequences and a single Turonian sequence were defined. Complete descriptions of the sections used in this chapter are given in Appendix 2. 3.2.1. Sequence 1: Early-Mid Cenomanian 3.2.1.1. Albian–Cenomanian sequence boundary (Alb/Ce SB-1) The boundary between the Yagur dolomites, considered Albian (Lewy & Raab 1978), and the overlying indurated chalks of the Deir Hanna/Isfiyye Fms (Galilee/Carmel respectively; Fig. 1.2) is a discontinuity surface, described previously in northern Israel by Karcz (1959), Folkman (1969), Bein (1974), Kafri (1986) and Lipson-Benitah et al. (1997). Folkman (1969) and Kafri (1986) described pisolitic calcretes, monomictic conglomerates, oxidized Fe crusts, quartz grains and silicification and dedolomitization phenomena and concluded that this discontinuity surface represents sub-aerial exposure. Gardosh et al. (2006) associated this horizon with submarine canyon incision into deep-basinal deposits offshore. Biostratigraphic data based on ammonites and planktonic foraminifera (Avnimelech 1965; Lewy & Raab 1978; Lipson-Benitah et al. 1997) from chalks somewhat above this surface give an Early Cenomanian age; thus this surface is close to the Albian-Cenomanian transition. Pyroclastics overlying the surface in the Carmel were dated to 97.1±1.7 m.y., broadly Early Cenomanian (Segev et al. 2002). Regionally, this boundary was described as a sub-aerial erosion surface paved by conglomerates in central Jordan (Abed 1984), and as an emergent surface in Lebanon (Ferry et al. 2007). In central and southern Israel, the Alb/Ce SB-1 corresponds to a discontinuity characterized by a shell bed of Pycnodonte vesiculosa (Lewy & Weissbrod 1993; Braun & Hirsch 1994). Approximately coeval hiatuses and sea-level fall have been recorded elsewhere, e.g., in SW England (Simmons et al. 1991), the Anglo-Paris basin, and as far as Crimea, Kazakhstan, Turkmenistan and Iran (Gale et al. 1996), as well as from the North American western interior basin (Gröcke et al. 1998). 3.2.1.2. The Early Cenomanian TST (Ce TST-1) The Ce TST-1 corresponds to the lower-middle part of the Deir-Hanna Fm of the Galilee and to the Isfiyye and Beit-Oren Fms of the Carmel (Fig. 3.3). This system tract is composed of one high-order UC-7 deepening-upward cycle bounded at the base by the erosional discontinuity of the Alb/Ce SB-1, and at the top by Early Cenomanian maximum-flooding interval (Ce MFI-1). Ammonites and planktonic foraminifera suggest that the age of the Ce TST-1 is Early Cenomanian (Fig. 3.3). The Cenomanian TST-1 is composed of well-laminated, fine-grained, mostly calsisiltitic mudstones and wackestones with rare pelagic biota, with some pyrite and glauconitic grains
42
Chapter 3: Sequence stratigraphy, N. Israel
(FT-1 in Table 2.1; Pl. 4d). These basinal deposits originated from two sources: fine-grained bioclasts and carbonate mud originating from remote proximal environment, and pelagic carbonates from the overlying water column. Both were deposited on a hypoxic basin floor, presumably where the oxygen minimum zone (OMZ) impinged on the slope. Facies-thickness variations of Ce TST-1 across northern Israel, from Manara in the NE to Isfiyye in the SW, are simplified and scaled in Figure 3.4a. The Ce TST-1 is relatively thin and condensed in the NE Galilee (14m at Manara), thickens toward the NW Galilee (50m at Betzet), and reaches maximal thickness in the Carmel region (up to 120m; Bein 1974). Subsidence rate gradually increased towards the Carmel in the southwest. Similarly, the maximum-flooding interval (Ce MFI-1) at the top of the systems tract (see below) is of relatively shallow depth in the NE Galilee, based on occurrences of orbitolinid foraminifera, and deeper in the Carmel region in the southwest, with planktonic foraminifera and glauconite. 3.2.1.3. Early Cenomanian maximum-flooding interval (Ce MFI-1) The Ce MFI-1 consists of a limestone bed intercalated within the basinal chalk of the Deir Hanna Fm of the Galilee, and occurring as part of the ‘Beit Oren’ limestone bed within the chalk succession of the Isfiyye-Arkan Fms in the Carmel region (Fig. 3.1). This bed is a few centimeters to a few meters in thickness, overlying the basinal laminites of the Ce TST-1. Orbitolina sefini occurs in this bed in the NE Galilee (id, M. Simmons, 2007, pers. comm.), and Early Cenomanian ammonites and planktonic foraminifera occur below and above this bed in the Carmel, suggesting that it is still in the Early Cenomanian (Fig. 3.3). The facies of the Ce MFI-1 corresponds to basinal shell beds and skeletal lag concentrations of FT-7 and FT-8 (Table 2.1). In the Rakit/Beit Oren and Mt. Kedumim sections of the Carmel and Nazerath Hills (Fig. 3.1) the Ce MFI-1 appears as a highlybioturbated, cemented basinal shell bed (‘Beit Oren’ shell bed in the Carmel, Bein 1974), capped by glauconite (Weiler 1968; Lipson-Benitah et al. 1997). In the NW and central Galilee (e.g. Kziv-east, Betzet sections; Fig. 3.1) it occurs as skeletal or peloidal hiatal concentrations with bioeroded bioclasts and pelagic microfossils (FT-7 – Pl. 2c, Pl. 4e) or as packed or loosely packed shell-concentrations (Kziv-east section; FT-8 – Pl. 4f). In the proximal NE Galilee (Manara section) the Ce MFI-1 occurs as a limestone bed rich in flat orbitolinids, serpulids and large bivalves (FT-8 – Pl. 4b). The matrix of the Ce MFI-1 contains pelagic micro-elements, suggesting that it was formed in the basinal environment. Bioturbation, colonization by macrobenthos, and formation of lag concentrations by winnowing, reflect aeration and oxygenation of the Ce TST-1 hypoxic bottom, accompanied by decreased sedimentation.
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Chapter 3: Sequence stratigraphy, N. Israel
Facies configuration of the Ce MFI-1 can be interpreted as three strike-parallel zones on the basin floor (Fig. 3.4). The proximal facies zone is represented by the flat orbitolinid horizon of the NE Galilee. According to Vilas et al. (1995) and Hottinger (1997), flat orbitolinids are better adapted to water cloudiness, low illumination and increased water depth, than conical forms. Simmons et al. (2000) and Di Lucia et al. (2007) showed that flat orbitolinids characterize transgressive system tracts, with the flattest forms found around maximum-flooding surfaces. Thus, flat orbitolinids at MFI-1 at Manara indicate increased nutrients and phytoplankton at the surface, causing turbidity and low light levels on the sea floor. Other authors also attributed flat orbitolinids to increased water cloudiness, either due to terrigenous influx or eutrophic conditions in transgressive contexts (e.g. Immenhauser et al. 1999; Pittet et al. 2002; Al Juboury et al. 2006). The intermediate facies belt of MFI-1 is represented by the fragmental/peloidal hiatal concentrations of the Betzet, Kziv-west and Kziv-east sections. In this zone, weak bottomcurrents caused winnowing of TST-1 fines, leaving fragmental-bioclastic or peloidal lag concentrations. In the Kziv-east region, the oxic sea-floor was subsequently colonized by bivalves and bioeroding organisms. The lower slope/basin facies zone of the Ce MFI-1 occurs in the Nazareth Hills and Carmel. It is characterized by bioturbated and biodegraded macrobenthic shell material indicating aeration of the sea-floor, decelerated sedimentation and probably also increased flux of nutrients (cf. Kidwell 1986; Hallock 1988). Glauconite formed in this distal zone as hypoxia encroached (Porrenga 1967; Carson & Crowley 1993). 3.2.1.4. The Early-Mid Cenomanian RST (Ce RST-1) The Lower-Mid Cenomanian systems tracts of northern Israel include the MidCenomanian highstand (Ce HST-1) and the Mid-Cenomanian forced-regressive system tract (Ce FRST-1) (Figs. 3.1, 3.2). These system tracts are combined into a ‘Regressive System Tract’ (RST) sensu Embry (2002) (Ce RST-1) for the purpose of time-boundary definition and palaeogeographic interpretation, but will be also addressed separately. The Ce RST-1 is bounded below by the Early Cenomanian MFI (Ce MFI-1) and above by the Mid-Cenomanian sequence boundary (Ce SB-2; Fig. 3.2). This succession corresponds to the upper Deir Hanna Fm and the mid-ramp/peritidal facies of the Sakhnin Fm in the Galilee, and to the Zikhron and Arkan Fms in the Carmel region. The presence of Orbitolina sefini in Ce MFI-1 at Manara (NE Galilee) indicates that the Ce RST-1 begins in the Early Cenomanian. Radiometric dating of pyroclastics at the Ce SB2 in the southern Carmel, (95.4±0.5 m.y; Segev et al. 2002) and occurrence of the ammonites Calycoceras sp. (Ident. Z. Lewy, pers. comm. 2007) and Protacanthoceras sp. (Freund 1958) of the A. jukesbrownei zone (Kennedy & Jolkičev 2004) above the Ce SB-2
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Chapter 3: Sequence stratigraphy, N. Israel
(above the Sakhnin Fm) in the Galilee, suggest that the Ce RST-1 terminated in the MidCenomanian A. rhotomagense Zone. The Mid-Cenomanian HST (Ce HST-1; Figs. 3.1, 3.2, 3.3) is the lower part of Ce RST-1. In most of the Galilee and southern Carmel, Ce HST-1 is a progradational type-1 low-order shallowing-upward cycle (CC1 – Fig. 2.10a), 60 to 200m in thickness (e.g. Pl. 5a). The midramp zone of this cycle (FT-12 in Table 1, Pl. 5a) is highly-productive, providing lime-mud both downslope to the outer ramp/basin, and by wave and current transport, also to the inner ramp. The sharp boundary in the type-1 low-order cycle, separating lower basinal laminites from overlying bioturbated mid-ramp facies reflects the transition from the zone of impingement of the oxygen minimum (OMZ) to the mixing zone. This boundary represents a marked ecological transition between non-productive sea-floor in the OMZ to highlyproductive mid-ramp in the mixing zone, a transition that could generate distally-steepened ramp morphology. The Mid-Cenomanian FRST (Ce FRST-1) (Figs. 3.2, 3.3) is the upper part of Ce RST-1. The Ce FRST-1 is represented by mass-transport deposits in the Sakhnin Fm. Their base corresponds to the 'basal surface of forced-regression' (BSFR sensu Hunt & Tucker 1992, 1993). The upper boundary corresponds to the Mid-Cenomanian sequence boundary (Ce SB-2), at the transition to overlying lowstand or transgressive deposits (Fig. 3.2). Mass-transport features of the Ce FRST-1 include debrite breccias, channelized and non-channelized turbidites, and shear-planes and translational slides, all described and discussed in detail in Chapter 4. Facies, geometries and sedimentary configuration of the Ce RST-1 (HST-1 and FRST-1) are presented in Figure 3.5. The CC1 cycle of the Ce HST-1 distally-steepened ramp (Fig. 2.10a) is present in most of the Galilee and in the southern Carmel region (Fig. 3.5a; examples in Pl. 5a-c). Blocks and skeletal grains derived from the ramp were redeposited as debris-flow breccias, deformed and sheared slabs, and channelized and non-channelized turbidites in the western Galilee. Toward the northeast Galilee the peritidal HST ramp passed distally into a mid-ramp productive zone (Fig. 3.5a,b). In the southern Carmel, the MidCenomanian peritidal ramp of the Zikhron Fm (Pl. 5b,c) passed toward the north into MidCenomanian basinal deposits of the Arkan Fm (Figures 3.2 and 3.5c). This transition is associated with ramp disintegration and formation of synsedimentary breccias. These facies changes are reflected in the transition from the Hotem-Carmel to the Rakit-Beit Oren sections (see also Segev & Sass 2006). NE-SW thickness-facies changes in the Ce RST-1 are summarized in Figure 3.5d. The peritidal system of the Galilee passes gradually toward the south into thicker and deeper outer-ramp deposits with slumps and debrites (FT-3 at Mt. Kedumim). Farther to the south, this outer ramp facies shallows once more into the peritidal zone, represented by the Hotem-
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Chapter 3: Sequence stratigraphy, N. Israel
Carmel section of the southern Carmel region. A more distal expression of this transition is shown in Figure 3.5c. In summary, the facies distribution in the Ce RST-1 indicates mid-ramp to peritidal palaeohighs in most of the Galilee and southern Carmel. The slope of the shallow ramp of the Galilee faced north, west and south. The slope of the shallow ramp of the southern Carmel faced north (and west; Bein 1974). These two highs were separated by a deeper zone of increased subsidence that extended roughly E-W across the central Carmel to the southern Galilee. 3.2.1.5. The Mid-Cenomanian sequence boundary (Ce SB-2) The Mid-Cenomanian sequence boundary separates the Ce RST-1 from the overlying TST or LST (Figs. 3.1-3.3). Pyroclastics above the Ce SB-2 in the southern Carmel (top Zikhron Fm) were dated to 95.4±0.5m.y (Segev et al. 2002), corresponding to the MidCenomanian A. rhotomagense ammonite Zone. The ammonite Calycoceras sp., recovered above the sequence boundary at Manara, supports this age (Fig. 3.3). The Ce SB-2 has proximal and distal expressions: Proximal expressions of the Ce SB-2 were recorded in the eastern Galilee (Manara section and Kadarin region, Fig. 1.1). Ce HST-1 at Manara corresponds to a dolomitized midramp rudist-gastropod facies (FT-12), 20m thick, developed on the north-facing slope of the Galilean ramp (Fig. 3.5a). The upper few centimeters of this mid-ramp facies at Manara are distinguished by the presence of sphaeroidal calcitic/silicic or partly-silicified vadose pisoids, empty or infilled dissolution vugs, pore-filling speleothems (flowstones), and concentric pendant calcites (Fig. 4.4). These vadose diagenetic features indicate subaerial exposure of the Ce HST-1 ramp. Further evidence for sub-aerial exposure from the proximal ramp are teepee structures reported by Bogoch et al. (1994) from the Sakhnin – Bina Fm transition in the eastern Galilee (Kadarim region, Fig.1.1). The subaerial unconformity at Manara is paved by a thin reddish-black ferruginous pavement (FT-9) overlain by basinal TST chalks. This mineralized crust terminated the final phase of void creation and fill, by sealing open cavities and inhibiting water percolation. Rapid sea-level rise at the TST phase, and flooding of the sequence boundary led to sediment starvation, condensation and mineralization. Longman (1980) stated that preservation of vadose cements requires rapid transgression or subsidence, supporting the claim that the ferruginous crust is a transgressive feature. The discontinuity interval terminating the Ce HST-1 ramp at Manara, represents therefore sub-aerial unconformity combined with transgressive surface (Fig. 3.2). Distal expressions of the Ce SB-2 are at the western Galilee and Carmel (Yanuch and Rakit/Beit-Oren sections, Fig. 3.1). On the distal ramp slopes of the western Galilee and
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Chapter 3: Sequence stratigraphy, N. Israel
central-northern Carmel the Ce SB-2 is expressed by truncation of Ce HST-1 basinal sediments (e.g. Fig. 3.5c). Bein (1974) showed this truncation at the top of the chalk complex (Arkan Fm) of the Carmel. A coeval truncation was found in the western Galilee at the transition from upper Deir Hanna chalk laminites to bioturbated limestones of the Yanuch Fm (Yanuch section, Pl. 1c). 3.2.2. Sequence 2: late Mid - to Late Cenomanian 3.2.2.1. The Mid-Cenomanian lowstand system tract (Ce LST-2) Lowstand deposits are confined to the Sakhnin Fm of the distal Sheikh-Danun and Kzivwest sections (Fig. 3.2). They are placed above debrites of the Ce FRST-1 and represent progradation above the Ce SB-2 at the first stage of relative sea-level rise (cf. Hunt & Tucker 1993). The lowstand tract is composed of UC-3 subtidal shallowing-upward cycles (Fig. 2.8) and is overlain by a transgressive surface at the transition to overlying laminites of the Ce TST-2 (below). 3.2.2.2. The late Mid- to Late Cenomanian transgressive system tract (Ce TST-2) and the Late Cenomanian maximum-flooding interval (Ce MFI-2) The Ce TST-2 corresponds to the lower Yanuch Fm of the western Galilee (Yanuch section), lower Bina Fm of NE Galilee (Dishon section) and the bedded chalky part of the Yanuch Fm at Manara (considered Deir Hanna Fm by Kafri 1991; Sneh & Weinberger 2003) (Figs. 3.1-3.3). The Ce TST-2 is composed of limestones and chalks bounded at the base by the Ce SB-2 /transgressive surface and at the top by a maximum-flooding interval (Ce MFI2). The ammonites Calycoceras sp. and Protacanthoceras sp. of the late Mid-Cenomanian Acanthoceras jukesbrownei zone (Kennedy & Jolkičev 2004) were recovered from the lower part of this tract (Freund 1958; Glikson 1966). Proximal sections are more poorly constrained. Benthic foraminifera from the lower part of the proximal Ce TST-2 at Dishon (Fig. 2.5; cycle DO2 in Fig. 3.1) belong to the Pseudorhapidionina dubia TRZ spanning the Middle- to Late Cenomanian (Aguilera-Franco 2001, 2005). Thus, correlation of proximal- to distal sections is based on a combination of facies considerations and stacking patterns, supported by low-resolution biostratigraphic correlation. In both proximal and distal settings, the Ce TST-2 begins with slow carbonate deposition or submarine omission, expressed by the heavily micritized grains, microbial lumps, coated grains and oncoids at the base of the Ce TST-2 at Dishon (FT-18 at base UC-8 deepeningupward cycle), and by the ferruginous pavement at Manara (FT-9). Facies configuration of the Ce TST-2 is in Figure 3.6. In the proximal sections (e.g. Betzet, Dishon; Fig. 3.1) the Ce TST-2 is thin, composed of shallow subtidal UC-6 cycles,
47
Chapter 3: Sequence stratigraphy, N. Israel
and lagoonal to outer-ramp UC-8 deepening upward cycle. The Ce TST-2 becomes much distal toward the Carmel region where it is composed of fine grained pelagites (FT-6 at OrenValley section). Compared to the thin and proximal UC-8 cycle at Dishon, the Ce TST-2 is much thicker and deeper in the northern Galilee, where it is composed of laminated calcisiltitic chalks (FT-1 at Kedesh and Manara). The rapid thickening from Dishon in the south to Kedesh and Manara in the north (Fig. 3.1) reflects increased subsidence to the north, possibly due to block down-faulting (Fig. 3.6a). The basinal zone of the northern Galilee extends westward along the depositional strike, to the Sarach Valley and Adamit regions (Fig. 3.6b). The basinal facies of the Carmel region is relatively condensed, and is more pelagic. The maximum-flooding interval, terminating the Ce TST-2 (Ce MFI-2), is marked only on the Galilee paleohigh. It forms well laminated pelagites (FT-1) at the termination of the relatively proximal UC-8 and UC-10 deepening-upward cycles (Fig. 2.9) of Dishon and Yanuch. 3.2.2.3. Late Cenomanian regressive system-tract (Ce RST-2) Late Cenomanian system tracts of northern Israel are the Late Cenomanian highstand (Ce HST-2) and the Late Cenomanian forced-regressive system tract (Ce FRST-2) (Fig. 3.1, 3.3), combined here into a "regressive system tract" (Ce RST-2) sensu Embry (2002), for the purpose of time-boundary definition and palaeogeographic interpretation. Ce RST-2 is bounded at the base by the Late Cenomanian maximum-flooding interval Ce MFI-2, and at the top by the Late Cenomanian sequence boundary (Ce SB-3 of western Galilee, described below) or by an amalgamation of the Ce SB-3 with the younger latest Cenomanian sequence boundary (Ce SB-4) (see below). The Ce RST-1 is Late Cenomanian in age (Fig. 3.3). The Ce HST-2 at the beginning of the Ce RST-2 corresponds to the Yanuch and Bina Fms in the Galilee, and to part of the Muhraqa Fm at Oren Valley section in the western Carmel (Fig. 3.3). The presence of latest Mid-Cenomanian ammonites in the previous TST, and occurrences of the benthic foraminifera Cisalveolina fallax in the upper part of the Ce HST-2, suggest that this system tract is Late to latestCenomanian in age. In the more distal western Galilee, Ce HST-2 is topped by a basal surface of forced regression (Ce BSFR-2) at the transition to overlying forced-regression system tract (Ce FRST-2) (Fig. 3.1). Ce HST-2 is composed of a type-2 lower order cycle (CC2) forming stressed, low productivity homoclinal ramp (Fig. 2. 10b). A cross section across this tract is in Figure 3.7a. In the central and NW Galilee it is thin, and composed of peritidal UC-1 cycles (e.g. at Betzet). This peritidal zone passes into thicker and deeper subtidal UC-3 cycles (outer-ramp to shoreface) toward the SW (Yanuch, Beit Ha'Emek) and NE (Dishon). These thickness--facies changes show that the central Galilee was peri-tidal and elevated, while to
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Chapter 3: Sequence stratigraphy, N. Israel
the S-SW and N-NE the ramp was deeper and more rapidly subsiding. Further SW in the Carmel (Oren Valley section, Fig. 3.1), this system is composed of deep basinal fine grained bioclastic pelagic ooze (FT-6). The Late Cenomanian forced regression system tract (Ce FRST-2) of the western Galilee and Carmel is the main subject of Chapter 5 and is briefly discussed here. This system tract is at the upper part of the Ce RST-2 (Figs. 3.1, 3.3), corresponding to parts of the Yanuch and Muhraqa Fms. Facies of the Ce FRST-1 corresponds mainly to calcarenitic slope clinoforms in the Galilee (FT-5) and to turbidites (FT-2) in the Carmel range (see Table 1). FT-5 occurs in the Galilee in an isolated clinoform set 35-100m in thickness, with inclinations of 22°-35° to the W-SW. The clinoforms are spectacularly exposed from Kishor to the Shagor canyon 5 kilometers to the south (Fig. 1.1). The base of the Kishor–Shagor clinoform set is sharp and erosive, and is considered to be the Late Cenomanian basal surface of forced regression (BSFR-2 in Fig. 3.7a). The top of the clinoform set is eroded by sub-aerial unconformity and then covered by Cenomanian limestones of the Pelech sequence (see below). The Kishor-Shagor clinoform set rapidly becomes steeper and thicker toward the south. Thickness is 0 at Yanuch–Beit-Ha’Emek, 35m at Kishor and > 100m at Hamra (Figs. 3.1, 3.7a). The isolation of the clinoform set, and the exposure surface at its top, are criteria for recognition of forced-regression (e.g. Plint 1988; Hunt & Tucker 1993; Posamentier & Morris 2000). Structural control on deposition may be inferred by the rapid thickness changes of the clinoform set. These relations are explained by syndepositional downfaulting creating accommodation space immediately filled by thick, steeply-dipping calcarenitic clinoforms (Fig. 3.7a). The calcarenite grains originated from shoreface erosion of UC-3 cycles on the nearby shelf (Beit-Ha’Emek, Yanuch). They were transported toward the S-SW across the fault scarp. The toe of this slope can be found in the distal Carmel ~40km to the S-SW, where 25-30m of graded, rippled turbiditic calcarenites occurs in the middle part of the Muhraqa Fm (Fig. 3.7a). 3.2.3. Sequence 3: Late Cenomanian 3.2.3.1. Late Cenomanian sequence boundary (Ce SB-3) and Pelech sequence In the Hamra Valley of western Galilee, near the village of Pelech (PL in Fig. 1.1), a succession composed of 35m of well-bedded limestones (Pl. 6a) lies above the Ce HST-2 and is bounded between two discontinuities (Fig. 3.8). This sequence corresponds to outerramp and mid-ramp facies of the FT-10 and FT-17 (Table 2.1). The presence of Cisalveolina fallax in the underlying Ce HST-2 suggests that this sequence corresponds to the latest Cenomanian M. geslinianum Zone (cf. Aguilera-Franco et al. 2001).
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Chapter 3: Sequence stratigraphy, N. Israel
The basal discontinuity is a ferruginous marl bed paved by decimeter-scale iron concentrations root casts (Pl. 6b). This surface is the Late Cenomanian sequence boundary (Ce SB-3). The upper discontinuity is the latest Ce sequence boundary (Ce SB-4; discussed below). Detailed lateral relations between the Pelech sequence and the adjacent units are fully discussed in Chapter 5, but this locally-exposed sequence lies above the Ce HST-2 and onlaps the Kishor–Shagor clinoform set in the west (Fig. 3.8). The basal discontinuity, i.e. the Ce SB-3, reflects sea-level fall at the end of the Ce RST-2, and the Pelech sequence represents a subsequent rise. Part of the Pelech sequence was removed by erosion during latest Cenomanian exposure event (Ce SB-4) and subsequent Early Turonian transgression. As a result, the Ce SB-3 locally amalgamates with Ce SB-4 (e.g. Fig. 3.8). 3.2.3.2. The latest Cenomanian sequence boundary (Ce SB-4) The Ce SB-4 is at the transition from the Cenomanian to the Early Turonian in northern Israel. It is marked by sub-aerial exposure and a hiatus in the Galilee, and by submarine omission in the Carmel region. Pedogenic and karst features (FT-20 Table 1) occur at the top of the Ce RST-2 and ‘Pelech sequence’ of the western Galilee (Fig. 3.8). The forced-regressive clinoforms of the Galilee are truncated by an irregular karst surface paved by ferruginous concentrations (Pl. 6c). This surface is covered with a thin reddish crust with pedogenic pisoliths, circumgranular cracks and alveolar septal fabric (Pl. 5d). This crust is the remnant of calcrete terminating Ce RST-2, most of which was removed by submarine erosion in the course of the following transgression. Similarly, the top of the ‘Pelech sequence’ of the western Galilee is an irregular erosional surface reminiscent of lapiés (Pl. 5e; Fig. 3.8). These features indicate sub-aerial exposure and hiatus prior to Early Turonian drowning. Generally in the Carmel region, the Cenomanian–Turonian transition is within the Muhraqa Fm, which has caprinid rudists in its lower part and hippuritids in the upper part (Buchbinder et al. 2000). In the Isfiyye section of the Carmel (Fig. 1.1, 3.9), the Ce SB-4 was found as an erosion surface paved by submarine ferruginous crust (FT-9) at the top of Late Cenomanian forced-regressive debrites (Pl. 6d,e). This discontinuity is topped by hippuritid grainstones of Turonian age. Buchbinder et al. (2000) defined the Late Cenomanian sequence boundary in northern Israel between the Late Cenomanian Yanuch/lower-Muhraqa and overlying Yirka/upper Daliyya Fms (Fig. 1.2). A prolonged hiatus spanning 1.25my was indicated, extending from the upper part of the Late Cenomanian N. vibrayeanus Zone, until the middle part of the Early Turonian W. coloradoense Zone (Zone T4 of the Turonian of Israel; Freund & Raab 1969).
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Chapter 3: Sequence stratigraphy, N. Israel
In the western Galilee, the benthic foraminifera Cisalveolina fallax was recovered from below the sequence boundary (Fig. 3.3). Saint-Marc (1974) defined a C. fallax zone spanning the latest Cenomanian to Early Turonian. Ammonites of the Turonian zone T4 (Freund & Raab 1969) occur ~20 above the sequence boundary. The hiatus in this region is therefore probably not as long as the previous estimation. 3.2.4. Sequence 4: Early –Middle Turonian 3.2.4.1. Turonian transgressive system tract (Tu TST) The Tu TST (Figs. 3.1, 3.3) corresponds to the upper Muhraqa and Daliyya Fms in the Carmel, and to the uppermost Yanuch Fm, Yirka Fm, and part of the Bina Fm in the Galilee. It is bounded at the base by the Ce SB-4 and the Early Turonian transgressive surface and at the top by a thick maximum-flooding interval. The Tu TST is composed of deep basinal facies, turbidites, fragmental hiatal concentrations, hypoxic laminites and shoreface deposits (FT-1, FT-2, FT-7, FT-11, FT-15; Pls. 1g, 3c, 3f-I, 6f-h). In the Carmel, the Tu TST begins with UC-12 cycles (Fig. 2.9; described by Buchbinder et al. 2000), deepening upwards from the lower shoreface to the basin. These cycles pass toward the southern Carmel into shallower UC-8 cycles (Fig. 2.9) with shallower lagoonal aspects. This transition (Fig. 3.9) indicates shallower conditions in the southern Carmel. The Carmel cycles (UC12 and 8) pass upwards into basinal marls of the Daliyya Fm. In the western Galilee, the Tu TST begins above the latest Cenomanian sequence boundary (Ce SB-4) with a CC6 cycle associated with turbidites (e.g. at Hamra, Fig. 3.1; Pl. 6f-h), and with FT-15 retrogradational shoreface grainstones farther in the north (Yanuch, Beit-Ha’Eemek; Pl. 1e). In other parts of the Galilee Tu TST corresponds to hypoxic mid- to outer-ramp deepening-upward cycles UC-7, 9 and 11 of the Bina and Yirka Fms (Fig. 2.9). Geometry and facies of the Tu TST are shown in Figure 3.7a. In the central and NW Galilee the Tu TST is a condensed mid- to outer-ramp UC-9 cycle. At Dishon in the NE, it is represented by a CC-7 low-order deepening-upward cycle composed mainly of mid/outer ramp laminites (FT-11). At Yanuch in the SW the Tu TST is composed of a CC-6 low order deepening-upward cycle, and farther to the SW, at Hamra, it is composed of 80m of basinal marls (FT-1; Yirka Fm). Facies configuration of the Tu TST is in Figure 3.7c. The thin UC-9 cycle of the Galilee reflects condensation on a Galilean palaeohigh. This zone was bounded to the north and south by rapidly subsiding deeper regions. Outer-ramp laminites of the NE Galilee onlapped the Galilee from the north, and basinal marls onlapped the steep clinoformic slope of the Galilee from the south.
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Chapter 3: Sequence stratigraphy, N. Israel
3.3. Discussion 3.3.1. The mid-Cretaceous tectono-sedimentary framework of northern Israel The 'Levantine hinge- belt’ (Gvirtzman & Klang 1972; Bein & Gvirtzman 1977) extends from the southern coastal-plain of Israel to the Carmel region. Northwards, in the Galilee and southern Lebanon, the directional trend and facies effects of the 'Levantine hinge-belt’ were not reported. Walley (1998) shows a continuation of a similar hinge-belt in northern Lebanon along the western Lebanon flexure, and considered the Galilee and southern Lebanon an interruption of the hinge-belt trend, presumably reflecting the southwestern extension of the Palmyride basin.
The data and sequence stratigraphic interpretations presented above
contribute materially to this picture. The north-south depositional strike of the hinge-belt shifts to the E-NE, beginning north of the southern Carmel region. This shift and an elevated Galilean-high bounded by subsiding basins both to the north and south are expressed in system tracts Ce TST-1, Ce HST-1, Ce TST-2, Ce HST-2, Ce FRST-2 and Tu-TST (Figs. 3.4-3.7). The southern Carmel was also elevated but the central-northern Carmel was deeper. Data from the Ce HST-1 in Mt. Kedumin and southern Carmel (Figs. 3.1, 3.5) and from the Tu TST of the southern Carmel (Figs. 3.1, 3.5, 3.9), suggest that an E-NE trending basin/depocenter in the central-northern Carmel extended towards the southern part of the Galilee. North of the Galilee paleohigh, the depositional strike shifts once more towards the NE, based on data from the Ce HST-1, Ce TST-2, Ce HST-2 and Tu TST in that region. The structural control on facies and systems-tract geometry extends beyond this immediate region. Data of Saint-Marc (1972, 1974) indicate a thickening trend of the Cenomanian succession towards central Lebanon in the north (Fig. 6.2a), apparently a continuation of the subsidence trend beginning in the northernmost Galilee. The W-SW – E-NE structural trend of northern Israel corresponds to other large-scale depositional trends in the Levant. The SW-NE Palmyrid trend in Syria was active at least since the Triassic and possibly even Late Palaeozoic (Ponikarov et al. 1967; Brew et al. 2001). Similarly, Ferry et al. (2007) showed that depositional strikes of the Early Aptian Jezzine Fm and Late Albian Niha Fm of Lebanon trend E-NE. This trend is also sub-parallel to the depositional strike of the northern margin of the Arabian Plate in the mid-Cretaceous (SE Turkey, Carter & Gillcrist 1994), and to depositional strikes from west-central Jordan (Schulze et al. 2005) and northern Sinai (Bachmann & Kuss 1998; Bauer et al. 2003). These depositional trends have not previously been considered as dominating the mid-Cretaceous palaeogeography in Israel. The integration of all these trends into a complete regional framework for the Mesozoic of the northern Arabian margin may at this point be premature. Clearly, a comprehensive
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Chapter 3: Sequence stratigraphy, N. Israel
regional compilation of thickness and facies data is necessary to complete the tectonosedimentary configuration of the northern Levant margin in the mid-Cretaceous. 3.3.2. Controlling mechanisms and correlation across the Arabian Plate and beyond 3.3.2.1. The beginning of Cenomanian sequence-1 A plate-wide earliest Cenomanian maximum-flooding surface K120 was described by Sharland et al. (2001) at 98 m.y. It was recognized in SE Turkey, Lebanon, eastern Syria, Iraq, Iran, Kuwait, Saudi Arabia, Qatar, UAE, Oman and Yemen. In the Arabian Plate, the K120 interval largely corresponds to organic-rich bituminous shales, or carbonate mudstones and wackestones. A section consisting of organic-rich shales from the base of the Natih-E unit in Oman was chosen as a 'reference section' of this flooding event. The Early Cenomanian maximum-flooding episode of northern Israel (Ce MFI-1) correlates with K120 (Fig. 3.3). This is indicated by the presence of Orbitolina sefini in this interval at Manara and the occurrences of Early Cenomanian ammonites below and above it in the Carmel region. However, the Ce MFI-1 has a different sedimentological expression than the K120. It corresponds to the sharp facies transition from basinal stagnant laminites of the Ce TST-1 into bioturbated, bioeroded shell beds, bioclastic peloidal hiatal concentrations, glauconite-rich horizons and a flat-orbitolinids bed. Weak current winnowing and reworking and sea floor aeration are indicated. A similar shell bed occurs within the well stratified basinal succession of the Ein Yorkeam Fm in the northern Negev (Fig. 3.3). In north-central Jordan, Schulze et al. (2003) correlated the K120 MFS to the surface separating Early Cenomanian marls of the Naur-A unit, from the shallower rudist buildups and shallowing upward cycles of the Naur-B unit. Other Early Cenomanian maximum-flooding episodes from Europe (Caus et al. 1997; Wilmsen 2008) are somewhat younger than the K120 of the Arabian plate but may provide some explanations for the sedimentological differences between the shelly-bioclastic Ce MFI-1 of northern Israel and the bituminous K120 described from the Arabian plate. In the Sopeira Fm of SE Spain, Caus et al. (1997) showed basinal deposits capped by a maximumflooding bed near the transition from the Early- to the Middle Cenomanian. Despite apparent diachrony, this is also an Early Cenomanian maximum-flooding event with nearly identical facies to the Galilee, e.g., abundant orbitolinids occurring proximally. Another maximumflooding event was described by Wilmsen (2008) in the Early Cenomanian mid-M. dixoni Zone
of
northern
Germany.
It
corresponds
to
inoceramid
shell
beds
(Schloenbachia/Inoceramus virgatus bioevent) widely distributed in NW Europe. Despite time discrepancy, the Early Cenomanian shell beds of Israel and Europe were affected by similar mechanism of bottom reworking near or below storm wave base.
53
Chapter 3: Sequence stratigraphy, N. Israel
Sharland et al. (2001) attributed the K120 event in the Arabian plate to eustatic rise and local subsidence. The Cenomanian maximum-flooding intervals in Israel and Europe, however, record an additional, paleo-environmental control. 3.3.2.2. End of Cenomanian sequence-1 The Early to Mid-Cenomanian highstand systems tract of northern Israel (Ce HST-1) represents a progradational distally-steepened ramp system. In most of the Galilee and southern Carmel it is reflected by a type-1 lower order cycle (CC1) with stacking patterns typically expressed by the Sakhnin and Zikhron Fms (Fig. 2.10a). This low order cycle reflects filling of the accommodation space to sea-level (peritidal depths). This is a consequence of two synchronously operating factors: (a) efficient carbonate mud supply from a mid-ramp rudist-gastropod production zone (FT-12 in Table 2.1), feeding both the upslope, inner-ramp belt, and downslope and basin areas, and (b) reduction of the accommodation space due to eustatic fall (see discussion below). The development of a mollusk-dominated high carbonate-production zone on the mid-ramp indicates mesotrophic conditions. Termination of that cycle was first by gravity collapse (Ce FRST-1), and subsequently by ramp emergence (Discussed in Chapter 4). This Early-Mid Cenomanian regressive event of northern Israel can be identified in the northern Negev, southern Israel (Fig. 3.3). Lewy (1990) described an Early to MidCenomanian regressive event (Zafit Fm), bounded at the top by subaerial exposure surface (Lewy & Avni 1988) and overlain by transgressive deposits of the Avnon Fm. Age, and facies stacking pattern of the Zafit Fm are similar to those of the type-1 lower order cycle of northern Israel (Fig. 2.10). An Early to Mid-Cenomanian regressive event was also recorded from other locations in the Arabian platform (Fig. 3.3) and Europe. In west central Jordan, Schulze et al. (2003, 2004) considered the rudist-bearing shallowing-upward cycles of the Naur-D unit as a regressive HST. The Naur-D is capped by a bored ferruginous hardground (sequence boundary CeJo2) and is covered by the drowning succession of the Fuheis Fm. Philip et al. (1995) recognized a Lower- to Mid-Cenomanian shallowing-upward HST pattern in the EarlyMiddle Cenomanian Natih-E unit of the Oman Platform. Van Buchem et al. (2002) assigned this interval to sequence-1 of the Cenomanian, terminated by a Mid-Cenomanian exposure surface (type-1 sequence boundary) 10m below top Natih-E unit and Grélaud et al. (2006) associated this boundary to eustatic fall, incisions and deposition of forced-regressive wedges. Caus et al. (1997) considered the Mid-Cenomanian part of the Santa-Fe limestones of NE Spain as a progradational system, redeposited as breccias in the basin, and topped by a sub-aerial Type-1 sequence boundary at the Middle/Late Cenomanian boundary.
54
Chapter 3: Sequence stratigraphy, N. Israel
In the deeper marine environments of the Anglo-Paris basin, Robaszynski et al. (1998) described a major fall in sea-level in the late Early Cenomanian, in the later part of the M. dixoni Zone, followed by a strong transgression in the earliest A. rhotomagense Zone, somewhat earlier than the late Mid-Cenomanian transgressive episode recorded in northern Israel. Tröger (2003) showed a Mid-Cenomanian hiatus in a basinal succession in Saxony. Hancock (2003) reported a Mid-Cenomanian ‘eustatic low’ associated with 'missing succession', submarine erosion (erosional channels and debrites) and condensation' (C. inerme Zone to early T. acutus Zone), from SW and eastern England. Mid-Cenomanian hiati were also reported by Hancock (2003) from Northern Ireland, NW France and Germany. Baraboshkin et al. (1998) reported a Mid-Cenomanian stratigraphic hiatus and erosional surface from Crimea, northern Caucasus, the Russian platform, peri-Caspian region and Mangyshlak basin, where Middle and Late Cenomanian deposits are strongly condensed or missing. Hancock (2003) recognized these hiati also in the western interior seaway of America. The pattern of sedimentation in the late Early- to Mid-Cenomanian in marginal marine environments of the Arabian plate and parts of Europe is therefore regressive-progradational, terminating by subaerial or submarine omission. In the deep environments sedimentation is associated with condensation and hiati, and redeposition of shallow-water deposits (northern Israel, Spain and England). Hancock & Kaufman (1979) and Hancock (2003) associated this Mid-Cenomanian regressive event with eustatic fall. Van Buchem et al. (2002) attributed this regressive event and the following sequence boundary also to eustatic fall, also responsible for the lowstand wedges reported by Azzam (1997) from the subsurface of Abu-Dhabi. 3.3.2.3. The beginning of Cenomanian sequence-2 Sequence-2 of northern Israel begins with transgressive event Ce TST-2 in the late MidCenomanian. The age of its top is not well constrained. This interval is represented by starvation, non-deposition and proximal deepening-upward cycles on the Galilean palaeohigh. It was terminated by laminated pelagites (FT-1) reflecting hypoxia and increased distality (Ce MFI-2). This transgressive event can be correlated across Israel and the Arabian plate (Fig. 3.3). In the northern Negev, this succession begins above the Ce SB-2 (top Zafit Fm) with the deposition of aerated outer-ramp nodular limestones of the Avnon Fm indicating platform drowning. As in the Galilee, the transgressive succession terminates with a maximumflooding bed composed of laminated chalk with abundant small calcispheres and planktonic foraminifera (Wald 2004). The last occurrence of orbitolinids is just above it, suggesting that Ce MFI-2 is still Mid-Cenomanian in age.
55
Chapter 3: Sequence stratigraphy, N. Israel
In west-central Jordan, the equivalent drowning succession starts with dark bituminous marls and limestones of the Mid-Cenomanian Fuheis Fm above the CeJo2 sequence boundary (Schulze et al. 2003). This transgressive event may correspond with the Ce TST-2 of northern Israel but the overlying progradational sequence until CeJo3 (base and top Karak limestone) was not identified in northern Israel. The biostratigraphic data (Fig. 3.3) suggest that the Ce MFI-2 corresponds to the maximum-flooding in mid-sequence 4 of Schulze et al. (2003) in west-central Jordan. In the Oman Platform, the Mid-Cenomanian transgression is recognized above shallow-water lowstand deposits at the upper Natih-E unit with maximumflooding within the lower Natih-D unit (Sharland et al. 2001; Van Buchem et al. 2002). This MFS corresponds to the calcareous shales of the K130 MFS of the Arabian plate and may correspond to the laminated pelagites of the Ce MFI-2 of northern Israel. This Mid-Cenomanian drowning succession at the base of sequence 2 in northern Israel represents the beginning of a lengthy episode of eutrophication, spanning the entire Cenomanian and Early Turonian (cf. Buchbinder et al. 2000). Eutrophication is expressed by blooms of pithonellid calcispheres (dinoflagellate cysts), phosphatic grains, intense bioerosion of molluscan bioclasts, silicic cements especially in micro-borings, and condensation phenomena associated with reduced photic penetration. Compared to the much more efficient carbonate production of the thick underlying Ce RST-1, the Ce TST-2 is highly-condensed (Fig. 3.6a), and no ‘carbonate factory’ was established. In the Cenomanian of the Oman Platform (Van Buchem et al. 2002), increased nutrients and clay dispersion are associated with platform flooding. Similarly, eutrophication controlled the stratigraphic architecture and facies of sequence-2, and seemingly the beginning of sequence-3, of Schulze et al. (2003) in central Jordan (Fuheis Fm). Schulze et al. (2004) invoked both sealevel rise and deposition of dysoxic bituminous successions in local basins to explain the Jordanian succession, while Van Buchem et al. (2002) advocated increased clay injection for water cloudiness. 3.3.2.4. Latter part of Cenomanian sequence-2 Late Cenomanian changes in northern Israel are expressed in the Ce RST-2 and the overlying Pelech sequence (Figs. 3.2, 3.3). Depositional processes were affected both by sea-level fluctuations and localized tectonic movements. Late Cenomanian aggradation began with the establishment of a CC2 cycle reflecting the establishment of a homoclinal ramp (Ce HST-2). Sea level fall and normal faulting in the Galilee led to the transformation of this ramp into an open non-rimmed shelf with a steep clinoform-bearing slope facing S-SW, and with toe-of-slope turbidites distally in the Carmel region (Ce FRST-2; Fig 3.7). This shelf-slope system was exposed for the first time during
56
Chapter 3: Sequence stratigraphy, N. Israel
the Late Cenomanian as evidenced by the Late Cenomanian sequence-boundary (Ce SB-3; Fig. 3.3). Shortly afterwards, the exposed shelf-slope system of the western Galilee was drowned below the storm wave base, forming the drowning succession of the Pelech sequence (Figs. 3.3, 3.8), partly removed by erosion in a second phase of emergence in the latest Cenomanian (Ce SB-4). The Late Cenomanian sequence boundary (Ce SB-3) is coeval with subaerial exposure recorded by Voigt et al. (2006) in the M. geslinianum Zone of northwest Europe, correlated by them to other surfaces in the Anglo-Paris basin, northeast Germany, Crimea, Ukraine, Mangyshlak (Kazahstan) and SE India. As in the Galilee, this surface was subsequently submerged, but a second exposure surface, as found in northern Israel (Ce SB-4) was not reported. Equivalents of the Late Cenomanian sequence boundary, the Pelech sequence and the latest Cenomanian sequence boundary of northern Israel are recorded from Wadi Feiran of Sinai. In this region, Kassab & Obaidalla (2001) show a basal Late Cenomanian erosional surface truncating part of the M. geslinianum ammonite zone, overlain by a 7.5m thick succession, and then with a second hiatus at the Cenomanian/Turonian boundary. The Late Cenomanian sequence boundary may correspond to the CeSin6 reported by Bauer et al. (2003) from Sinai, while their CeSin7 may correspond to the overlying latest Cenomanian sequence boundary. The Late Cenomanian sequence boundary may also correspond to the slightly younger CeJo4 sequence boundary reported by Schulze et al. (2003) from Jordan and to the CeUp sequence boundary at the top of the Tamar Formation of the Negev (Buchbinder et al. 2000), both within the Late Cenomanian N. vibrayeanus Zone (Fig. 3.3). The termination of the Late Cenomanian homoclinal carbonate ramp system of the Galilee (Ce HST-2) is unique, as drowning (by the Pelech sequence) was preceded by forced regression and subaerial exposure and followed by a second subaerial exposure event (latest Cenomanian sequence boundary). Regression and platform exposure at the end of the Cenomanian in Israel and Sinai were suggested by Flexer et al. (1986), Lewy & Avni (1988) and Bauer et al. (2003) on the basis of ferruginous crusts and borings, but robust indications for subaerial exposure such as karst, calcrete, palaeosols or fluvio-deltaic deposits were not reported. For example, the top of the latest Cenomanian in the Negev and central Israel is characterized by ferruginous burrowed or brecciated horizons that are more likely to be indicative of subaquatic omission (see recent example of Heck et al. 2007, ancient example of Immenhauser et al. 2000). Furthermore, in many marginal carbonate systems around the Tethys, subaerial exposure of Late Cenomanian platforms was not recognized and platform termination resulted from drowning due to sea-level rise, or eutrophication and anoxia (Jenkyns 1991; Philip & Airaud-Crumiere 1991; Gušic & Jelaska 1993; Philip et al. 1995; Drzewiecki & Simo 1997; Caus et al. 1997; Schlager 1999; Buchbinder et al. 2000; Scott et al. 2000; Bauer et al. 2001, 2003; Schulze et al. 2004).
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Chapter 3: Sequence stratigraphy, N. Israel
The first Late Cenomanian subaerial exposure event (Ce SB-3, Fig. 3.3) was influenced by eustatic fall at the M. geslinianum Zone (Flexer et al. 1986; Voigt et al. 2006), and also by local faulting in the Galilee (Fig. 3.8). On the other hand, the younger subaerial latestCenomanian sequence boundary (Ce SB-4) seems to result from local tectonic uplift of the Galilee and does not have equivalents in Europe. Late Cenomanian tectonic movements associated to local uplifts were also reported from northern Sinai by Bauer et al. (2003), and localized subaerial unconformities attributed to local tectonic events were reported from the Cenomanian on the eastern Arabian plate and the northern margins of the Tethys (Van Buchem et al. 1996 and Scott et al. 2000 – Oman; Borgomano 2000 – Apulian platform; Wilmsen 2000 – Cantabrian platform; Mouty et al. 2003 – northeast Palmyrides). 3.3.2.5. Sedimentary events across the latest Cenomanian sequence boundary and the manifestation of the OAE-2 The latest Cenomanian sequence boundary of the Galilee represents both a subaerial unconformity and transgressive surface onlapped by Early Turonian deepening-upward cycles UC-7, UC-8, UC-9, UC-11 and UC-12 (Figs. 2.8, 2.9). These Early Turonian cycles represent initiation of drowning caused by sea-level rise and onlap of deep-water facies onto the Galilee palaeo-high (Tu TST in Fig. 3.7). Above the latest Cenomanian sequence boundary, the Early Turonian Daliyya marls of the Carmel contain heterohelicid foraminifera, dinoflagellates, evidence for high vegetative productivity, high concentrations of total organic carbon and pyrite (Honigstein et al. 1989). The Daliyya marls were considered by Honigstein et al. (1989) to have formed under ecological stress and hypoxia, and were related to the second oceanic anoxic event of the Tethys (OAE-2; Arthur et al. 1987). The correlative marls in the CC6 cycles of the Yirka Formation, onlapping the Galilee from the south (Fig. 3.7), contain similar features indicating increased nutrients and hypoxia (facies type 1). The welllaminated and thin-bedded CC7 and UC9 cycles of the Bina Formation (facies type 11), onlapping the Galilee paleo-high from the north (Fig. 3.7) represent hypoxia and stagnation of Early Turonian sea-floor deposits as well. These well stratified, well laminated mudstones and wackestones are characterized by sparse but consistent occurrences of rotaliid foraminifera, especially gavelinellids (facies type 11 in Table 1, Fig. 3). Rotaliid foraminifera began to dominate oxygen-poor bottoms from the Albian, and small benthic foraminifera including gavelinellids were used as deep-water hypoxic markers of the CenomanianTuronian OAE in the Tarfaya basin, Morocco (Gerhardt et al. 2004). This well laminated rotaliid-bearing facies in the Bina Formation (facies type 11) therefore also reflects hypoxia related to the OAE-2. The succession of events above the Mid-Cenomanian sequence boundary (Fig. 3.10) shows that Early Turonian hypoxia related to the OAE-2 of the Tethys was at the culmination
58
Chapter 3: Sequence stratigraphy, N. Israel
of a lengthy trajectory of eutrophication. In the Ce TST-2, mass occurrences of pithonellid calcisphaeres, phosphatic grains and condensation indicate increased nutrients in the water body. During the subsequent Ce HST-2 the carbonate system failed to keep-up with rising sea-level as a result of low auto-production of skeletal carbonate (as reflected in the UC-3 cycles; see chapter 2), hence development of a homoclinal ramp (Fig. 2.10b). The Late Cenomanian sequence boundary reflects final deterioration of this non-productive homoclinal ramp at a time of Late Cenomanian emergence. The subsequent Pelech sequence, depleted in skeletal grainstones but rich with pithonellid calcispheres, reflects further eutrophicationinduced reduction of skeletal production leading to drowning of the system. This sequence was uplifted and subaerially-exposed within the latest Cenomanian but the overlying Early Turonian ammonite marls and laminites reflect culmination of eutrophication and hypoxia of the OAE-2. This chain of events (Fig. 3.10) tracks the gradual reduction in production of skeletal carbonate through the Late Cenomanian due to increased eutrophication.
3.4. Summary 1. Facies types of the Cenomanian-Turonian succession of northern Israel are integrated into three Cenomanian sequences and one Turonian sequence. Paleoenvironments were shown to be highly variable, ranging from deep basinal environments to peritidal and subaerial exposure. 2. The origin of major Cenomanian-Turonian depositional events that can be correlated across the Arabian plate and Europe is eustatic. The Late Cenomanian eustatic rise was to some extent decoupled and masked by local tectonism and uplift in northern Israel. Sedimentary overprints of bottom winnowing, oxygenation of hypoxic sediments, eutrophication or hypoxia reflect palaeoceanographic influence on some of the depositional events. Eustatic and/or palaeoenvironmental imprints were revealed for the first Cenomanian subaerial exposure; Early Cenomanian maximum flooding and oxygenation of hypoxic sea-floor; Mid-Cenomanian highstand progradation followed by forced-regression and mass-transport toward the basin; Mid-Cenomanian subaerial exposure; increased Late Cenomanian eutrophication throughout sea-level rise; Late Cenomanian subaerial exposure; latest Cenomanian/Turonian eutrophication and anoxia of the OAE-2 (“Bonarelli”) Event. 3. Two episodes of carbonate ramp demise were recorded: (a) An abrupt deterioration of productive prograding ramp in the Mid-Cenomanian by sea-level fall, gravity collapse, subaerial exposure and transgression, and (b) Deterioration of a non-productive aggradational homoclinal ramp by sea-level fall, faulting, subaerial exposure and drowning in the Late Cenomanian.
59
Chapter 3: Sequence stratigraphy, N. Israel
4. Throughout the Late Cenomanian, the carbonate system of the Galilee experienced a continuous decrease in skeletal grains production due to increasing eutrophication. These conditions culminated in the Early Turonian by extreme hypoxia related to the OAE-2. 5. Facies-thickness trends of Cenomanian-Turonian systems-tracts show that the southern Carmel and much of the Galilee were elevated during the Cenomanian and Turonian stages. The Galilee and southern Carmel paleo-highs were separated by a subsiding trough extending from the central-northern Carmel toward the southern Galilee. The Galilee paleo-high was bounded at the north by a deeper zone of rapid subsidence, extending across the northernmost Galilee and into Lebanon. This sedimentary configuration represents a shift of the north-south depositional strike of the midCretaceous Levantine hinge-belt toward the east-northeast. This trend corresponds to other large-scale tectono-depositional features in the Mesozoic Levant. 6. Late Cenomanian normal faults and a latest Cenomanian sequence boundary reflect tectonic activity and uplift of the Galilee to above sea-level near the end of the Cenomanian. Late Cenomanian faulting was associated with transformation of Late Cenomanian homoclinal ramp into a steep margin. Locally, calcarenitic clinoform unit formed steep slope beyond this margin, facing to the south-southwest. A concomitant, turbiditic toe-of-slope developed distally in the Carmel to the south.
60
61
62
63
64
65
66
67
68
69
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Chapter 4: Mid-Cenomanian ramp system
Chapter 4 Deterioration of Mid-Cenomanian ramp system in northern Israel: mechanical collapse, subaerial exposure, and submarine omission 4.1. Introduction The sequence stratigraphic position and tectono-sedimentary configuration of the Early to Middle Cenomanian phase of the carbonate system of northern Israel were summarized in Figure 3.5. The Early–Middle Cenomanian RST-1 (Ce RST-1) is progradational, incorporating both the Early to Middle Cenomanian highstand tract (Ce HST1) and the Mid-Cenomanian forced regression system tract (Ce FRST-1). The early highstand phase was shown to represent a distally steepened ramp, with high skeletal production on the mid-ramp (carbonate factory) contrasting markedly to the laminatedhypoxic, poorly productive outer ramp and adjoining basinal environment. The termination of this ramp system took place during the Middle Cenomanian A. rhotomagense ammonite Zone, in a sequence of events including mechanical ramp collapse, subaerial exposure, formation of the Ce SB-2, and drowning. The details of these events involved depositional processes, geomechanical mechanisms or diagenetic evidence that will be presented here in detail. An integrated mechanism will be suggested for the events surrounding the Ce SB-2, explaining ramp termination, events associated with the sequence boundary itself, and commencement of the overlying sequence by drowning. 4.2. Mass-transport features of the Mid-Cenomanian FRST-1: ramp termination by mechanical collapse prior to the Ce SB-2 Synsedimentary breccias, shear folds, fissures, normal faults and boudins in the Sakhnin Fm were first described by Ron (1978) in the western Galilee, but were not considered mass transport slope phenomena. Instead they were considered to represent a highly localized event of vertical gravity collapse of interbedded lithified/non-lithified beds on a wide, shallow water shelf-lagoon platform. However, the spatial distribution of gravity collapse features across the Galilee in their sequence stratigraphic context, shown in Chapter 3 for the Ce FRST-1 (Fig. 3.5), suggests that these are in fact mass transport slope deposits. Synsedimentary deformation features associated with the Ce FRST-1 are found in the Sakhnin Fm of the Galilee and in the Zikhron Fm of the southern Carmel region. Some examples of Ce FRST-1 mass transport deposits of the Galilee are here described.
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Chapter 4: Mid-Cenomanian ramp system
4.2.1. The Namer Valley outcrops, NW Galilee: shear zones, translational slides and debrites The Sakhnin Fm of the Namer Valley appears to be stratified, but on closer examination, fragmented and brecciated blocks embedded in dolomitic matrix form largely concordant bodies within it. Characteristic sub-horizontal or slightly curved narrow zones composed of distinctive yellowish-foliated lithologies broadly parallel bedding planes, often separating these breccia bodies (Pl. 7a-e). These foliated lithologies may contain elongated, rotated or symmetrically- or asymmetrically-winged mudstone objects or boudins embedded in fine-grained matrix (Pl. 7 f). The shape of these objects indicates strong elongation and rotation due to subhorizontal shear, with some resemblance to mylonitic textures. Some of the elongated, rotated or symmetric/asymmetric winged mudstone objects resemble typical mylonitic σ and δ textures. In a few examples, the orientation of the foliations corresponds to typical mylonitic C/S foliations, a type of shear-band cleavage. C-type shear bands are parallel to shear zone boundaries and S-type shear bands are oblique (Passchier &Trouw 1998). Miller (1996) showed similar foliations along normal faults in the Turtleback fault system in California, and considered them ductile. Cladouhos (1999) termed similar shallow crust foliations “Flow Banded Gouge” (FBG). Presence of FBG features between Ce HST-1 beds formed under peritidal conditions in the Sakhnin Fm along the Namer Valley region reflects, therefore, post-depositional, subhorizontal movement along subhorizontal shear planes and other slightly curved shear planes. Breccia blocks of Sakhnin Fm with clasts of laminated dolomites of algal mat origin derived from inner-ramp Ce HST-1 sediments are also found emplaced in deformed matrix (Pl. 7g), bounded by FBG shear planes. In places, the matrix is of the autochthonous basinal laminites (FT-1, Table 2.1). In one Namer Valley outcrop, a fragmented bed with a recumbent fold (axial plane striking 200°-20°) (Pl. 7h) overlies a slightly curved shear plane (Pl. 8a). These features indicate block sliding along a curved plane toward the W-NW. Bed fragmentation above the shear plane (Pl. 8b) appears to be a consequence of sub-horizontal movement of semilithified blocks along wavy curved shear planes with minor block rotations. Ellis & McClay (1988) showed that one factor dictating the slip along listric extensional faults is the degree of curvature of the detachment. Movement along the low-angle curvature of some of the shear planes would then result in only minor brecciation, but a wavy morphology would cause more damage. Block movements along sub-horizontal shear planes or shear planes with minor rotations closely correspond to translational slides sensu Cook & Mullins (1983).
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Chapter 4: Mid-Cenomanian ramp system
Translational slides of the Namer Valley region pass to the west into non-bedded breccia encased in laminar matrix (e.g. Pl. 7g). Similar breccias were found in the Ce FRST1 of the Betzet Valley, Kziv Valley section and Sheikh Danun section (e.g. Pl. 8c,d; Fig. 3.2), interbedded with dolomitized matrix. Their non-channelized morphology and their location to the west, suggest that they represent debris sheets beyond the edge of the Mid-Cenomanian distally-steepened ramp. 4.2.2. Adamit region, NW Galilee: A channel levee complex A notable mass transport feature occurs also in the Sakhnin Fm of the Betzet Valley east of Adamit (Eilon section, Fig. 4.1), above the bioclastic horizon of Early Cenomanian MFI-1. Well defined lenticular dolomite bodies characteristic of the Sakhnin Fm are incised within basinal facies at the top of the Deir-Hanna Fm (Pl. 9a). These bodies are composed of well stratified or laminated, thin bedded and well graded dolomitized calcarenites, composed of molluscan-foraminiferal debris. The graded calcarenites clearly show Bouma (1962) Ta division (S3 division of Lowe 1982) overlain by parallel laminae of Bouma Tb division (Pl. 9b) and thus were transported as turbidites. These bodies are overlain by laterally extensive finer-grained dolomitized deposits composed of carbonate silts and muds which may be also laminar and graded, rippled or bioturbated, and show the Bouma Ta, Tb, Tc and occasionally Td division (Pl. 9c). The lenticular dolomite bodies are interpreted as the calcarenitic fill of basinwarddraining slope channels. The finer-grained accumulations above and beyond the channel reflecting low density calciturbiditic overbank spills, covering a wide area and forming broad overbank wedges (e.g. Mutti & Normak 1991). 4.2.3. Yanuch region, western Galilee: slide scar and turbidites Near Yanuch in the western Galilee, the CC1 cycle (Fig. 2.10a; Sakhnin Fm) normally found above the laminites of the Deir Hanna Fm, and beneath bioturbated limestones of the Yanuch or Bina Fms, is completely absent. Freund (1958, 1965) suggested that the Sakhnin dolomites are of shallow-water platform facies, and their absence at this level in the western Galilee is a result of locally increased subsidence and formation of a narrow intrashelf basin (the ‘Yirka basin’). An oblique-planar surface truncates the Deir Hanna laminites at this stratigraphic position (Pl. 1c). Microfacies on and near the surface show elongated surface-parallel lithoclasts, of which some are geometrically similar to winged- or stair-stepping objects common in mylonitic shear zones, and some resemble ’σ-objects’ (Passchier & Trouw 1998) (Pl. 9d,e). This oblique truncation of the Deir Hanna laminites appears to be a shear plane.
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Chapter 4: Mid-Cenomanian ramp system
Reappearance of the Sakhnin Fm laterally in the form of dolomitized turbidites, suggests that the deformed horizon at Yanuch represents a slide scar. Truncation of semi-consolidated basinal laminites of the Deir Hanna Fm suggests that it occurs beneath the original position of material that was mass-transported. The Yanuch Fm above the shear plane is composed of subtidal cycles (UC3) and is free of internal deformation and is thus autochthonous. Stratigraphic relations between the Yanuch section, where this shear plane is located at the stratigraphic position of the absent Sakhnin Fm, and the nearby Gat-Yanuch section, where the Sakhnin Fm is present at that level as a dolomitized turbiditic succession (Pl. 9f), are shown in Figure 4.2. Above the Gat-Yanuch turbidites, stratified fine grained, pelagic calcisphere-rich wackestones of the Yanuch Fm are present, and above them, the ammonite-bearing marls of the Yirka Fm. The slide plane at Yanuch section and the turbiditic succession at Gat-Yanuch section are at the same stratigraphic position and are therefore genetically connected. They represent erosional and depositional phases of slope mass transport of the Ce FRST-1. 4.2.4. The Betzet mega-block, NW Galilee The Ce RST-1 succession mapped as Sakhnin Fm at Betzet, is anomalously emplaced within FT-1 laminated basinal facies. Continuity of Ce MFI-1 between the Betzet section and Eilon-Adamit section 2500m to the east (Fig. 4.1) shows that a complete but internally deformed CC1 cycle (shear-zones, debrite, breccias) at Betzet terminates sharply against a basinal succession topped by a deep-water channel complex (Pl. 9a-c). West of the CC1 cycle at Betzet, breccias derived from that cycle are embedded in basinal laminites. The CC1 cycle embedded in the Betzet section can best be explained as a displaced, internally deformed mega-block transported along a detachment surface similar to the slide scar exposed at Yanuch. 4.2.5. Model for mass-transport across the distally steepened slope East to west facies transitions across the Galilee, from Dishon to Sheikh Danun (Fig. 4.3), show how the proximal part of the ramp at Peqi'in and Deir El Assad collapsed and passed into distal mass-transported deposits toward the west. The Mid-Cenomanian RST-1 at Dishon is composed of autochthonous peritidal UC-1 cycles. Similar cycles at the same stratigraphic position originally occurred to the west at Peqi’n and Deir El-Assad. However, the Ce RST-1 peritidal cycles of Peqi’in and Deir El-Assad are represented by breccias, shear folds, boudins and shear zones. At Yanuch to the west, brecciated and deformed peritidal cycles of the Sakhnin Fm are absent, but a slide scar (Yanuch slide scar) occurs at this stratigraphic position (see above and Fig. 4.2). Further toward the west, relatively thin (few meters in thickness) turbidites and debrites appear in the Gat/Yanuch, Kziv-west and
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Sheikh Danun sections, representing mass-transport of inner-ramp deposits into the basin at the west. The spatial distribution of these slope features have already shown and discussed in Chapter 3 (Fig. 3.5). The synsedimentary deformation features described above occur as part of the Mid-Ce FRST-1. They reflect mechanical collapse and downslope gravity mass transport of both lithified beds and unconsolidated grains originating on the inner ramp. These allochthonous sediments were then deposited on various positions on the slope of the distally-steepened ramp system. They represent the beginning of the termination of the first Cenomanian sequence of northern Israel prior to the Ce SB-2.
4.3. Diagenetic features of the Mid-Cenomanian Sequence Boundary (Ce SB-2) and transgressive surface Mid-ramp conditions prevailed during the Ce HST-1 in the relatively distal section at Manara, and the peritidal succession characterizing the upper part of this sequence did not develop. Maximum progradation is represented by a bioturbated mid-ramp facies colonized by rudists, gastropods and other mollusks. Special diagenetic features characterize the uppermost few centimeters of this cycle at Manara. Most marked is a black-reddish ferro(manganese?) pavement. A series of diagenetic features were identified at this level (Fig. 4.4): (1) microsparitic mosaics formed by aggrading recrystallization of micrite; (2) dissolution vugs; (3) pore-filling bladed rim calcites and equant calcitic mosaics within dissolution vugs; (4) fibrous chalcedonic overlays, spherulitic chalcedony and mega-quartz crystals; (5) pore-filling flowstone speleothems, and concentric pendant calcites showing preferred growth and uniform extinction; (6) silicified or partly silicified spheroidal pisoids, some slightly deformed; (7) angular micritic lithoclasts; and (8) well rounded mold peloids, some with planktonic foraminifera. The host rock upon which these diagenetic elements were superposed, was the bioturbated mid-ramp facies of the Mid-Ce HST-1. The host rock is rich in calcitic and aragonitic skeletons, especially of radiolitid rudists and gastropods. The original matrix was micritic, peloidal/pelletal or fine-grained calcisiltitic, with contributions of pelagic grains. Primary pores were absent and no early-marine phreatic cements such as micritic rims or isopachous fibrous cements were formed. The earliest diagenetic imprints are of freshwater origin (see below). Four
diagenetic
phases,
interrupted
by
dissolution
discontinuities
can
be
distinguished as follows (Fig. 4.4): Diagenetic phase 1 marks a shift from marine to freshwater-phreatic conditions, with leaching of aragonitic shells, especially gastropods and parts of radiolitid rudists. This phase
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Chapter 4: Mid-Cenomanian ramp system
resulted from influx of undersaturated freshwaters from the partly-exposed tidal flats of the Ce HST-1, forming an active freshwater phreatic zone. Leached pores were partially or completely filled by freshwater cements, following a typical sequence beginning with bladed calcitic rims, followed by clear equant, occasionally drusy calcites (Fig. 4.4a, b). Recrystallization of micrite due to freshwater contact is attributed to this stage. Finally, cements were partly silicified. Diagenetic phase 2 took place in the freshwater vadose environment. Dissolution truncated the pore-filling equant cements of phase 1, and new dissolution vugs were formed. Precipitation in phase 2 is by fibrous chalcedonic overlays, a typical pore-filling cement (Hesse, 1989), and by equant mega-quartz mosaics and spherulitic chalcedony (Fig. 4.4c,d,e) that partly or completely filled dissolution vugs. Diagenetic phase 3 commenced with dissolution of earlier siliceous cement, forming dissolution cavities, subsequently filled by vadose-type cements including pendant calcites, vadose and cave pisoids and flowstones, some partly silicified (Fig. 4.4e,f,g). They were precipitated above the water table when percolating freshwaters became oversaturated with respect to calcite. Diagenetic phase 4 commenced with in the vadose zone with partial dissolution of calcitic pendant cements and vadose pisoids, and formation of new dissolution vugs that remained empty (Fig. 4.4e). The empty vugs mark termination of calcite cement precipitation. These features are overprinted by development of a ferro-manganese encrustation, paving the upper surface (Fig. 4.4h), inhibiting water penetration from above by sealing open cavities, and enabling preservation of remaining vadose cements and empty dissolution vugs. Mineralization at this final stage is attributed to drowning, and condensation (see discussion).
4.4. Discussion 4.4.1. Mechanism of slope collapse The Ce SB-2 at the top of the Ce RST-1 was shown in Chapter 3 to be correlative to a Mid-Cenomanian eustatic fall of sea level. Liquefaction and water escape during a continuous fall of relative sea-level can lead to mechanical collapse (e.g. Spence & Tucker 1997). On the other hand, mechanical collapse could also be enhanced by factors unique to the paleoecological context of the depositional regime at Ce RST-1. The CC1 cycle (Fig. 2.10a) is the vertical cyclic expression of the Mid-Cenomanian distally-steepened ramp in most of the Galilee and the southern Carmel. It is characterized by a sharp boundary transition between basinal hypoxic laminites at the base, to overlying rudist-gastropod mid-ramp carbonate factory facies. This upward trend reflects the transition
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from an expanded oxygen minimum zone (OMZ) impinging on the mid-ramp, to an oxic mixed layer (Fig. 2.10a). This vertical trend also suggests that progradation of the ramp, controlled by the mid-ramp carbonate factory in the mixing zone, was constrained distally by OMZ impingement, causing distal steepening at that point. Basinward-progradation took place during the entire Ce HST-1, indicating that the upper surface of the OMZ was continuously displaced seaward, possibly in harmony with falling sea-level. A significant deceleration of falling of the top OMZ, or long-term stabilization of the OMZ, could contribute to cessation of progradation resulting in accumulation of mid-ramp carbonates by aggradation, forming an unstable sediment pile. Overloading on the mid-ramp caused the ramp-slope system to become mechanically unstable, and subsequently, to collapse. This series of events (Fig. 4.5) took place near the end of the Ce RST-1, when sea level stabilization and deceleration of OMZ fall took place. The ramp collapsed, and at the transition to the subsequent transgressive stage (Ce TST-2; Fig. 3.3), laminated OMZ facies again began to onlap the ramp. Coincidence of two factors was therefore responsible for the mechanical collapse of the Mid-Cenomanian ramp system of northern Israel: (a) pressure release and pore-water escape in the course of Mid-Cenomanian eustatic sea level fall and (b) sediment aggradation and overloading on the mid- ramp due to deceleration in the rate of the falling OMZ. 4.4.2. Mid-Cenomanian ramp termination: a composite discontinuity The distally steepened carbonate ramp system was terminated in much of the Galilee by mechanical processes.
Continuing sea level fall then led to a period of subaerial
exposure and formation of the Ce SB-2, and as the ramp was again flooded, drowning was associated with submarine omission phenomena and formation of a transgressive surface. The four-part diagenetic sequence at the termination of the Ce RST-1 at Manara demonstrates gradual change from marine mid-ramp depositional setting to freshwater phreatic and vadose diagenetic settings, indicating exposure of the Ce HST-1 above sealevel. Further evidence for subaerial exposure from more proximal parts of the ramp are teepee structures reported by Bogoch et al. (1994) from the Sakhnin - Bina Fm transition in the eastern Galilee (Kadarim region, Fig. 1.1). However, limited penetration of the minor karst and pedogenetic features, and absence of extensive paleosols or fluvio-deltaic deposits, imply that emergence above the Ce SB-2 was probably relatively short. The ferromanganese crust that terminates diagenetic phase 4 (Fig. 4.4h) is characteristic of ramp drowning as a response to rapid rise of relative sea-level, and indicates sediment starvation and condensation. In the view of Longman (1980), preservation of vadose cements requires rapid transgression or subsidence. Therefore, the Fe-Mn crusts
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Chapter 4: Mid-Cenomanian ramp system
on the Ce SB-2 cannot be considered lateritic features (cf. Lewy & Avni 1988), unless good pedogenetic evidence is present. Composite discontinuity surfaces with subaerial features fused to ferruginous crusts of marine origin were also reported by Immenhauser et al. (2000) from the Middle Cretaceous of northern Oman ("polygenic discontinuity surfaces") and by Calner (2002) from the Middle Silurian of Gotland, Sweden. Encrustation of submarine surfaces, even in the deep ocean, by ferromanganese mineralization is often reported from ancient neritic and bathyal submarine hardgrounds (e.g., surfaces described by Scholle & Kennedy 1974; Lewy 2002) and from the subrecent record such as on the surface of the Marion Plateau, offshore northeast Australia (Heck et al. 2007). These crusts develop slowly on the sea floors (1-3mm/myr) and therefore represent condensation (Yang et al. 2006). Presence of pyroclastics on the Ce SB-2 in the southern Carmel region (Segev et al. 2002) suggests that there may have been some tectonic control on the position of the Ce SB-2 in the Carmel region, but any such effect would have been highly localized. 4.5. Summary The Mid Cenomanian ramp system was first terminated by mechanical collapse, then exposed subaerially, and finally became a submarine omission surface due to flooding: 1. The forced regressive system tract prior to the Ce SB-2 includes parts of the Sakhnin Fm in the Galilee and the Zikhron Fm of the southern Carmel. This system tract is dominated by a mass-transport complex, reflecting mechanical collapse of the MidCenomanian distally-steepened ramp of northern Israel. 2. Mass transport features of the Mid-Cenomanian forced regressive system tract include a variety of allochthonous slope deposits such as shear-zone lithologies (foliations, boudins, shear-folds, ‘mylonitic’ micro-features), translational slides, slidescars, debrites and channel-levee turbidites. 3. The mechanical collapse of the Mid-Cenomanian ramp system of northern Israel is attributed to coincidence of two factors: (a) pressure release, fluid escape and liquefaction during the course of the Mid-Cenomanian eustatic fall; and (b) stabilization of OMZ impingement on the mid-ramp, transition from a progradational to an aggradational carbonate factory, and consequently, overloading of the ramp at the locus of distal steepening. The mechanical collapse of the steepened edge was therefore controlled in two ways by the Mid-Cenomanian eustatic fall. 4. The Mid-Cenomanian ramp system is capped by a composite discontinuity surface composed of four superposed diagenetic phases. The earlier phases indicate gradual transition from marine to fresh-water vadose diagenesis and sub-aerial exposure, but the final phase formed during the first stage of renewed ramp flooding, and represents submarine omission and condensation.
78
79
80
81
82
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Chapter 5: Late Cenomanian shelf-to-basin system
Chapter 5 Growth-fault-controlled calcarenitic shelf-margin apron across the Cenomanian/Turonian boundary, western Galilee, northern Israel 5.1. Introduction The final 1.25m.y. of the Late Cenomanian, until the beginning of the Turonian, are represented in northern Israel by four genetic carbonate units, viz: a progradational Late Cenomanian highstand system tract (Ce HST-1); a Late Cenomanian forced regression system tract (Ce FRST-2); a short late Cenomanian sequence (Pelech Sequence); and an Early Turonian transgressive system tract (Tu TST). Sedimentary configuration of these system tracts and main events across the Cenomanian-Turonian boundary, were presented in Chapter 3 (Fig. 3.7). Late Cenomanian and Early Turonian successions in the western Galilee were incorporated into the Yanuch and Yirka Fms (Freund 1959), and in the Carmel region into the Muhraqa and Daliyya Fms (Picard & Kashai 1958). Stacking patterns of genetic units (See Chapter 2,) show that the Yanuch Fm consists mostly of meter-scale shallowing-upward cycles of UC3 type (Figs. 2.2, 2.10b), reflecting inefficient shoreface progradation on a homoclinal ramp. The Yirka Fm consists of deepening-upward cycles of UC7 type (Figs. 2.2, 2.9). The upper part of the Muhraqa Fm, and the Daliyya Fm in the Carmel, are similarly formed of deepening upward cycles, of UC12 and UC8 type (Rakefet and Isfiyye sections, Fig. 3.9). The focus of the present chapter is on the detailed structural and sedimentary evolution of these Late Cenomanian system tracts in the western Galilee and Carmel regions. In the western Galilee, these tracts are represented in outcrops sometimes by laterally discontinuous, sometimes deformed, sediment bodies indicating a complex sedimentary history with synsedimentary structural features indicating compression, extension and shear. Unravelling the complexity of these features reveals details of sedimentary-structural evolution of the shelf-margin in this key interval, especially during genesis of the Late Cenomanian forced-regressive system tract (Ce FRST-2). The findings show that Ce FRST-2 developed in the western Galilee on a steep, southward-facing, faultcontrolled calcarenitic slope, with calciturbidites extending as far as the Carmel region to the S-SW. Remarkably, this calcarenitic ‘shelf-margin apron’ system has features of sedimentary geometry and structural control in common with siliciclastic shelf-margin deltas.
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Chapter 5: Late Cenomanian shelf-to-basin system
5.2. Late Cenomanian clinoform body in the western Galilee 5.2.1. Distribution and characteristics of the clinoform body The upper part of the Yanuch Fm in the western Galilee, south of the Yirka Fault (Fig. 1.1), is composed of calcarenitic clinoform body, well exposed at Kishor-Blaya, Hamra Valley and the Mt. Gamal (Fig. 5.1). Its contact with the underlying, cyclic part of the Yanuch Fm (Fig. 2.10b) is an erosional surface as observed in the clino-succession of Kishor (Pl. 10a,b). The clino-succession at Kishor (Pl. 10a) is 35-40m in thickness, with initial sedimentary inclinations of 22°-25°. At Hamra Valley some 3km to the south it is more than 100m thick, with steeper inclinations, 30° to 35°. At Mt. Gamal, two kilometers farther south, the succession is similar but highly truncated. Further to the south, small additional outcrops are found in the Shagor Canyon (Fig. 5.1). The clinoforms are well-stratified or laminated, in beds 5-30cm in thickness, and composed of grainstones or rudstones with bioclasts of rudists, oysters, echinoderms, dasyclad algae, mud peloids and microbial lumps (e.g. Pl. 10c-f). Grain fabric is usually mudfree; the grains are well-sorted, mostly angular and coarse-grained, but only rarely graded. At Kishor, well-sorted and cross-bedded grainstones form the upper 17m of the clinoform body (Pl. 10g). At Hamra, convex-upward truncation surfaces a few meters wide are filled with thin bedded or massive grainstones (Pl. 10h). Petrographic examination shows that neomorphism often affects skeletal grains and matrix (e.g. (Pl. 10d,e), but where less intense, original cements and early diagenetic stages can be recognized (e.g. Pl. 10c). Early marine rim cements around carbonate grains, and blocky calcite in remnant pores, are common (Pl. 10f), as well as fabric-selective interparticle pores (Pl. 10e). Late features are penetrating neomorphism, and microstalactitic cements (Pl. 10e). The upper boundary of the clinoform body is a discontinuity surface, an amalgamation of the Late Cenomanian sequence boundary (Ce SB-3), Latest Cenomanian sequence boundary (Ce SB-4) and Early Turonian transgressive surface (Fig. 3.8). The diagenetic overprints corroborate field evidence for exposure at this disconformity. On top of the clinoforms at Mt. Gamal there is an irregular karst surface covered by a thin calcrete deposit (Pls. 5b, 6c). The deepening-upward Yirka Fm overlies the discontinuity. 5.2.2. Genesis of clinoform features These clinoforms were mapped by Freund (1965) as member CTb of the Yanuch Fm. They were considered as fore-reef talus of ‘rudist reefs’, growing along the margin of a local, NNE-trending intrashelf basin (the ‘Yirka basin’). A different model for genesis of these features is here presented.
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Chapter 5: Late Cenomanian shelf-to-basin system
Textural characteristics indicate that some of the calcarenites forming the clinoforms were transported primarily by submarine grain-flow (Lowe 1976). Mud-free non-graded textures occur when dispersive pressure is the primary grain-supporting force. Thin beds and laminae found in grain-flow deposits occur in the clinoforms, where dispersive pressure alone cannot support grain accumulations thicker than 5cm. Under sub-aqueous conditions they can be somewhat thicker (Leeder, 1999), but generally, thicker clino-beds are better interpreted as sandy debris-flows (sensu Shanmugam 1997). Steep depositional slopes (20°35°) of the Galilean clinoforms suggest repeated over-steepening to angle of repose, prone to failure and grain flow. Sediment-filled convex-upward truncation surfaces found at Hamra, represent channels incised into such an over-steepened calcarenite slope, exceeding the angle of repose for coarse calcarenite. These channels were then filled by sandy debrites. Diagenetic features and the calcrete surface at the top indicate that the clinoforms were initially deposited in a marine environment with active flushing of pore-waters, buried for a time in the marine-phreatic zone, and finally, sub-aerially exposed.
5.3. Special features of the clinoform body of the western Galilee Special features of the clinoform body of the western Galilee will be described from north to south, beginning at Yanuch and Beit-Ha'Emek sections, to Kishor, Hamra Valley, and Mt. Gamal outcrops (Fig. 5.1). Full data from these sections are given in Appendix 2. 5.3.1. Significance of the Yirka Fault The Yirka Fault (Figs. 1.1, 5.1) has a marked morphological expression, transecting the western Galilee from east to west, and belongs to a regional system of faults considerably younger than the Cretaceous (Freund 1970). North of the Yirka Fault, at Yanuch and Beit-Ha’Emek, typical UC3 cycles of the Late Cenomanian highstand form most of the Yanuch Fm, and no clinoforms are present beneath the Yirka Fm. South of the fault, thick clinoform bodies overly the UC3 stack at Kishor, Hamra and Mt. Gamal. The Kishor clino-succession is actually in contact with the fault, with SW-facing syndepositional inclinations dipping perpendicular to the strike of the Yirka Fault. This young fault therefore appears to be coincident with a structural feature that strongly influenced the Late Cenomanian clino-succession. 5.3.2. Special features of the Kishor-Blaya and Hamra clino-successions 5.3.2.1. Description of special featutes at Kishor-Blaya and Hamra The UC3 cycles of the lower part of the Yanuch Fm at Kishor are inclined 12° to the NE. The overlying clinoforms, in contact with the Yirka Fault, are 35-40m in thickness and dip
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Chapter 5: Late Cenomanian shelf-to-basin system
22°-25° to the SW (Pl. 10a). The upper 17m of the clinoforms are well-sorted grainstones, with ripple and trough cross-lamination. Along their strike toward the west the Kishor clinoforms interfinger with deep-basinal pelagic facies of the Yanuch Fm (Pl. 11a). Approximately 1000m southwest of the Kishor-Blaya outcrops, the Hamra clinoforms are more than 100m in thickness, with steeper inclinations, up to 35°, dipping SW. 5.3.2.2. Late Cenomanian events at Kishor-Blaya and Hamra The clinoform prism, ranging in thickness from 0 north of the Yirka Fault, to 35-40m at Kishor and to more than 100m at Hamra (Fig. 5.2) must reflect subsidence to the south. This subsidence took place over the hangingwall block of an E-W - trending Late Cenomanian growth-fault located at the exact position of the present Yirka Fault (Fig. 5.1). Block movement coincided with progradation of shoreface deposits north of the Yirka Fault (UC3 cycles), resulting in transport of bioclastic grains across the Yirka Fault into the accommodation space created over the hangingwall block. Opposite inclinations between the NE-facing UC3 cycles underlying the clinoforms and the SW-facing clinoforms indicate that the hangingwall block rotated to the NE (Fig. 5.3). Interfingering of the Kishor clinoforms along their strike to the west with basinal deposits of the Yanuch Fm (Pl. 11a) indicate that hangingwall subsidence of at least 35-40m was continuous from shoreface to basinal depths. At Hamra, the much thicker clino-succession (>100m) points to even greater subsidence toward the south, requiring additional Cenomanian fault (Fig. 5.3). 5.3.3. Special features of the Hamra Valley outcrops 5.3.3.1. Description of special features at Hamra Valley In the Hamra Valley cross section (Fig. 5.4), a Late Cenomanian syncline (Pelech syncline) a few hundreds meters wide is developed in the upper part of the Yanuch Fm in the easternmost outcrops (Pl. 11c). Well-stratified limestones of open-marine outer- to mid-ramp facies are folded into this syncline, but upwards, younger beds become gradually less tilted. The synclinal succession is bounded by discontinuities forming a complete sequence (Pelech sequence, Chapter 3). It is bounded at the base by the Late Cenomanian sequence boundary (Ce SB-3), and at the top by the latest Cenomanian sequence boundary (Ce SB4). The Pelech syncline is in sharp vertical contact against the Hamra clino-succession to the west, and the lithofacies forming the syncline also overlaps the clinoforms (Fig. 5.4; Pl. 12a). Further to the west, the Hamra clino-succession together with the overlapping synclinal deposits are faulted (Fig. 5.4; Pl. 11b), with the down-faulted block inclined antithetically to the fault plane. The entire section is truncated by Ce SB-4, and Early Turonian transgressive deposits of the Yirka Fm are present above. The Yirka marls were neither folded nor faulted, but become much thicker toward the west.
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Chapter 5: Late Cenomanian shelf-to-basin system
The Yirka Fm of the Hamra outcrops belong to three subunits (Fig. 5.4): a basal unit 16m thick, of marl-limestone couplets (Pl. 6f) of basinal-pelagic origin; 4m of graded turbiditic grainstones intercalated with pelagic marls (Pl. 6g); and at the top, pelagic marls bearing ammonites (Fig. 6h) indicating Early Turonian zone T5 and higher. 5.3.3.2. Late Cenomanian events at the Hamra Valley An interpretative model for the exposed features of the Hamra Valley is in Figure 5.5. The sharp vertical discontinuity between the Pelech syncline and the Hamra clinoforms is a syndepositional fault scar. Decreasing inclinations toward younger beds in the Pelech syncline marks it as a Late Cenomanian growth structure. Deceleration of syndepositional folding was followed by spillover of synclinal deposits onto the Hamra clinoforms in the west. Features exposed at Hamra can be explained conservatively by proposing a single composite detachment in the subsurface, with a steep ramp zone underlying the Pelech growth syncline (Fig. 5.5). The Pelech syncline can be explained as a hangingwall syncline overlying a curved ramp zone with complex listric detachments (c.f. Ellis & McClay 1988). Movement on the bounding fault of the Pelech syncline formed new accommodation space above the hangingwall that was rapidly filled by coarse-grained bioclasts, forming steep clinoforms. Late Cenomanian eustatic fall terminated both deposition and overload-related synsedimentary subsidence, forming the erosional Ce SB-3. Calcarenitic deposits were eroded from above the footwall, but a considerable thickness of the Hamra clino-succession was preserved over the hangingwall. The subsequent sea level rise led to infilling of the Pelech growth syncline by the stratified deposits of the Pelech sequence. As rate of folding fell behind sedimentation rate, sediments spilled over the Hamra clinoforms to the west. Renewed faulting took place just prior to the latest Cenomanian relative sea-level fall. Subaerial erosion then formed a latest Cenomanian sequence boundary (Ce SB-4). This discontinuity was apparently due to local uplift of the Galilee (see Chapter 3) in the face of ongoing Late Cenomanian eustatic rise. Ultimately, highly transgressive basinal facies of the Yirka Fm overstepped the clinoformic slope. 5.3.4. Special features of the Mt. Gamal outcrops 5.3.4.1. Description of special features at Mt. Gamal region Discontinuous blocks in this region are shown in the cross-section along the Yizhar Valley (Fig. 5.6). The easternmost part of the transverse is in the stratified UC3 cycles of the lower Yanuch Fm, dipping 10°-12° to the W-SW. They are overlain to the west by 35m of calcarenitic clino-beds dipping 10°-15° to the E-NE. The UC3 cycles are separated from the overlying clinoforms by discontinuities marked Ds1 and Ds2 (Fig. 5.6; Pl. 12b). The clinoforms are truncated by an internal angular and erosional discontinuity marked Ds5, and
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Chapter 5: Late Cenomanian shelf-to-basin system
topped by an additional set of clino-beds inclined up to 35°, to the W-SW (Pl. 12b,c). Continuing westwards, the clino-succession again terminates sharply by discontinuity marked Ds2, against well-stratified UC3 cycles, dipping up to 30° to the NE. These NE-tilted beds terminate against a block of gently-inclined SW-facing beds of the same UC3 cycles, upon a discontinuity marked Ds3. Farther to the west, another discontinuity labeled Ds4 separated the NE-tilted block from another UC3 block in the west. Truncating all the discontinuous blocks along this traverse is an irregular erosion surface labeled Ds6, paved by ferruginous concentrations and remnant calcrete (MY section, Fig. 5.6). The discontinuity Ds6 is overlain by the Early Turonian pelagic Yirka Fm. 5.3.4.2. Late Cenomanian events at the Mt. Gamal region The morphology of discontinuity Ds3 corresponds to a high angle reverse fault (Pl. 12d). Discontinuity Ds4 corresponds to a normal fault formed farther to the west. A model for reconstructing the Cenomanian synsedimentary development of the Mt. Gamal region in five stages is shown in Figure 5.7. In stage 1, stratified block composed of UC3 ramp cycles was sliding along rotational detachment and was tilted to the NE forming a reverse fault. The sliding block was over-riding another block in the west (Ds3) and a normal fault was formed (Ds4). Rotational sliding left a depression above the hangingwall that was filled with bioclastic clinoformic sands forming Ds1 and Ds2. In stage 2, further movement along the listric detachment caused re-rotation of the clinoformic beds. The additional accommodation space was filled by SW-facing bioclastic clinoforms of the upper clinosuccession forming Ds5 at their base. In stage 3, a further block rotation tilted the clinoforms top by 12° to the NE. In stage 4, the entire system was sub-aerially exposed, Ds6 was formed (Ce SB-4), and the highly porous clino-succession was cemented. Resistance conferred by this cementation then led to differential erosion and accentuation of the clinosuccession as a topographic high on the erosion surface. In stage 5, the Turonian transgression led to overstepping of pelagic sediments of the Yirka Fm. 5.3.5. Late Cenomanian turbidites, debrites and shear-zones in the western Carmel region The middle part of the Muhraqa Fm in the Carmel range overlies Mid-Cenomanian chalks of the Arkan Fm (Lipson-Benitah et al. 1997; Segev et al. 2002) (Fig. 2.2) and unconformably underlies Turonian limestones with hippuritid rudists. It therefore reflects Late Cenomanian events, coeval with Late Cenomanian events in the western Galilee. The lower part of the Muhraqa Fm in the western Carmel is composed of calcisphere-rich wackestone (FT-6 in Table 2.1, Fig. 2.7) and the younger part is composed mostly of stratified grainstones and wackestones, and breccias with meter-sized blocks. The uppermost,
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Turonian part of the Muhraqa Fm, with hippuritid rudists (Buchbinder et al. 2000), was not recognized in this region. Relevant outcrops from the middle part of the Muhraqa Fm in the western Carmel were studied at the Oren and Megadim valleys, Sefunim quarry, and Isfiyye (Fig. 1.1). 5.3.5.1. Megadim- and Oren valley outcrops The fine-grained calcispheric packstones of the lower part of the Muhraqa Fm (FT-6), are overlain by 25-30m of well-stratified grainstones (Pl. 13a). The grainstones are peloidal/bioclastic, mostly graded, usually well sorted, and occasionally rippled or scoured (Pl. 13b,c). 5.3.5.2. Late Cenomanian events at Megadim and Oren valleys The thick well-stratified grainstone succession of the Megadim and Oren valleys correspond to highly amalgamated, mostly high density proximal turbidites, stacked mostly as Ta/Tb Bouma divisions, occasionally with Tc/Td low-density units (e.g., Pl. 13c). Rare decimeter-scale ripple cross laminations correspond to Tc divisions. Rarity of intervening fine grained pelagic limestones suggests high-frequency turbidite deposition. Similar but dolomitized turbidites are found in the mid-Muhraqa Fm of the Rakefet Valley in the southeastern Carmel (Fig. 1.1). 5.3.5.3. The Sefunim quarry section The sedimentary features of the Sefunim outcrops are schematically summarized in Figure 5.8. The base of the exposed part at the Sefunim quarry consists of coarse breccia with blocks containing unbroken radiolitid rudists, gastropods, bryozoans and green calcareous algae. The blocks are decimeters to a few meters in length, randomly oriented or imbricated (Pl. 13d). Matrix consists of coarse bioclastic material, occasionally laminated or graded, penetrated by flame structures of fine grained sediment (Pl. 13e). This part of the section is overlain by a massive bed composed of bioclastic debris, truncated by a sharp erosive contact (Fig. 5.8). The contact is offset by a normal listric fault and is overlain by wellstratified wackestones that tend to be thicker above the hangingwall (Pl. 13f). The stratified wackestone above the fault plane are deformed into steeply-inclined shear folds, covered by sub-planar non-folded beds, and then by a bed-parallel zone with meter-scale boudins, 1.52m in thickness (Pl. 13g). On the southern side of the Sefunim Valley this well stratified wackestone succession is inclined 24°-26° to the SW (Pl. 13h, marked 'A'). The stratified wackestones terminate at the top by a highly irregular surface ('SZ' in Pl. 13h). Above this surface, a further section of stratified wackestones is inclined in the opposite direction, i.e. approximately to the NE
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(marked 'B'), and draped and folded on the irregular SZ plane. Upwards, inclinations of the ‘B’ block beds gradually decrease, and ultimately become sub-horizontal (‘C’ in Pl. 13h). 5.3.5.4. Late Cenomanian events at Sefunim The meter-scale, occasionally imbricated blocks of the lower Sefunim outcrop are best interpreted as coarse debrites. The blocks were emplaced just prior to, or coevally, with coarsegrained turbidites forming the calcarenitic matrix. During deposition, the debrite blocks were semi-consolidated, as suggested by flame structures penetrating the calcarenitic matrix. Above the debrite succession are few meters of massive calcarenites answering to the description of sandy debris-flows (cf. Shanmungan 1997). Shear folds and meter-scale boudins in the wackestones of the upper part of the succession indicate mechanical instability, and downslope transport of open-marine stratified wackestones. Overloading caused by block emplacement resulted by growth faulting of underlying deposits. Thickness variations above the listric fault show that movement along it occurred during accumulation of the stratified wackestones. The NE-facing inclinations of the lower part of the ‘B’ block (Pl. 13h) suggest shear, gliding and rotation of the stratified wackestones above the curved surface ‘SZ’. 5.3.5.5. The Isfiyye section At Isfiyye in the eastern Carmel, pelagic limestones bearing the Late Cenomanian ammonite Calycoceras sp. (id. Z. Lewy, pers. comm.) are overlain by meter-sized breccias with calcarenitic blocks embedded in pelagic matrix (Pl. 6d). This unit is encrusted by a thin ferruginous crust, which is overlain by grainstones with planar bedding or hummocky crossstratification, intercalated with laminated pelagic marls with abundant pelagic microfossils. Some hippuritid rudists are present among the grainstones. 5.3.5.6. Late Cenomanian events at Isfiyye Late Cenomanian basinal pelagic chalks of Late Cenomanian age at Isfiyye were interrupted by emplacement of meter-sized calcarenitic megablocks. Examination of the calcarenites within these blocks reveals similarity in composition and fabric to the calcarenitic clino-successions of the western Galilee. These mega-clasts were re-deposited by debris flows downslope into the basin. The thin reddish crust indicates submarine omission. It is overlain by Turonian grainstone-marl couplets forming deepening-upward cycles.
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5.4. Discussion 5.4.1. The Late Cenomanian structural-depositional slope configuration of northern Israel Regional integration of the above-described cross sections (Fig. 5.9) shows that a homoclinal ramp system of the Late Cenomanian HST-2, represented by the UC3 cycles, was downfaulted at Kishor by approximately E-W striking normal fault. This Late Cenomanian normal fault is the incarnation of the much younger Yirka Fault. Faulting at Kishor caused irregularly sub-horizontal subsurface detachment, basinward to the SW. Subsidiary block rotations along the detachment at Kishor, Hamra Valley, and Mt. Gamal caused development of a chain of small extensional basins above the hangingwalls. Near Pelech, a hangingwall growth syncline developed above a ramp zone in the detachment. New accommodation space formed south of the Late Cenomanian Yirka Fault was rapidly filled by bioclastic sands generated on the shoreface north of the fault (UC3 cycles). Opposite dips between the Kishor clinoforms and the underlying beds indicate block rotation. The sediment-filled basins forming the slope system of the Galilee passed to the SW into a toe-of-slope mass transport complex including sandy turbidites, slides and debrites in the Carmel. 5.4.2. Evolution of the Late Cenomanian slope ─ toe-of-slope system of N. Israel Evolution of the Late Cenomanian slope ─ toe-of-slope system of northern Israel as a ‘shelfmargin apron system’, in respect to relative sea-level changes and tectonic events is summarized here and in Figures. 5.10 and 5.11: 5.4.2.1. Stages 1 and 2: Formation of open margins (Fig. 5.10) Disintegration of the Late Cenomanian HST-2 homoclinal ramp was the result of sedimentweakening and increased shear-stress caused by tectonic block downfaulting along the Yirka Fault and coeval Late Cenomanian eustatic fall (see Chapter 3; Voight et al. 2006). Block down-faulting at this stage resulted in differentiation of the carbonate system into three depositional bands: (1) a proximal shelf with non-rimmed open-margin, on the footwall north of the Late Cenomanian Yirka Fault; (2) a composite steep slope in the Kishor–Shagor region, underlain by fault-induced intraslope basins; (3) a toe-of-slope system in the S-SW, reaching the Carmel region. 5.4.2.2. Stage 3: Development of slope-apron foresets and base-of-slope turbiditic bottomsets (Fig. 5.10) Intersection of the fairweather wave base with the footwall shelf during base level fall stimulated bottom reworking and increased sand transportation downslope to the south. As
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shown by Plint (1988), basinward shoreface migration during sea level fall often triggers erosion of the downlap surface, here subsequently obscured by sub-aerial exposure and transgressive shoreface ravinement. Skeletal shoreface grains bypassed the non-rimmed open shelf-margin, and accumulated, coevally with block rotation (viz, evidence from Mt. Gamal outcrops), in new accommodation space created on the hangingwalls. Transportation of grains on the slope was by grain-flows, sandy debris-flows and to a lesser extent by turbidity flows. The skeletal grains accumulated on the hangingwalls as foresets dipping 20° to 35° to the W-SW, forming an upper-slope-apron. Foreset accumulation on the upper slope was coeval with intersection of fair-weather wave-base with the shelf floor. Sea-level at this stage was thus 10-20m above the top of the footwall shelf, based on recent analogues. Foresets accumulated in at least three fault-controlled extensional intraslope basins, at Kishor–Blaya, Hamra and Mt. Gamal, while the Pelech basin is compressional. The slope system was observed at least 6km downslope, from the Yirka Fault at Kishor to the Shagor Valley in the south (Fig. 5.1); its continuation as a steeply dipping progradational slope toward the S-SW is not known. However, the upper slope-apron foresets of the western Galilee pass distally into subhorizontally-stratified toe-of-slope turbiditic bottomsets in the Carmel region, 35 to 40 km to the S-SW. These bottomsets extend at least 11-12 km along the depositional strike, from the Sefunim region in the central Carmel to the Rakefet Valley in the southern Carmel (Fig. 1.1), forming an extensive turbidite sheet. The bioclastic- and peloidal grains of which they are composed originated on the footwall shelf of the western Galilee. Failure on oversteepened slopes in the Galilee initiated downslope mass movement events, of unconsolidated grains flowing toward the Carmel. Debrites and consolidated stratified blocks with shear zones and boudins are embedded within the Carmel turbidite sheet. Toe-of-slope deposits in the Carmel region overlie late Mid- or Late Cenomanian basinal successions, and are topped by Turonian deepening-upwards cycles (Isfiyye). Their age is therefore bracketed as Late Cenomanian, coeval with the slope-apron of the western Galilee. Unsurprisingly, blocks within the debrites are often composed of the same coarse grainstones found on the slope-apron foresets of the western Galilee (e.g. Isfiyye debrites, Pl. 6c). 5.4.2.3. Stage-4: Sea level fall, followed by drowning of the slope-apron (Fig. 5.10) The sequence, overlying the Ce SB-3 and truncated by the Ce SB-4, is a window into the latest Cenomanian depositional phase. The Pelech sequence is folded into syndepositional hangingwall syncline above a ramp segment of complex listric detachment. The Ce SB-3 reflects eustatic sea-level fall (see Chapter 3) to below the shelf-break of Kishor, with subaerial exposure, erosion and removal of calcarenites from the Pelech
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synclinal basin. This stage was followed by eustatic rise during which outer- and mid-ramp muds of the Pelech sequence accumulated within the Pelech syncline and beyond. 5.4.2.4. Stage-5: tectonic uplift and faulting (Fig. 5.10) The Ce SB-4 is a second episode of subaerial erosion. Subaerial exposure and erosion were recorded in the western Galilee (Mt. Gamal, Hamra Valley) but were not recorded in the distal Carmel region. There (Isfiyye outcrops), the Ce SB-4 corresponds to a ferruginous crust indicating submarine omission without evidence for exposure. As subaerial exposure features of this stage are localized in the Galilee and cannot be widely correlated across the marginal Tethys (Chapter 3), tectonic uplift restricted to the Galilee is indicated. The faulting of spillover deposits of the Pelech sequence at Hamra may belong to this episode. 5.4.2.5. Stage-6: Termination by Drowning (Fig. 5.11) Four retrogradational stages record drowning of the Galilee–Carmel slope–toe-ofslope system: (1) the first retrogradational stage (stage 6.1 in Fig. 5.11) is recorded by Early Turonian UC8 and UC12 deepening-upward cycles of the Muhraqa and Isfiye sections (Figs. 2.9, 3.9). At this stage, the Carmel fluctuates between the basin and the lower shoreface and the elevated Galilee is still subaerially exposed. (2) The second retrogradational stage (stage 6.2 in Fig. 5.11) corresponds to deposition of Early Turonian basinal marls of zones T4 and T5 of the early Turonian Daliyya Fm in the Carmel, and the lower part of the basinal Yirka Fm in the Galilee. This stage is attributed to partial onlap on the Galilean slope, while the footwall shelf is still exposed. (3) The third retrogradational stage (stage 6.3 in Fig. 5.11) corresponds to flooding of the footwall shelf in the Galilee, reworking of the shelf floor by wave base and formation of retrogradational shoreface deposits of top Yanuch Fm on the footwall shelf. Some reworked carbonate grains produced at this stage were redeposited as turbidites beyond the shelf-break (in the Yirka Fm of the Hamra region). Absence of turbidites at this stage in the more distal Mt. Gamal basin suggests ponding of turbidites in the Hamra basin and possible flow along the basin strike. (4) The final stage (stage 6.4 in Fig. 5.11) corresponds to further rise of sea level with deposition of deep water marls over the footwall shelf. 5.4.3. Comparison of the Galilean carbonate shelf-margin apron with siliciclastic deltaic systems In slope aprons of the Little Bahama Bank (Mullins & Neuman 1979; Hine et al. 1981; Hine 1983), and in ancient systems (e.g., Swinchatt 1967; Pomar et al. 2005), calcareous sand from a linear source is transported off-bank by storms and waves and ultimately
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accumulates on a smooth slope or inside intra-slope depressions. Sedimentary geometries in these systems are controlled by inherited topographies (reefs/buildups/Pleistocene remnant topography) and only a few ancient examples show thick growth-fault controlled skeletal sand-aprons such as those described above. The few examples are from the Jurassic of the Pre-Alps (Abbots 1989) where oolitic sands, 800 to 1000m in thickness, accumulated as huge turbiditic lobes, and from the Cretaceous of south Italy, where bioclastic aprons accumulated on the faulted slope of the Apulian platform during highstands (Borgomano 2000). The geometry, stacking patterns and mechanism of formation of the Late Cenomanian slope clinoforms of the western Galilee are in some ways more similar to those described in siliciclastic shelf-margin deltas (e.g., Edwards 1981; Porębski & Steel 2003; Cummings & Arnott 2005). Common features of the both are: (1) thick and steep grainy clinosuccessions located between a proximal sediment source and a distal-basin; (2) seaward steepening and thickening of clinoforms (30-35m, 20°-25° in the proximal Kishor to at least 80-100m, 30°-35° in the more distal Hamra Valley), (3) Growth-fault control on the formation of thick clino-successions; (4) bounding of calcarenitic slope deposits by erosion surfaces merging proximally; (5) deposition of clino-successions in a forced regression setting. Moreover, the sedimentary architecture of the Galilee–Carmel slope – toe-of-slope system is comparable to that of Gilbert-type deltas (e.g. Colella 1988). The upper slope clinosuccession in Kishor-Blaya (Galilee) is stacked as lower foresets topped by cross-bedded topsets (see Kishor section in Fig. 5.2; Pl. 10g), while at the distal Carmel sub-horizontallystratified turbidites form bottom-sets (e.g. Pl. 13a). The similarities of the stacking-patterns between carbonate sand lobes, and those of fluvial-controlled Gilbert-type fan deltas were also reported by Vecsei (1998) from the upper Cretaceous of the Maiella Platform of Italy. The sheeted turbiditic deposits of the Carmel also recall siliciclastic deltaic bottomsets, or carbonate outer-aprons or sand-rich slope-aprons often developed at the base of slopes, especially during falling stages of sea level (e.g., Mullins & Cook 1986; Reading & Richards 1994). This system can best be described by a combination of two depositional models: (1) off-bank sand transportation from a linear sediment source (shoreface) located on an open margin, as in carbonate-apron models (Mullins & Cook 1986); (2) growth-fault controlled sedimentary geometries, facies distribution and stacking patterns, as in shelf-margin deltas (e.g., Porębski & Steel 2003). The slope- to base-of-slope system of the Galilee–Carmel corresponds therefore to a 'shelf-margin apron', with topsets, foresets and remote bottomsets. The purely carbonate calcarenitic lithology of the shelf-margin apron controls some of its key characteristics. Sediment production and transport was controlled mainly by
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autochthonous biogenic processes, and shoreface and slope dynamics are often modified by early cementation because of their calcitic/aragonitic mineralogy. Carbonate slopes tend to be steeper than siliciclastic foreset slopes (Schlager & Camber 1986), and mud-free grainsupported grainstones and rudstones may be as steep as 40° (Kenter 1990). The steep gradient of the Galilean slope-apron was therefore primarily due to the coarse grainy fabric. Furthermore, the earliest cements recorded in these grainy slope-deposits are early marine phreatic fibrous rim cements (Pl. 10f), that would have enhanced long-term slope stabilization and steep gradients. 5.4.4. 'Fill and Spill' sedimentary dynamics The down-dip configuration of the slope basins in respect to a prograding linear sediment source in the north (Yanuch, Beit HaE’emek) suggests that complete fill of the upslope basin (the proximal Kishor–Blaya basin) must be completed before grain progradation continued into the next downslope basin (Hamra and Mt. Gamal basins). This progressive down-dip filling process corresponds to the “fill-and-spill” mechanism described by Satterfield & Behrens (1990) and Beauboeuf & Friedmann (2000) from siliciclastics of the northern Gulf of Mexico. The exact mechanism of basin fill and the ensuing sediment spill, as well as the connection of this process to the continuously falling fairweather wave base, is reflected in sedimentary structures in each small basin (Fig. 5.9). The SW-dipping foresets filling each basin are overlain by the cross stratified part at the proximal Kishor-Blaya basin, and by a channelized part at the distal Hamra basin (Fig. 5.2). Within each basin, vertical stacking reflects the upward shift to equilibrium. In the proximal Kishor–Blaya basin, sediment transport into the downslope basin (Hamra) was by shoreface progradation, indicating that the continuously falling fairweather wave base intersected the sedimentary interface during spill. On the other hand, in the more distal Hamra mini-basin, spilling took place through slope channels.
5.5. Summary 1. The E-W striking Yirka Fault of the western Galilee was active during Late Cenomanian times. Rotational downfaulting along this fault toward the south formed the Late Cenomanian shelf-edge of the Galilee, previously characterized by homoclinal geometry. Tectonic faulting was simultaneous with Late Cenomanian eustatic fall, resulting by sediment instability and development of a composite detachment toward the south. Subsidiary block movements along the detachment
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formed at least three extensional basins above hangingwalls and a single syndepositional hangingwall syncline. 2. The new basins formed south to the Yirka Fault were filled by coarse calcarenitic grains originated on the footwall-shelf in the north. Carbonate grains were deposited as W-SW facing clino-successions, becoming much thicker and steeper to the distal south, with evidence supporting syn-deformational basin fill. Slope oversteepening triggered mass-movements of unconsolidated grains, redeposited on the toe-of-slope of the Carmel as an extensive turbiditic sheet. Other types of mass transport features, such as debrites, shear zones and internally deformed slides also characterize the toe-of-slope in the Carmel range. 3. The Late Cenomanian structural infrastructure divided northern Israel into a wave base-controlled footwall shelf north of the Yirka Fault, a fault-controlled composite calcarenitic slope system in the western Galilee, and a turbiditic toe-of-slope in the Carmel in the SW. 4. The loose carbonate grains forming the upper slope foresets of the Galilee filled the small slope basins to equilibrium point and were then spilled over the basin edges into the next downslope basin. Sediment spill in the proximal regions was by shoreface progradation and in the more distal regions by mass-movements of grains inside channels. 5. The purely carbonate Galilee–Carmel slope – toe-of-slope system show similarities with siliciclastic shelf-margin deltas and can be best described by a combination of two depositional models: the carbonate-apron model and the siliciclastic shelf-margin deltas model. This system corresponds therefore to a 'shelf-margin apron', with topsets, foresets and remote bottomsets.
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Chapter 6 Sedimentary configuration and tectonic framework of the midCretaceous northern Arabian platform, and their paleoceanographic implications
6.1. Introduction The Levantine margins of the Neotethys (Fig. 6.1) were shaped by Late Permian, Middle to Late Triassic, and Early Jurassic rifting events (Garfunkel 1998). Normal faults associated with this rifting strongly influenced the depositional strike of the Levant margin. Cessation of rifting, lengthy tectonic quiescence, and slow passive margin-type subsidence prevailed from the Mid-Jurassic to the Middle Cretaceous (Garfunkel 1998; 2004). Post MidJurassic margin geometry exercised control on carbonate facies transitions across this margin (Gvirtzman & Klang 1972; Bein & Gvirtzman 1977), along a narrow belt termed the ‘Levantine hinge-line’ or ‘hinge-belt’. E-W trending proximal-to-distal facies transitions for the mid-Cretaceous occurred perpendicular to this N-S trending rimmed shelf-margin. They were described from the Carmel region in the north to the southern coastal-plain of Israel (Fig. 6.1) (Bein 1971,1974; Bein & Weiler 1976; Sass & Bein 1982). A sharp transition took place along this zone, with shallow water carbonates and carbonate buildups in the east passing into deep water laminites, turbidites and contourites of the Talme Yafe Fm (Cohen 1971; Bein & Weiler 1976) to the west. North of the Carmel region, in the Galilee, the continuation of the hinge belt was not described. The mid-Cretaceous carbonate system was considered a wide carbonate platform transected by intra-platformal basins (Freund 1965; Kafri 1972 1991). Walley (1998) considered the Galilee and southern Mt. Lebanon to be a zone of interruption in the hinge belt, which reappears a few tens of kilometers to the north, parallel to the N-S trend of the western Lebanon flexure. This ‘zone of interruption’ was considered a western continuation of the southern Palmyride rift basin of Syria, this continuity is more apparent upon restoration of late left-lateral movement along splay faults belonging to Dead-Sea transform system. In the present chapter the sedimentary-tectonic configuration of the northern Levantine carbonate system are reconstructed and paleoenvironmental implications are analyzed. The high-resolution sequence stratigraphic framework for the Galilee and Carmel of northern Israel (see Chapter 3) is here extended to Lebanon and Syria. Facies and thickness trends were found to diverge from the N-S – directed hinge-belt, and depositional strikes of major sedimentary-tectonic elements of the Cenomanian-Turonian carbonate
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system in this region trend W-SW─E-NE. Additionally, the sedimentary-tectonic framework gives rise to large-scale fluctuations in the position of the Cenomanian-Turonian OMZ on and off the ramp corresponding to two eustatic rises of sea level: in the upper Mid-Cenomanian and in the Late Cenomanian.
6.2. Thickness patterns of the Cenomanian succession in the northern Levant Figure 6.2 presents SSW-NNE and SW-NE directed cross sections showing thickness patterns of the Cenomanian succession in the northern Levant. Cross-section 6.2a extends from the Carmel in the south, across the western and northern Galilee, and into central Lebanon and western Syria in the north. Cross-section 6.2b extends from the Carmel in the southwest, across the Palmyride basin in Syria to the NE. The cross sections are bounded at the base by the Alb-Ce sequence boundary and at the top by a disconformity at which the latest Cenomanian sequence boundary is placed in Israel. Disconformity surfaces at these horizons are also described in Lebanon (Ferry et al. 2007) and Syria (Mouty et al. 2003). The Cenomanian succession shown in cross-section 6.2a is more than 400 m thick at the Carmel (Kashai 1966), is relatively thin in the Galilee (140m at Betzet), and is more than 500m thick in Mt. Sannine of central Lebanon (Saint-Marc 1972). Farther to the north, in the Safita and Slenfe sections of western Syria, the Cenomanian is around 350m in thickness (Krasheninikov 2005). The Cenomanian succession in cross-section 6.2b similarly begins with the 400m of the Carmel (op. cit), reaching 700m in the region of Bloudan (Ponikarov et al. 1967; Krasheninikov 2005). Although the Bloudan section is located in the central AntiLebanon range, it is east of the Serghaya fault (Fig. 6.1), the easternmost left-lateral fault of the Dead-Sea transform system (DST) (Gomez et al. 2001). Restoring 105km of total postCretaceous left-lateral shear along the DST faults (Freund et al. 1970), the Bloudan section and also sections in the Palmyrides are shown near their original mid-Cretaceous position in Figure 6.1b. NE of the (restored) Bloudan section, the Cenomanian at Mt. Abu-Zounar of the SW Palmyrides is ~330m in thickness, and thins to the NE to 225 to 300m (Palmyride sections from Mouty et al. 2003). Thickness changes may be attributed to variations in productivity of the carbonate system, stratigraphic condensation, limitations of accommodation space, or variations in tectonic subsidence. In this chapter, the facies-based, high-resolution sequence stratigraphic framework of northern Israel presented in Chapter 3 (Figs. 3.1, 3.2) is extended to central Lebanon, western Syria and the Palmyrides using published data (Fig 6.3). The sources of these data
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are Dubertret (1937), Ponikarov (1967), Saint-Marc (1972 1974), Mouty & Al-Maleh (1983), Mouty et al. (2003), Krasheninikov (2005) and Nader et al. (2006). In order to narrow the information gap caused by different resolutions inherent in this type of data, the sequence stratigraphic technique suggested by Embry (2002), using broad-based transgressiveregressive sequences, is employed. This technique, also used in Chapter 3, uses maximumflooding surfaces and sequence boundaries as correlation surfaces, and integrates highstand, forced-regressive and lowstand system tracts into an inclusive 'Regressive System Tract' (RST). Five such transgressive-regressive systems tracts (Ce TST-1; Ce RST1; Ce TST-2; Ce RST-2; Tu TST) are identified in Israel and extended regionally as in Figure 6.3.
6.3. Sedimentary configuration in the Ce TST-1 6.3.1. Characteristics of Ce TST-1 Ce TST-1 of northern Israel is an Early Cenomanian outer-ramp to basinal openmarine, pelagic and hypoxic system composed of well-bedded to laminated, sometimes pyritic, calcisiltite wackestones and packstones (see Chapter 3). In the NE Galilee, Ce TST-1 is relatively thin and of a proximal aspect, passing laterally into thicker and more distal facies towards the Carmel region in the SW, and towards the west (Fig. 3.4).
6.3.2. SW-NE traverse: Ce TST-1 along the Carmel–Palmyride trend The Isfiye Fm chalks of central and northern Carmel region are 120m thick and belong to the Early Cenomanian R. brotzeni Zone (Bein 1974; Lipson-Benitah et al. 1997). They transgressively overlie the Albian/Cenomanian sequence boundary. The Isfiye chalks are well-bedded, platy, or laminated, with some pyrite and glauconite, representing a relatively-starved, hypoxic environment on the distal slope or basin. Figure 6.4a shows the Carmel succession passing toward the NE into a 280m thick succession of pelagic limestone-marl couplets, represented by the Bloudan section (Krasheninnikov 2005). This section is dated by planktonic foraminifera of the Early Cenomanian Rotalipora brotzeni Zone and the early Mid-Cenomanian R. reicheli Zone. The restored Bloudan succession passes toward the Palmyrides into the marly lower part of the Abu Zounnar Fm (Fig. 6.3; marne d’Abou Zounnar of Mouty & Al-Maleh 1983) with Early Cenomanian ammonites Acompoceras spathi and Mantelliceras sp. This part of the Abu-Zounnar Fm was considered by Mouty et al. (2003) as deep subtidal, the deepest part of the Cenomanian succession of the Palmyrides. The base of the Abu Zounnar Fm is transgressive over the upper Zbeide Fm, and is overlain by shallow water bioturbated
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limestones and dolomitic limestone of the upper Abu Zounnar Fm. Generally the Abu Zounnar marls become thinner and more proximal toward the NE Palmyrides: At Abu Zounnar in the SW, the thickness is ~100m in deep subtidal facies, while at Mt. Satih in the NE this unit is only ~20m thick, dolomitized, and apparently more proximal (Fig. 6.3). The Carmel-Palmyride trend of the SW-NE traverse is summarized in Figure 6.4a. The Early Cenomanian succession is hypoxic and relatively condensed in the Carmel and Galilee, and passes into more oxic, basinal-marl couplets at Bloudan, deep subtidal facies in the SW Palmyrides, and shallow-subtidal in the NE Palmyrides.
6.3.3. S-N Traverse: Ce TST-1 in the Carmel – Galilee – central Lebanon - NW Syria trend (Fig. 6.3) The 120m-thick basinal Isfiye chalks of the Carmel become thinner towards the NE Galilee. They pass to 50m of basinal chalk at Betzet, then to 15m of laminated basinal limestones at Manara. At Mt. Sannine of central Lebanon to the north, the lower ~130m of the Sannine Fm is Early Cenomanian, based on benthic foraminifera (Saint-Marc 1972), with 30-35m of marly-limestones at the base, corresponding to the Early Cenomanian TST (Ce TST-1). The Early Cenomanian maximum-flooding surface should be placed at the transition to the overlying regressive dolomites. Ce TST-1 is basinal and much thicker at Mt. Nusseiriyeh, western Syria. In the Safita section of southern Mt. Nusseiriyeh (Fig. 6.3) 20-30m of platy-marls of the Maisra Mb (Member) (Kozlov 1966) bears the Early Cenomanian ammonite Mantelliceras mantelli. The Maisra Mb overlies bioclastic rudist limestone with Eoradiolites liratus and E. syriacus, identical to the Albian-Cenomanian transition in the Carmel, Manara and parts of Lebanon (Fig. 6.3). Above the Maisra marls occur 100m of open-marine, fine-grained, well-bedded chert-bearing limestones also attributed to Ce TST-1. This basinal succession is bounded at the top by the Mid-Cenomanian, regressive ‘Slenfe Limestone’ (Dubertret 1937). Approximately 70km north of Safita, at Slenfe of NW Syria (Fig. 6.3), more than 100m of Early Cenomanian marls and bedded limestones directly overlie latest Albian bioclastic rudist limestone. The Early Cenomanian succession is also basinal in this region, with planktonic foraminifera of the Early Cenomanian R. brotzeni zone, and the Early Cenomanian ammonites Sharpeiceras laticlavium and Mantelliceras mantelli. As at Safita, this succession is again overlain by the regressive Mid-Cenomanian ‘Slenfe Limestone’. The Carmel-Galilee-W Syria trend of the S-N traverse is summarized in Figure 6.4b. A relatively condensed facies in the Galilee and central Lebanon passes into thicker, deeperwater facies both towards the Carmel in the south and towards western Syria in the north.
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6.3.4. Ce TST-1 basin configuration in the northern Levant (Fig. 6.4c) Reduced thicknesses in the NE Galilee, central Lebanon (Mt. Sannine) and NE Palmyrides are consistent with relatively elevated and condensed facies. Thicker accumulations of basinal deposits indicate two rapidly subsiding regions. One extended south of the NE Galilee–central Lebanon elevation, from the southwesternmost Palmyrides to the Carmel. The second was north of the NE Galilee–central Lebanon elevation, in northern Lebanon and western Syria.
6.4. Sedimentary configuration in the Ce RST-1 6.4.1. Characteristics of Ce RST-1 Ce RST1 of northern Israel is a prograding system composed of a highstand, and forced regressive system tracts. The resolution of data from published work on Lebanon and Syria is insufficient for such a distinction and the entire system is treated as a single regressive system tract, following Embry (2002). Fully developed successions in northern Israel show a thick shallowing-upward cycle with typical stacking patterns indicating gradual filling of accommodation space to sea-level (Fig. 2.10a). This relatively shallow water succession is bounded between deep subtidal to basinal facies of the Ce TST-1 below (of the Deir Hanna Fm), and upper Mid-Cenomanian to Late Cenomanian drowning succession of the Ce TST-2 above (of the Yanuch or Bina Fms). Both the underlying Ce TST1 and overlying Ce TST2 are dated by ammonite and planktonic foraminiferal biostratigraphy, so Ce RST1 is constrained to an Early to Middle Cenomanian age.
6.4.2. SW-NE traverse: Ce RST-1 along the Carmel –Palmyride trend The Palmyrides-Carmel trend of the Ce RST1 is summarized in Figure 6.5a. In the central and northern Carmel, Ce RST1 corresponds to the Early- to Mid-Cenomanian part of the chalk succession of the Arkan Fm (Segev et al. 2002). It is fully basinal, and intercalated with volcanics (e.g. Sass 1980; Lipson-Benitah et al. 1997). Considerable submarine erosion at the top of this system truncates the Arkan Fm (Bein 1974; Lipson-Benitah et al. 1997). This truncation is the off-ramp basinal expression of the Ce SB-2 (Mid-Cenomanian sequence boundary) that becomes a subaerial unconformity in the northern Galilee (Fig. 3.5c). The Carmel succession passes to the NE into the thick shallow subtidal succession of Mt. Kedumim of southern Galilee, and farther to toward the NE the succession is the 120m of shallow-subtidal fossilerous limestones and dolomitic limestones at Bloudan, SW of the
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Palmyrides (Fig. 6.3). The shallow water limestones of Bloudan overlie basinal limestones of the R. reicheli planktonic foraminiferal zone and are therefore Middle to Late Cenomanian in age. This succession passes to the Palmyrides in the NE into the upper part of the Abu Zounnar Fm (Falaise d’Abou Zounnar of Mouty & Al-Maleh 1983), also composed of bioturbated fine to coarse grained bioclastic mudstones and wackestones, and dolomitic limestones. Mollusk-sponge debris, corals, pellets and dolomitized oolites were reported.
6.4.3. S-N Traverse: Ce RST1 along the Carmel – Galilee – central Lebanon – western Syria trend The S-N Carmel-Safita trend is summarized in Figure 6.5b. Ce RST1 of the southern Carmel (Zikhron Fm, Hotem Carmel section; Fig. 6.3) is composed of mid-ramp and peritidal dolomites. These shallow water dolomites pass toward Mt. Kedumin of the southern Galilee into very thick succession composed of ~200m of well-stratified shallow subtidal dolomites of the Sakhnin Fm, passing further northwards into the mid-ramp to peritidal system in the Galilee. In the north Galilee and southern Carmel (Hotem Carmel section), Ce RST1 is a distinct shallowing-upward cycle that ultimately equilibrates with sea level in the peritidal environment (CC1 cycle in Fig. 2.10a). The peritidal system of the Galilee passes toward Manara in the north into a deeper mid-ramp rudist-gastropod 'carbonate-factory' facies 20m in thickness. This Galilean trend is represented in Figure 6.3, by the transition from Dishon to Manara section. The mid-ramp facies of Manara passes northwards toward Mt. Sannine of central Lebanon, into a 200m-thick succession of dolomites, considered as upper Early to MidCenomanian by Saint-Marc (1972). The Early Cenomanian part of this succession is composed of 95m of white fine-grained massive dolomite with quartz geodes and the overlying succession is well-stratified fine-grained dolomites. The vertical stacking pattern revealed at Sannine, of massive dolomites overlain by well-stratified dolomites is essentially the same as at Ce RST1 of Mt. Kedumim of the southern Galilee (Fig. 6.3). In both localities, stacking patterns and thicknesses are identical and the facies is very similar, composed of well stratified dolomites, occasionally brecciated, with insubstantial peritidal features. Facies in both regions represent subtidal mid-slope environments. North of Mt. Sannine, at Safita and Slenfe of Mt. Nusseiriyeh, western Syria, the Early - to Mid-Cenomanian succession of Ce RST1 is represented by the subtidal 'Slenfe limestone' (Dubertret 1937). At Safita it is described as 'coquinal' (presumably bioclastic), while more distally at Slenfe to the north, it is finer grained (Krasheninikov 2005).
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6.4.4. Ce RST1 basin configuration in the northern Levant (Fig. 6.5c) The southern Carmel and parts of the Galilee were two relatively elevated peritidal platforms separated by a rapidly subsiding basin extending from a basinal Carmel to a midslope at Mt. Kedumim and to a shallow subtidal Bloudan and SW Palmyrides (Fig. 6.5a). The peritidal high of the southern Carmel (Hotem-Carmel section; Fig. 6.3) bounded this basin at the SW and probably passed gradually updip toward the NE into the exposed Hamad Uplift (Mouty & Al-Maleh 1983). The Galilean peritidal high bounded the Carmel – SW Palmyride basin at the north, but also descended northward, towards the rapidlysubsiding region of central Lebanon (Sannine). The thick successions of shallow subtidal carbonates of both Mt. Sannine in the north and Mt. Kedumim in the south accumulated in the rapidly subsiding regions south and north of the Galilean peritidal high. West of Mt. Kedumim, the subtidal slope passed into a stagnant basin represented in the Carmel by the Arkan chalk succession. The thick Mt. Sannine section of central Lebanon thins towards Safita and Slenfe at the north, becoming an outer ramp facies represented by the ‘Slenfe limestones’ (Fig. 6.5b).
6.5. Sedimentary configuration in the Ce TST-2 6.5.1. Characteristics of Ce TST-2 Ce TST-2 in northern Israel accumulated above the Ce SB-2 (Mid-Cenomanian SB) of the Galilee and Carmel (Fig. 6.3) and it is late Mid-Cenomanian to Late Cenomanian in age (Chapter 3, Fig. 3.3). The transgressive facies of this system is variable across the Galilee, onlapping a collapsed ramp bathymetry formed during the previous Ce RST-1. It generally corresponds to open marine subtidal or basinal facies bounded by MidCenomanian regressive dolomites at the base, and Late Cenomanian regressive limestones or dolomite at the top. Condensation due to sediment starvation characterizes this stage in both proximal and distal environments in the NW and central Galilee. Local features include occasional massive influx of pithonellid calcispheres, bioerosion, bioturbation and grain bioerosion, occurrences of phosphatic grains, ferruginous encrustations and silicification.
6.5.2. SW-NE traverse: Ce TST-2 along the Carmel –Palmyride trend Ce TST-2 in the west-central Carmel (e.g. western Oren Valley) corresponds to the lower part of the Muhraqa Fm (Carmel composite section; Fig. 6.3), overlying a submarine truncation surface (Bein 1974) that is the distal expression of the Middle Cenomanian
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sequence boundary (Ce SB-2). It is composed of 35 to 40m of fine-grained packstones with abundant pelagic microfossils deposited on a distal slope or basin floor. Ce TST-2 in the southern Carmel (Hotem Carmel) and southern Galilee (Mt. Kedumim) is mapped as the lower part of the Bina Fm. It overlies well-stratified peritidal to mid-slope dolomites of the Ce RST-1 (Sakhnin Fm; Fig. 6.3) and consists of bioturbated finegrained wackestones with echinoderm debris, hyaline benthic foraminifera (mostly rotaliids) and pithonellid calcispheres of open-marine outer-ramp facies. In the restored Bloudan section, which is the projection of this part of the succession to the NE, there was insufficient data to allow recognition of a distinct Ce TST-2. Further to the north in the Palmyrides, Ce TST-2 corresponds to the lower Abtar Fm (la sèrie stratifiée of Mouty & Al-Maleh 1983), dated by Mouty et al. (2003) as early Late Cenomanian. Although this unit was interpreted by Mouty et al. (2003) as “littoral”, slightly deeper than, and transgressive over the upper Abu Zounnar Fm, the lower Abtar Fm at Satih and Abu-Zounnar (Fig. 6.3) was described as well-stratified or laminated with open-marine pelagic fossils such as pithonellids and gavelinellid foraminifera, all characteristic of slope deposits. Marine hardgrounds with bored surfaces and limonitic crusts are common. Pronounced stratification, lamination and abundant hardground features may reflect frequent omissions. The Ce TST-2 in the Carmel – Palmyride trend is summarized in Figure 6.6a. The pelagic basinal facies of the west-central Carmel passes toward the E-NE into the more proximal bioturbated outer-ramp of the southern Galilee (and southern Carmel as well). This bioturbated facies passes to the NE into a stagnant mid-ramp facies in the Palmyrides.
6.5.3. S-N traverse: Ce TST-2 in the Carmel – Galilee – central Lebanon – western Syria trend The 30-40m of deep-basinal Ce TST-2 of the lower Muhraqa Fm of the Carmel (described above) passes to the Galilee in the north into a few meters of condensed, deepening-upward cycle that pass toward Manara (Fig. 6.3) in the north, into at least 60m of well-stratified to well-laminated basinal chalks. The transgressive chalks of Manara pass toward Mt. Sannine of central Lebanon into the upper part of the Mid-Cenomanian succession (Saint-Marc 1972) overlying regressive dolomites of the Ce RST-1 (Fig. 6.3). This part of the Mt. Sannine succession is 55-65m in thickness, and composed mainly of bedded, fine-grained or bioclastic limestones with chert bands and nodules. Slope foraminifera such as orbiotolinids and hyaline rotaliids were reported. Chert-bearing limestones from a similar stratigraphic position in the Sannine Fm of
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northwestern Lebanon were considered basinal by Nader et al. (2006). Radiolarians, pithonellid calcispheres and some pyrites are found in this facies, as well as an estimated high content of organic matter. The Ce TST-1 at Sannine thickens toward Safita in the north, to more than 100m of thin-platy to thick bedded limestones with chert and clayey horizons, echinids, nautiloids (Eutrephoceras sublaevigatum) and the late Mid-Cenomanian ammonite Calycoceras boulei. Outer-ramp or basinal origin is proposed for this succession that directly overlies the regressive ‘Slenfe limestones’ (Fig. 6.3). To the north at Slenfe, a similar Middle- to LateCenomanian basinal succession occurs. West of the basinal successions of Mt. Sannine–Slenfe–Safita, toward the northern Anti-Lebanon, the late Mid- to Late Cenomanian succession thins into 15-25m of bioturbated fine-grained bioclastic limestones of the Al-Koroum Mb. These limestones overlie bioturbated dolomites related to Ce RST1, and bear echinoids, bivalves and late Middle- to Late Cenomanian ammonites Neolobites fourtaui and Neolobites vibrayeanus. They were deposited in an open-marine outer ramp. The N-S Carmel – Galilee – Mt. Sannine – Slenfe trend is summarized in Figure 6.6b. The proximal but condensed sections of the Galilee were elevated compared to deeper basinal environments of the Carmel in the south, and thicker basinal succession of the northern Galilee, Lebanon and western Syria.
6.5.4. Ce TST-2 basin configuration in the northern Levant (Fig. 6.6c) The relatively proximal and condensed Galilean paleohigh was bounded both to the south and north by basinal regions. The deeper southern region is represented by basinal facies in the Carmel in the SW, shallowing gradually toward the Galilee in the north and toward the Palmyrides in the NE. The deeper northern region extended from Manara to westcentral Syria, with much thicker and more basinal successions, commencing already in the Galilee. Increased subsidence is invoked for both basinal trends.
6.6. Sedimentary configuration in the Ce RST-2 6.6.1. Characteristics of Ce RST-2 As for Ce RST-1, the Late Cenomanian highstand system tract Ce HST-2 and the overlying latest Cenomanian forced regressive system tract (FRST), are integrated into the Ce RST-2, following Embry (2002).
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Ce HST-2 is a homoclinal ramp composed of peritidal cycles in the NW and central Galilee (e.g. Betzet) and becoming deep-subtidal towards the southwestern and northeastern Galilee (Fig. 3.7a). In the Ce FRST-2 the homoclinal ramp metamorphosed into an openmargin shelf with a thick S-SW–facing calcarenitic apron (details in Chapter 5; Figs. 5.10, 5.11). Late Cenomanian synsedimentary faults developed beneath and within this apron, with significant southward thickening of slope-clinoforms accumulating on the down-faulted blocks. Loose carbonate grains were transported downslope, deposited as an extensive turbitide sheet in the toe-of-slope of the Carmel.
6.6.2. S-N traverse: Ce RST-2 in the Carmel – Galilee – Central Lebanon trend The Carmel-Galilee-Mt. Sannine trend is summarized in Figure 6.7a. The deepsubtidal ramp of the Ce RST-2 of the NE Galilee passed to central Lebanon in the north into 225m of progradational limestones of Mt. Sannine (Saint-Marc 1970, 1972; Fig. 6.3), with an upper rudist-bearing part considered to represent the source of a thick lower part composed mainly of bioclastic debris. This succession is regressive over the underlying Ce TST-2 of Mt. Sannine.
6.6.3. SW-NE traverse: Ce TST-2 in the Carmel – Palmyride trend The lateral equivalent of the allochthonous slope deposits of the Carmel region toward the far NE is unclear since sedimentary data from the southern Galilee and the restored Bloudan section is missing. The allochthonous slope deposits of the Carmel were sourced in the western Galilee to the north, and a connection to facies in the NE cannot be confirmed. Figure 6.7b therefore expresses the assumed facies configuration along the Carmel – Palmyride trend. The distal slope of the Carmel passes in some way to the upper part of the Late Cenomanian Abtar Fm of the Palmyrides, considered by Mouty et al. (2003) to be a low-energy subtidal facies regressive over the lower Abtar Fm. Facies in this unit becomes more proximal toward the NE Palmyrides.
6.6.4. Ce RST-2 basin configuration in the northern Levant (Fig. 6.7c) Ce RST-2 of the Galilee developed on a peritidal – subtidal shelf with a faultcontrolled slope system composed of bioclastic calcarenitic clinoforms. North of the Galilean shelf, bioclastic calcarenites were similarly shed into the rapidly subsiding depocenter of central Lebanon forming the thick bioclastic accumulation of Mt. Sannine. The basinal facies of the Carmel was bounded to the north by the elevated Galilee and passed to the NE into
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‘deep-subtidal’ (outer-ramp), then ‘shallow-subtidal’ (mid-ramp) facies of the Palmyrides, bounded to the SE by the Rutbah High. Features of the paleohigh of the Galilean Ce RST-2 are traceable to the upper Abtar Fm of NE Palmyrides. Facies of both regions are shallow-subtidal (also peritidal in the Galilee). A subaerial unconformity at the Cenomanian-Turonian transition interval terminates Ce RST-2 in both regions. In the western Galilee, pedogenetic and karst features are present. In the NE Palmyrides, erosional features and deposition of quartz sandstone are reported (Ponikarov et al. 1967; Mouty et al. 2003; Krasheninikov 2005). These features indicate that both regions were relatively elevated, with the Galilee being the western extension of a NE Palmyride paleohigh (Fig. 6.7c).
6.7. Sedimentary configuration in the Turonian TST 6.7.1. Geometry of Turonian TST in northern Israel The 70-80m of open-marine and pelagic deposits of the Tu TST-1 of the western Galilee become highly condensed on the Galilean paleohigh, but then thicken downslope toward the eastern Galilee (Dishon) where they are ~60m in thickness (Fig. 6.8a; discussed in Chapter 3).
6.7.2. The Early Turonian succession of northern Israel and northern Levant – earlier model Freund (1961, 1965) showed that in northern and southern Israel Early Turonian ammonites are concentrated in several NNE-SSW to NE-SW narrow belts, some described as wedging-out lenses. These belts were considered intra-platform synclines or basins. One of them is the Yirka intraplatform basin extending from the Carmel via western Galilee into eastern Lebanon. Thinning-outs of lithosomes, or ammonite-free/depleted zones, were interpreted as elevated shallow-water margins or shorelines of the intraplatform basins. On the other hand, marginal shallow-water environments should bear shallow-water features such as peritidal flats, lagoonal foraminiferal facies, shoreface systems or biogenic buildups. Spatial distribution of ammonites is usually controlled by distance from shore (as suggested by Freund 1961) but might be also depend on post-mortem transport, changes in the rate of background sedimentation (e.g. sedimentary dilution), local diagenetic effects such as distructive late dolomitization, or paleoecology (e.g. Reboulet et al. 2005). Ammonite accumulations in thin pelagic units occurring with secondary mineralization characterize
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environments of sediment starvation well offshore. In fact, ammonites are rarely found in nearshore shell beds.
6.7.3. Basin configuration in the Turonian TST in the northern Levant The widespread distribution of ammonites in the Early Turonian is a regional phenomenon and Turonian ammonite-bearing facies appear also in the NW Palmyrides (Fig. 6.8b). The highly-pelagic Early Turonian open-marine facies of the northern Levant (Carmel, western Galilee, Lebanon; Fig. 6.8b) passes toward the SW into Early Turonian dolomites on the margins of the Rutbah uplift. These marginal dolomites are occasionally described by Ponikarov et al. (1967) and Krasheninikov (2005) as thin-bedded or well-stratified. They appear analogous to the stratified, well-laminated gavelinellid-bearing facies of the Tu TST in the eastern Galilee (Pl. 3f-i), deposited on the margins of a Galilean paleohigh (Fig. 6.8a,b). The Early Turonian facies-thickness configuration (Fig. 6.8b,c) is consistent with a homoclinal ramp, with relatively shallow marginal hypoxic facies in the SE and pelagic basinal facies in the NW (Fig. 6.8b). Homoclinal geometry of the Early Turonian system is supported also by subsurface seismic interpretation from offshore southern Israel showing westward-dipping ramp geometry (Gardosh 2002). The Early Turonian open-marine pelagic facies extends from the Carmel and western Galilee (Daliyya and Yirka Fms) toward the E-NE, passing into marginal facies in the SW Palmyrides. A Carmel – SW Palmyride trough was bounded from the SE by the Hamad High in Syria, and by the southern Carmel elevation in Israel. Facies of the Early Turonian Daliyya Fm in the southern Carmel region (Muhraqa section; Fig. 1.1) indicates that this region was relatively elevated in respect to the central and northern Carmel (Chapter 3; Fig. 3.9)
6.8. Discussion 6.8.1. Basin configuration: synthesis 6.8.1.1. The continuation of the ‘Sinai-Carmel hinge-belt' north of the southern Carmel Consistent distribution of facies and thicknesses in each of the five system tracts are integrated in Fig. 6.9 to show a comprehensive tectono-sedimentary configuration for the Cenomanian-Turonian of the northern Levant. E-W facies changes across the midCretaceous ‘Sinai-Carmel hinge belt’ were interrupted beginning from the central Carmel region, and north to this region N-S trending facies changes are apparent. This change in
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trend is attributed to branching of the ‘Sinai-Carmel hinge belt’ into two E-NE–trending segments, bounding a Galilean paleohigh. The southern branch traced the depositional strike of a depocenter that extended from the central-northern Carmel, via southern Galilee, into the SW Palmyrides. Toward this region, relatively shallow water ramp carbonates of the Galilee, southern Carmel and NE Palmyrides became deeper and occasionally thicker. The northern branch shifts to the E-NE starting in the NW Galilee, tracing the depositional strike of a carbonate ramp system that was relatively shallow in the SE (NW-central Galilee), and deeper toward Lebanon and western Syria in the far NNE. These depositional trends are consistent in the five Cenomanian-Turonian system tracts found here.
6.8.1.2. The elevated margins of the Carmel–SW Palmyrides basin and their trends The elevated part of the Galilee and the NE Palmyrides share similar facies and/or thickness patterns: (a) In the Ce TST-1 the Galilee-central Mt. Lebanon and the NE Palmyrides are deep subtidal with comparable thicknesses of 15-36m. (b) In the Ce RST-1 the Galilee is shallow subtidal and peritidal and the NE Palmyrides are shallow subtidal. (c) In the Ce TST-2 facies on the Galilean paleohigh changed from shallow lagoonal to pelagic outer-ramp, and low carbonate production and condensation were common. Similarly, shallow ‘littoral’ facies characterizes Ce TST-2 of the Palmyrides, where omission surfaces and hardgrounds are common. (d) In the Ce RST-2 the Galilee was mainly shallow subtidal to deep subtidal; the NE Palmyrides are shallow subtidal as well. In both regions a subaerial unconformity surface terminated Ce RST-2 at the Cenomanian-Turonian transition. (e) In the Tu TST well stratified to laminated facies were deposited in the northwestern and northeastern Galilee, as well as in the NE Palmyrides. These comparable features imply a genetic connection between the carbonate systems of the Galilee paleohigh and NE Palmyrides, and are the basis for extending the paleohigh from the Galilee to the NE Palmyrides (Fig. 6.10). The directional trend of this paleohigh is parallel to the trend of Carmel – SW Palmyride basin. A unique, very shallow water peritidal carbonate succession developed on the Galilean high during Mid-Cenomanian (Ce RST-1), Late Cenomanian (Ce RST-2) and Late Turonian (Tu RST; discussed in Chapter 2, Fig. 2.10c), suggesting that part of the Galilee was slightly more elevated compared to other parts of the Galilee–NE Palmyrides paleohigh. The Carmel – SW Palmyride depocenter/basin was bounded from the SE by the southern Carmel elevation and by the Hamad Uplift (Mouty & Al-Maleh 1983) (Fig. 6.9). The latter is a large-scale and long-lived feature trending NE-SW, sub-parallel to the Palmyrides. It was defined as a NE-trending paleohigh on the northern side of the Rutbah Uplift. Deep
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drillings in the Al-Hamad region of SE Palmyrides and Iraq indicate that Cretaceous deposits are absent in this region. Mouty et al. (2003) considered this region to have been sub-aerially exposed from the Late Triassic to the Cretaceous. Evidence from the southern Carmel show that this region was relatively elevated compared to the central and northern Carmel: (a) in the Ce RST-1 the southern Carmel region (Hotem Carmel) shallows to near sea level (peritidal cycles), but passes gradually northwards into basinal deposits of the Carmel basin (Fig. 3.5c); (b) in the Ce TST-2, bioturbated mid-ramp limestones above the Ce SB-2 in the southern Carmel region pass northwards to basinal pelagic facies in the lower Muhraqa Fm; (c) in the Tu TST, deepeningupward cycles with shallow inner-ramp beds were deposited in the southern Carmel region (Muhraqa section), but to the north these are deep subtidal cycles (Fig. 3.9). It appears that the along-strike continuation of the Hamad uplift to the southwest became shallow-subtidal to peritidal in the southern Carmel.
6.8.1.3. Trend of the elevated margins of the northern basin Data presented above show significant thickening and deepening of systems tracts from the Galilee towards the far north (Figs. 6.2, 6.4-6.8). This trend indicates that the Galilee was elevated with respect to the rapidly subsiding central Lebanon and western Syria, and that the Cenomanian depositional strike north of the Galilean paleohigh was directed to the E-NE. Scarcity of data regarding the Cenomanian of the Aleppo Plateau in the far NE (Fig. 6.1) limits the extent to which the depositional strike of the Cenomanian ramp can be extended. On the other hand, there is abundant data for the Early Turonian, and their welldefined lithologies reveal transitions from marginal deposits to deeper water ammonitebearing deposits in the northwestern Palmyrides (Figs. 6.8b). This data strongly suggests that the Cenomanian-Turonian ramp following this strike continued from the Galilee far to the NE, as shown in Figure 6.9.
6.8.1.4. Configuration of the Cenomanian-Turonian carbonate system in the northern Levant SW-NE trends of large-scale depositional features are shown in Figure 6.9. These include the southern Carmel – Hamad Uplift elevation, the Carmel – SW Palmyride depocenter/basin and the Galilee – NE Palmyride elevation. Trends are consistent throughout the Cenomanian-Turonian, suggesting long-term tectonic control on the depositional strike. Similar depositional trends are shown for the Late Albian and Early Aptian of Lebanon (Ferry at al. 2007), and are also seen in the isopach maps of Ponikarov et al. (1967) and Brew et al. (1999, 2001) for the Late Paleozoic and Mesozoic in Syria.
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The long-term consistency of the SW-NE depositional trend strongly supports the view that long-lived tectonics dictated trends of major carbonate systems in the northern Levant.
Broadly, the Cenomanian-Turonian carbonate system formed a complex ramp,
descending to the NW, and from the Galilee also toward the SW (Fig. 6.9). An open ramp configuration is supported by independent data. Brew et al. (2001), McBride et al. (1990) and Sawaf et al. (2001) show that the entire Late Cretaceous succession of the Palmyrides thins towards the Rutbah (Hamad) Uplift, but thickens toward the Palmyride axis, and remains thick beneath the Aleppo Plateau to the N-NW. Relatively minor thinning on the Aleppo High was recorded by Sawaf et al. (2001) (e.g. their fig. 6). This thickness pattern is in agreement with a Cenomanian-Turonian open-ramp configuration, but contradicts the Palmyride trough configuration recorded for the Triassic–Early Cretaceous (e.g. Brew et al. 2001). Kazmin (2005) explained the transition from confined trough to more open-margin configuration by changes in the nature of the Palmyride rift in the Late Cretaceous.
6.8.2. Paleoceanographic implications of the tectono-sedimentary framework 6.8.2.1. Dynamics of the oceanic OMZ in the Cenomanian-Turonian of the northern Levant Evidence for hypoxia is abundant in many parts of the Cenomanian-Turonian succession of northern Israel. The most common indicators are laminated or platy mudstones and fine-grained pelagic packstones (chalks, limestones and marly chalks). In the Cenomanian they constitute most of the Deir Hanna Fm and parts of the Yanuch Fm; in the Turonian they form large parts of the Daliyya and Yirka Fms. Benthic faunas are rare or absent in Cenomanian and Turonian laminites of northern Israel that contain pelagic microfossils, pyrites and glauconite grains. Lipson-Benitah et al. (1997) and Honigstein et al. (1989) described such facies in the Cenomanian and Turonian of the Carmel region (Daliyya Fm). Well-laminated or platy rocks constitute major parts of the mid-Cretaceous Sannine Fm of Lebanon (e.g. Hückel 1970; Hemleben 1977; Nader et al. 2006; Ferry et al. 2007) and parts of the Cenomanian-Turonian succession in western Syria and the Palmyrides (Ponikarov 1967; Krasheninnikov 2005). In northern Israel, laminites were deposited on slopes or basins where the oceanic OMZ impinged on the sea floor. Some deposits in Lebanon, Carmel and Galilee reflect hypoxia in the water body itself, indicated by high productivity episodes and phytoplankton blooms in the surface waters (Hemleben 1977; Honigstein et al. 1989). Transitions between different types of thin bedded rocks to their equivalent nonlaminated and bioturbated carbonate facies marks the boundary between OMZ carbonate
125
Chapter 6: Mid-Cretaceous, N. Levant
facies to mixing zone ramp carbonates. Superposition of these stages in each system tract will show long-term dynamics of the OMZ in respect to relative sea level (Fig. 6.10). In the Ce TST-1, hypoxic laminites prevailed in northern Israel (Fig. 3.4). This facies may have extended to the (restored) Bloudan region in the east, where pelagic-basinal marllimestone couplets are described. Relatively rich benthic macrofaunas described from Bloudan suggests that some parts of this TST were deposited under more aerated conditions, but platy textures and blooms of opportunistic heterohelicid foraminifera suggests that oxygen stress was still a factor. In the Palmyrides, the Ce TST-1 (lower part of the AbuZounnar Fm) is not laminated, but composed of deep subtidal marls with abundant benthic macrofauna, and no hypoxia is indicated. Therefore, the top of the OMZ of the Ce TST-1 (line 'A' in Fig. 6.10) begins impinging on the carbonate system between central Lebanon– Bloudan and the SW Palmyrides. In northern Israel, the end of the Ce TST-1 is a maximum-flooding interval that represents bottom aeration, bioturbation, and occupation of the sea floor by macrobenthos. This event marks contracting and shifting of the OMZ toward the west (line 'B' in Fig. 6.10). Somewhat later, the OMZ again expanded, marking the beginning of the Ce RST-1. In the Ce RST-1, hypoxic laminites extended from central–northern Carmel to the westernmost Galilee (e.g. Rosh Hanikra region). East and north of this facies belt, welloxygenated bioturbated mid-ramp deposits and well-stratified peritidal and shallow-subtidal sediments were deposited on the Galilee paleohigh, on the southern Carmel paleohigh, in the central Lebanon and southern Galilee subtidal accumulations, in western Syria and the Palmyrides. The top of the OMZ retreated at that time to the west, tracking a eustatic MidCenomanian sea-level fall recorded for that time (Chapter 3) (line 'C' in Fig. 6.10). In the Ce TST-2, hypoxic laminites and platy limestones were deposited north of the Galilean paleohigh (e.g. Kedesh, Manara and Sarach Valley, northernmost Galilee; Fig. 1.1) and on part of the Galilee paleohigh as shallow- inner-ramp hypoxic cycles (UC6 cycles, Fig. 2.8). Basinal laminites of Mid-Late Cenomanian age were also reported from western and central Lebanon (Mt. Sannine). They pass to the east into fine-grained bioturbated bioclastic limestones of the Al-Koroum Mb of northern Anti-Lebanon (restored position). To the SW of the Galilee paleohigh (Carmel range), basinal facies of this system tract is highly-pelagic but not laminated. The top of the OMZ of the Ce TST-2 therefore intersected the ramp between Mt. Sannine to the (restored) Anti-Lebanon, and traversed the Galilee north of Dishon and Yanuch (line 'D' in Fig. 6.10). In the Ce RST-2, slope calcarenites of western Galilee were emplaced within platy chalks or laminites of the Late Cenomanian Yanuch Fm that onlapped the Galilee paleohigh
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Chapter 6: Mid-Cretaceous, N. Levant
(e.g. top Yanuch section; Fig. 2.2, 6.3). At Dishon in the eastern Galilee, laminites are absent, and the upper Abtar Fm of the SW Palmyrides is completely bioturbated. Top-OMZ of the Ce RST-2 transected the ramp on the Galilean paleohigh between Yanuch and Dishon (line 'E' in Fig. 6.10). In the Tu TST, basinal hypoxic facies was recorded by Honigstein et al. (1989) in the Daliyya Fm of the Carmel. Similar ammonite-bearing marls are of Early Turonian marls of the western Galilee, Lebanon, Anti-Lebanon and NW Palmyrides (Mt. Billas and Dhour; Fig. 6.1). Hypoxic conditions in the Tu TST overlapping the Galilee paleohigh are indicated by laminated condensed gavelinellid-bearing mudstones at Dishon (Fig. 6.8a). Some Early Turonian successions from the Palmyride region were also described as well stratified or thinly bedded (Krasheninikov 2005), suggesting hypoxia there as well. The expanded OMZ during the Tu TST is therefore probably more widespread than that described from the Cenomanian (line 'F' in Fig. 6.10).
6.8.2.2. OMZ dynamics and Cenomanian drowning events In the Early Cenomanian (Ce TST-1; line 'A' in Fig. 6.10) the OMZ was relatively expanded. Above the OMZ, bioturbated deep subtidal deposits developed in Palmyrides, but no mid-ramp carbonate factory developed. In the Mid-Cenomanian (Ce RST-1) the OMZ retreated hundreds of kilometers toward the west (line 'C' in Fig. 6.10) following a Mid-Cenomanian sea level fall culminating in subaerial exposure (evidence in Chapter 3). A bioturbated rudist-gastropod facies formed a carbonate factory that produced sufficient carbonate to equilibrate with sea level on the Galilee high, and to fill the rapidly subsiding slopes north and south of the Galilee paleohigh (Mt. Sannine and Mt. Kedumim subtidal accumulations; Fig. 6.5). At the beginning of the upper Mid-Cenomanian (Ce TST-2) the OMZ again expanded tens of kilometers up-ramp to the SE (line 'D' in Fig. 6.10). OMZ expansion at this stage was linked to a eustatic rise described in Chapter 3. In the latest Cenomanian (highstand part of the Ce RST-2) sea level was relatively high (Chapter 3) and the OMZ was stable or shifted slightly to the west (line 'E' in Fig. 6.10). Above the OMZ, bioturbated deposits were formed in the Carmal, Galilee and Palmyrides. A relatively stable, expanded OMZ prevailed from the late Mid-Cenomanian to the end of the Cenomanian (lines ‘D’ and ‘E’ in Fig. 6.10). The stability was maintained by constant input of nutrients and organic matter to the water body. Mass occurrences of pithonellid calcispheres, phosphatic grains and bioerosion occur within laminated deposits,
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Chapter 6: Mid-Cretaceous, N. Levant
as well as within bioturbated mixing-zone deposits of that time. Such bioturbated 'mixing zone' deposits may also reflect high nutrient conditions but less oxygen depletion than the hypoxic laminites. This may be the first indication of establishment of an upwelling regime. The OMZ reached maximum expansion in the earliest Turonian (line 'E' in Fig. 6.10) and hypoxia affected deep as well as shallow water and marginal environments. Eutrophication of the water body is especially prominent in the distal open marine surface waters (e.g. Daliyya Fm - Honigstein et al. 1989) but hypoxia also affected more proximal marginal environments, evidenced by laminated and thin bedded mudstones, with gavelinellid and occasional mass occurrences of fecal pellets and some phosphatic grains. The Early Turonian interval was characterized therefore by maximum expansion of the OMZ. The highly-expanded OMZ at that time reflects climax of a continuous OMZ expansion beginning in the late Mid-Cenomanian (the path ‘D’ to ‘F’ in Fig. 6.10) and is coeval with the second Oceanic Anoxic Event of the Tethys (OAE-2). OMZ expansion began with the Mid-Cenomanian transgression (Fig. 6.10), when sea-level rise was accompanied by eutrophication of the water body. The evolution of the carbonate system of the Galilee toward the OAE-2 was discussed in Chapter 3 and is summarized in Figure 3.10.
6.8.3. Controls on systems tract geometry and facies Systems tract geometries, facies type and direction of depositional strikes were primarily governed by the large-scale, long-lived structures of the northern Levant basin (Fig. 6.10). This substrate, controlled spatial differences in tectonic subsidence and consequently, accommodation space on structural highs and lows. In most cases, active faulting on the margins of structural highs was not encountered. Notable exceptions are faulting in the northern flank of the Galilee in the late Mid-Cenomanian (Fig. 3.6a), and the fault-controlled shelf-margins of the southern flank of the Galilee paleohigh in the Late Cenomanian (Figs. 5.10, 5.11). Other factors controlling systems tract geometry and facies type/distribution are superposed on the large-scale structural pattern. These include Mid-Cenomanian and Late Cenomanian eustatic falls and rises, and variations in sediment production rate controlled by nutrient input, trophic state and degree of oxygenation of the water body. Large-scale geometries of the Mid-Cenomanian regressive system tract (Ce RST-1; Fig. 6.5) match wedging-out geometries reported from other regressive carbonate systems (e.g. Eberli & Ginsburg 1987; Hunt & Tucker 1992): Carbonate production zones of the MidCenomanian highstand were able to shed sufficient carbonate to fill rapidly-subsiding regions and to form the thickest parts of the regressive prisms. On the other hand, the distal thinning
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Chapter 6: Mid-Cretaceous, N. Levant
trend from Bloudan to the Carmel reflects sediment starvation far from the sediment production zone. The Late Cenomanian thick regressive calcarenitic prism of the Ce RST-2, extending from the Galilee to the Carmel (Fig. 6.7a) was not a result of high auto-production of carbonate. It formed in a relatively eutrophic water body, and an efficient 'carbonate factory' is absent (CC2 cycles, Chapter 2). The origin of the calcarenite grains composing this Late Cenomanian regressive prism was continuous, long-term mechanical reworking of the seafloor by shoreface waves during Late Cenomanian eustatic fall. In both Ce RST-1 and Ce RST-2, thinning toward paleohighs is caused by base levellimiting accommodation space on the paleohighs, and ultimately by subaerial truncation at sequence boundaries governed by sea level fall. The Palmyride-Carmel wedging-out prism of the Ce TST-1 (Fig. 6.4) is different. Sedimentation at this stage occurred entirely below storm wave base, and auto-production of skeletal carbonate was limited, unable to catch-up with rising sea level. Nevertheless, a macrobenthic community of bivalves and sea urchins was reported from the deep subtidal Palmyrides, with some skeletal production also in Bloudan (Mouty et al. 2003; Krasheninikov 2005). Maximal thickness of the Ce TST-1 sedimentary prism at Bloudan was achieved by carbonate supply from the Palmyrides production zone, with some auto-production and contribution of suspended pelagic fines. Thinning towards the Carmel was due to failure of auto-production in this hypoxic region, and sedimentation was exclusively of fine carbonate muds and silts from Bloudan and Palmyides in the NE, augmented by pelagic fines. The Galilee at this stage was elevated but still well beneath storm wave base, and within the OMZ, conditions not compatible with significant auto-production. Fines deposited within the OMZ in Lebanon and Syria probably also originated in the Palmyride production zones. High sea-level and eutrophication and hypoxia in the Ce TST-2 prevented the development of efficient skeletal production zones. The main sediment source at this stage was surface water, contributing fine-grained pelagic material. Nonetheless, limited production of calcarenitic skeletal grains was possible on tectonically-uplifted regions outside the OMZ, such as the NE Galilee (Dishon) and northern Anti-Lebanon (Al-Koroum Mb, see above). Relatively thin sedimentary prisms were formed, somewhat thicker in rapidly subsiding deeper zones closer to pelagic sources. These limitations controlled the geometry and facies of the relatively condensed, almost purely pelagic cover in the TST-2 of the Carmel, onlapping the Galilee paleohigh to the north (Carmal – Galilee trend of Fig. 6.6b). The much thicker accumulations of calcisiltite fines north of the Galilean high are from the subtidal skeletal production zone of the Anti-Lebanon (Al-Koroum Mb).
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Chapter 6: Mid-Cretaceous, N. Levant
The highly stressed, hypoxic water body in the Tu TST cannot account for thick accumulations of basinal carbonates (Fig. 6.8c). In fact, the peloidal/skeletal grains that form a major component of the hypoxic facies of the western Galilee (Yirka Fm) were redeposited. They formed on the Galilee high by shoreface reworking at the beginning of the Turonian transgression, and were transported in suspension or as turbidites (e.g. Pls. 1g,h, 6g) into the basin. Other anomalously thick accumulations of the Tu TST, such as at Nabi Chitt of Anti-Lebanon (100m), can be explained in a similar way.
6.9. Summary and conclusions 1. The local high-resolution sequence stratigraphic framework for northern Israel was extended to the region of the northern Levant, shown along two traverses: SW-NE Carmel-Palmyrides, and SSW-NNE Carmel-Galilee-Lebanon-W. Syria. 2. The S-N depositional strike of the mid-Cretaceous Levant margin, known as the ‘Sinai-Carmel hinge belt’, terminated just north of the southern Carmel. The sedimentary and structural configuration of all five Cenomanian-Turonian systems tracts of the Levant margin of northern Israel and beyond, consisted of two E-NE trending
sub-parallel
paleohighs
with
an
intervening
basin.
A
‘southern
Carmel─Hamad Uplift paleohigh’ in the south, inclined to the S-SW, was bounded to the north by a parallel ‘northern Carmel─SW Palmyride' basin. To the north, a parallel ‘Galilee─NE Palmyrides paleohigh’, inclined to the N-NW, descended to a basin extending well into Lebanon and western Syria. 3. The dynamics of OMZ impingement on the Cenomanian-Turonian ramp system of the northern Levant exerted control on the carbonate system. The OMZ was relatively expanded in the Early Cenomanian, and significantly contracted in the MidCenomanian, allowing the development of 'carbonate factory' zones on paleohighs. Long-term expansion of the OMZ, accompanied by eutrophication of the water body, accompanied late Mid-Cenomanian transgression. Maximum expansion of the OMZ corresponded to the Early Turonian OAE-2. 4. Large-scale system tract geometry and facies of the northern Levant were primarily controlled by the underlying structural framework. Superposed effects of sea level change, nutrient input, and consequently, position of the OMZ on the ramp, exerted significant control on the rate and locus of carbonate productivity, facies type, and on large-scale systems tract geometries.
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Chapter 7: Synthesis and summary
Chapter 7 Synthesis and summary The Cenomanian-Turonian carbonate succession of northern Israel was deconstructed into 20 facies types taken from 22 measured sections, including autochthonous and allochthonous basinal facies, a variety of outer- and mid-ramp facies types, shallow-marine inner ramp facies and features indicating subaerially-exposed environments. These facies units compose 12 types of high-order progradational or retrogradational cycles. Some of these could be integrated into eight recognizable low-order composite cycles. Some of these cycles, and particularly the deepening-upward cycles, reflect slow skeletal production, eutrophication and hypoxia. The high- and low-order cycles are the building blocks of systems tracts and sequences, assembled into a comprehensive sequence stratigraphy using bounding discontinuities and biostratigraphic control. The mid-Cretaceous carbonate system of northern Israel comprises three Cenomanian sequences and a single Turonian sequence. Major sedimentary events of the section in northern Israel can be correlated to the global oceanic system. Eustatic and palaeoenvironmental imprints are represented by earliest Cenomanian subaerial exposure; Early Cenomanian maximum flooding and oxygenation of hypoxic sea-floor; Mid-Cenomanian highstand progradation followed by forced-regression and mass-transport; Mid-Cenomanian subaerial exposure; Late Cenomanian eutrophication during sea-level rise; Late Cenomanian subaerial exposure; latest Cenomanian/Turonian eutrophication and gradual development of the OAE-2 (oceanic anoxic event). Late Cenomanian eustatic rise was locally masked by uplift and subaerial exposure. The Late Cenomanian-Early Turonian systems tract exhibits continuous deterioration of a poorly productive Late Cenomanian aggradational homoclinal ramp by increasing palaeoenvironmental stress. Although this process was overprinted by Late Cenomanian eustatic fall and local uplift of the Galilee, it culminated in the Early Turonian by eutrophication and hypoxia related to the Tethyan OAE -2. The Late Cenomanian shelf-edge in the Galilee was characterized by a dramatic transformation from homoclinal ramp to a steep-slope. This transformation involved rotational faulting along the E-W striking Yirka Fault, a Late Cenomanian fault reactivated much later. Faulting was simultaneous with Late Cenomanian eustatic fall, resulting in sediment instability and development of a composite detachment toward the south. Subsidiary block movements underlie at least three Late Cenomanian extensional basins and a syndepositional syncline. Syn-deformational fill in these basins was by calcarenites originating on the shelf north of the Yirka Fault, forming thick W-SW facing clino-successions that become thicker and steeper toward the south. Oversteepening of the calcarenitic slope
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Chapter 7: Synthesis and summary
triggered mass-movements of unconsolidated grains, redeposited in the basinal Carmel as an extensive turbiditic sheet, along with other mass-transport features, such as debrites, shear zones and internally deformed slides. The purely carbonate Galilee/Carmel slope/toeof-slope system have similarities with siliciclastic shelf-margin deltas and can be best described by a combination of two depositional models: the carbonate-apron model and the siliciclastic shelf-margin-delta model. This system corresponds therefore to a 'shelf-margin apron', with topsets, foresets and remote bottomsets. A Mid Cenomanian distally-steepened ramp system was first terminated by mechanical collapse, subjected to subaerial exposure, and finally flooded. Mechanical collapse of the ramp took place as a result of forced regression and is represented by a mass transport complex. Allochthonous slope deposits display features such as shear zone lithologies (foliations, boudins, shear-folds, ‘mylonitic’ micro-features), translational slides, slide scars, debrites and channel levee turbidites. Mechanical collapse is attributed to coincidence of two factors: (a) pressure release, fluid escape and liquefaction during the course of the Mid-Cenomanian eustatic fall; and (b) stabilization of OMZ impingement on the mid-ramp, transition from a progradational to an aggradational carbonate factory, and consequently, overloading of the ramp at the locus of distal steepening. Facies-thickness trends of the systems tracts were studied for northern Israel (Galilee and Carmel) revealing the depositional-structural framework of this region. A paleo-high extended from the southern Carmel to the E-NE, bounded to the north by a subsiding trough with the same strike, extending from the northern Carmel across the southern Galilee. An additional subparallel paleo-high in the Galilee to the north was bounded by a zone of rapid subsidence extending into Lebanon. This structural underpinning can be extended to central Lebanon, western Syria and the Palmyrides. A ‘southern Carmel─Hamad paleohigh’ in the south was bounded to the north by a parallel trough, the ‘Carmel─SW Palmyride basin’. To the north lies the ‘Galilee─NE Palmyride paleo-high’, descending again to a subsident region extending to the north into Lebanon and western Syria. This tectono-sedimentary framework of the northern Levant interrupts the S-N depositional strike of the mid-Cretaceous Levant margin, known as the ‘Sinai-Carmel hinge belt’, from southern Carmel northwards. The tectono-sedimentary framework responded to dynamics of OMZ on the Cenomanian-Turonian ramp system of the northern Levant. The OMZ was relatively expanded in the Early Cenomanian, inhibiting significant carbonate production on the ramp. It significantly contracted in the Mid-Cenomanian, allowing the development of carbonate factory zones on paleohighs. Coeval to late Mid-Cenomanian transgression, long-term expansion of the OMZ was accompanied by eutrophication of the water body. Maximum expansion of the OMZ was at the Early Turonian global OAE2 event.
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Appendix 1
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Plate 1 A – UC1 peri-tidal cycles. Base to top: deformed crinkly microbial laminites; sharp erosive contact at transition to next cycle; transgressive lag conglomerates at base of new cycle. Unit DS-1/ Ce RST-1, top Sakhnin Fm, Deir El-Assad section, W. Galilee. B – Bioturbated bioclastic-foraminiferal wackestone with gastropods (Acteonella), overlies Sakhnin Fm UC1 peritidal cycles, unit DS2/ Ce TST-2, lower Bina Fm, Deir El-Assad section, W. Galilee. C – Ce TST-1 basinal laminites (Deir Hanna Fm) in contact with Ce TST-2 bioturbated limestones (lower Yanuch Fm). Transition marked by erosive discontinuity; shear micro-textures present below surface (Pl. 9d,e). YN1 /YN2 transition, Yanuch Valley section, W. Galilee D – Photomicrograph of well-cemented grainstone, top UC3 cycle. Bioclasts, mainly oysters and rudists, some bored (arrows); composite grains (circled). YN3, Yanuch Fm, Yanuch Valley section, W. Galilee. E – Trough cross-laminated grainstones of the upper shoreface. Base Tu TST. Base subunit 4.1, top Yanuch Fm, Yanuch Valley section, W. Galilee. F – Photomicrograph of well-cemented grainstones; some bioclasts of inner-ramp origin: note benthic foraminifera Cuneolina sp. (arrow). Base subunit 4.1, Yanuch Fm, Yanuch Valley section, W. Galilee. G, H – Photomicrograph of intercalated laminae of fine-grained pelagic packstone (P) and bioclastic-peloidal grainstone (G). Tu TST, subunit YN4 of Yirka Fm, Yanuch Valley section, W. Galilee.
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Plate 2 A – Hummocky cross-stratification and thin bedding in grainstones. Tu RST, subunit BK3, Bina (Kishk) Fm, Beit-Ha'Emek Valley section, W. Galilee. B – Synsedimentary inclinations of peloidal grainstone beds, topped by hippuritid rudist shell-bed. Tu RST, subunit BK3, Bina (Kishk) Fm, Beit-Ha'Emek Valley section (Gat quarry), W. Galilee. C – Photomicrograph of packstone with broken and bored bioclasts, borings filled by chalcedonic cements. Part of 3m bed embedded in basinal chalks. Ce TST-1, subunit BZ1, Deir Hanna Fm, Betzet section, W. Galilee. D – UC6 cycles: Interbedded mudstone laminites (L) and bioturbated foraminiferal wackestones (B). Ce TST-2, subunit BZ3, Betzet section, NW Galilee. E – UC1 cycle. Base is massive-bioturbated dolomite (B) topped by crinkly microbial laminites (ML). Ce HST-2, subunit BZ3, Betzet section, NW Galilee.
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Plate 3 A, B – Part of UC10 deepening-upward cycle. Lower part (L) massive and bioclastic, separated from well-bedded and nodular middle part (M) by slightly erosive (transgressive) surface (S). Tu TST, subunit BZ4, Bina Fm, Betzet section, NW Galilee. C – Photomicrograph of mudstone laminites terminating UC10 cycle of photos ‘A,B’ above. Tu TST, subunit BZ4, Bina Fm, Betzet section, NW Galilee. D – Photomicrograph of poorly-washed grainstones, base UC9 deepening-upward cycle. Cortoids (C), oncoids (O), heavily micritized benthic foraminifera (B), microbial lumps (M) and aggregate grains (A). Ce TST-2, lower subunit DO2 of the lower Bina Fm, Dishon section, E. Galilee. E – Photomicrograph of fine-grained pelagic packstone (pelagite) from top UC9 cycle. Upper subunit DO2 of the lower Bina Fm, Dishon section, E. Galilee. F – Well-stratified laminites, Tu TST, subunit DO4, Bina Fm, Dishon section, E. Galilee. G,H,I – Photomicrographs of stratified laminites shown in photo ‘F’. Mudstones with planktonic forminifera (G), stereom plates of pelagic echinoderms (H) and gavelinellid foraminifera (I). Tu TST, subunit DO4, Bina Fm, Dishon section, E. Galilee.
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Plate 4 A – Limestone with Eoradiolites lyratus. Overlain by Deir Hanna laminites. Subunit MN1, Manara section, NE Galilee. B – Photomicrograph of packstone with flat large orbitolinids, serpulids and mollusk debris. Subunit MN2, Manara section, NE Galilee. C – Stratified planar, massive or graded calcarenite. Subunit MN1, Manara section, NE Galilee. D – Basinal laminites of Ce TST-1. Deir Hanna Fm, Kziv-East section, E. Galilee. E – Photomicrograph of poorly-washed peloidal packstone. Deir Hanna Fm, KzivEast section, E. Galilee. F – Bivalve shell-bed. Shells in life-position, highly bioeroded. Deir Hanna Fm, Kziv-East section, E. Galilee.
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Plate 5 A –Middle and upper parts of CC1 cycle. Middle part is massive bioturbated dolomitized mid-ramp limestones with unbroken rudists. Upper part is composed of stratified UC1 peritidal cycles. Sakhnin Fm, Deir El-Assad section, Beit Hakerem Valley, W. Galilee. B – Well stratified dolomites stacked in meter-scale UC1 peritidal cycles. Zikhron Fm, Hotem Carmel section, S. Carmel. C – Detail of peritidal UC1 cycle of the southern Carmel. Crinkly microbial laminites; V-shaped (desiccation) crack, brecciated storm bed. Zikhron Fm, Hotem Carmel section, S Carmel. D – Photomicrograph of remnant paleo-calcrete pavement at top of Mt. Gamal clinosuccession. Non-homogeneous fabric with circum-granular cracks, pisoids. E – Top of Pelech sequence. Irregular karst surface overlain by weathered horizon. Yanuch Fm, Hamra Valley section, W. Galilee.
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Plate 6 A – Well-stratified limestones (wackestones and packstones) of the Pelech sequence, Eastern flank of Pelech syncline. Yanuch Fm, Hamra Valley, Pelech section, W. Galilee. B – The Pelech SB, base Pelech sequence: Weathered yellowish soft marl bed paved by hardened surface with ferruginous concentrations. Yanuch Fm, Hamra Valley, W. Galilee. C – Irregular karst surface, top Mt. Gamal clinoform body. Surface paved by ferruginous concentrations and paleo-calcrete crust shown in Plate 5D. Yanuch Fm, Mt. Gamal, W. Galilee. D – Breccias of the Isfiyye section. Clasts composed of coarse-grained bioclastic grainstone. Matrix is chalky pelagic-basinal. Muhraqa Fm, Isfiyye sectin, E. Carmel. E – Irregular sub-planar surface, top of the Isfiyye breccia. Overlying are bioclastic hippuritid limestones of the upper Muhraqa Fm, Isfiyye, E. Carmel. F,G,H – Three facies units in the Yirka Fm of the Hamra Valley: lower part (F) is of interbedded wackestone-marl couplets of pelagic-basinal origin. Middle part (G) is sharp-based massive or graded grainstone beds intercalated with marl. Upper part (H) is grey or yellowish marls with ammonites. Yirka Fm, Hamra Valley section, W. Galilee.
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Plate 7 A,B – Dolomitized, yellowish, foliated indurated lithology with internal shear microfabric, emplaced along bedding planes (B) or irregular or wavy surfaces (C,D,E). 'Flow banded gouge', a synsedimentary shear-zone lithology. Sakhnin Fm, Betzet section region, NE Galilee. C,D,E – Photomicrographs of flow banded gouge. C- δ – texture: fragment of mudstone in foliated fine-grained dolomite. Shape indicates strong elongation and rotation due to sub-horizontal shear. D- σ–texture: elongated fragments of mudstone emplaced in finer-grained foliated dolomite. 1, 2 – fragments with elongation and typical wings. Fragments exhibit symmetric and asymmetric wings respectively. EC/S foliations, type of shear band cleavage. C-type shear bands are parallel to shear zone boundaries, S-type are oblique. Fabrics typical of mylonites. F – Zone of thin elongated rock slices deformed as boudins, separated by numerous dark foliations. Shear-zone, Sakhnin Fm, Namer Valley, NW Galilee. G – Dolomitized breccias. Clasts, probably of inner-ramp origin; matrix is laminar siltstone. Sakhnin Fm, Betzet, NW Galilee. H – Brecciated, folded beds, dolomitized clasts and matrix. Dashed line traces recumbent fold with axial-plane oblique to bedding. Sakhnin Fm, Betzet, NW Galilee.
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Plate 8 A – A curved shear plane; fold shown in Plate 7H and fragmented interval of photo 8B directly overlies this plane. Sakhnin Fm, Namer Valley, W. Galilee. B – Brecciated bed - note geometric matching of fragments and minor extension between clasts, embedded in dolomitized matrix. Sakhnin Fm, Namer Valley, W. Galilee. C – Mud-supported breccias; clasts are dolomitized mudstone, matrix lacks texture. Interpreted as debrite. Sheikh-Danun section, W. Galilee D – Dolomitized breccias; clasts composed of inner-ramp microbial laminites. Interpreted as debrites. Sakhnin Fm, Betzet Valley, W. Galilee.
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Plate 9 A – The northern flank of Betzet Valley near Adamit, NW Galilee. Base of slope is Late Albian Yagur Fm dolomites, overlain by basinal laminites (indurated chalks) of the Deir-Hanna Fm. These passes upwards into bioturbated limestones interbedded with basinal laminites of the Yanuch Fm. Two lenticular dolomitized bodies (Sakhnin Fm lithology) are incised in the basinal chalks of the Deir-Hanna Fm. Internal textures (Plates 9B,C) indicate they are channel fill. B – Dolomitized well-stratified, graded grainstones filling channels shown in Plate 9A, interpreted as Ta and Tb turbidite divisions sensu Bouma (1962). C – Well-laminated and ripple cross-stratified dolomitized calcisilts (Sakhnin Fm) overlying the two channels of Plate 9C. Succession consists Tb and Tc turbidite divisions sensu Bouma (1962), with truncated top, again overlain by two amalgamated Ta turbidite divisions, also separated by truncation surface. D,E – Photomicrographs showing shear features on the surface shown in Plate 1C. These winged objects of the Deir-Hanna facies show plane-parallel preferredorientation. F – Turbidites, Sakhnin Fm, Gat-Yanuch section, W Galilee.
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Plate 10 A – The Kishor clino-succession. Calcarenitic beds are 35-40m thick inclined to 25° to the SW. Clino-beds terminate against inclined downlap surface facing NE. Underlying the clinoforms is a UC3 cycle of Yanuch Fm and laminites of the DeirHanna Fm. Kishor, W. Galilee. B – Lower erosional contact of Kishor clino-succession upon well-bedded limestones of Yanuch Fm. Kishor, W. Galilee. C,D,E,F – Photomicrographs from Late Cenomanian clino-successsion. C - Coarsegrained bioclastic grainstones with mollusk bioclasts, mud peloids and microbial lumps, Kishor; D - Neomorphic bioclastic grainstone, top of Mt. Gamal clinosuccession; E - Vadose microstalactites (arrow) in dissolution vug, Mt. Gamal; FGrains coated by early-diagenetic bladed rim cement, then blocky calcite, Kishor. G – Ripple- and trough cross-laminated grainstones from top part of the Kishor clinosuccession. H – Channel, 4m wide, filled by well-sorted bioclastic sands, incised into Hamra clinoforms.
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Plate 11 A – Interfingering of Kishor clinoforms along-strike, with basinal facies of the Yanuch Fm in the west. Yirka ridge, W. Galilee. B – Syndepositional deformation in the Hamra Valley, western Galilee. To the east, well-stratified foresets of clino-beds dip 30°-35° to the SW, overlain by blocks of the Pelech synclinal sequence and relatively thin Yirka Fm. Towards the west, synsedimentary normal fault vertically offsets top Pelech beds and underlying clinoforms. Yirka Fm thickens directly above the fault plane, overlying Kishk Fm is sub-horizontal, unaffected by the fault. C – Western flank of the Pelech syncline, Hamra Valley, W. Galilee. Inclination decreases toward younger beds.
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Plate 12 A – Western termination of Pelech syncline against Hamra clino-succession. Synclinal beds (wackestones and packstones) terminate sharply against, then overlap, SW-dipping calcarenitic clino-beds. Hamra Valley, W. Galilee. B – Mt. Gamal clino-succession (Cl), bounded and internally divided by discontinuities (Ds1-Ds6; Ds1 not seen in this photo). Ds2 - the western contact of the clino-succession, with NE-tilted beds of the Yanuch Fm. Ds3 - sharp contact between NE- facing beds and W-facing beds. Probably a syndepositional reverse fault. Ds4 Syn-depositional normal fault in ‘Yanuch-type’ beds. Ds5 - internal angular discontinuity in the Mt. Gamal clinofom body, separating lower N-NE facing clinobeds from upper W-SW facing clinobeds. DS6: erosional discontinuity truncating the entire block system from above; paved by remnant paleo-calcrete. The Early Turonian Yirka Fm above Ds6 is thin (4-5m) above the clino-succession (Cl) and considerably thicker towards its flanks. C – Additional view of discontinuity surface Ds5 separating lower N-NE facing clino-beds from upper W-SW facing clinobeds. D – Close-up of Ds3. Tilted grainstones of the Yanuch Fm (Gr) overlie brecciated wackestones of the Yanuch Fm (Wc). The contact is interpreted as a Late Cenomanian reverse fault.
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Plate 13 A – Well-stratified grainstones of the Muhraqa Fm, Oren Valley, western Carmel. These grainstones are peloidal-bioclastic and mostly graded. B – Photomicrograph of peloidal-bioclastic grainstones of the well-stratified part of the Mid-Muhraqa Fm, western Carmel (Plate 13A). C – Turbidite from the well-stratified part of the Mid-Muhraqa Fm of the western Carmel (Plate 13A). Sample is graded (Ta, Tb), with rippled top (Tc). D, E – Sefunim quarry breccias, western Carmel. Imbricated blocks (Plate 13 D). Clasts (marked ‘C’) are coarse-grained bioclastic rudstones/floastones with unbroken rudists. The matrix (marked ‘M’) is coarse to fine but usually well-sorted bioclastic debris, occasionally laminated. Note flame structures (marked ‘f’) resulting from density differences between non-lithified clasts and calcarenitic matrix F – Sharp erosive surface (white and black arrows) separating bioturbated bioclastic packstones (B) from well-stratified wackestones (S). Note offset of contact by listric fault. Well-stratified upper part with recumbent-folds (grey arrow) and large-scale boudins seen in Plate 13G. Muhraqa Fm, Sefunim quarry, western Carmel. G –Discrete zone of large-scale boudins in a well-stratified part of the Muhraqa Fm (the same well-stratified zone marked 'S' in photo ‘F’). Muhraqa Fm, Sefunim quarry, western Carmel. H – The Sefunim Valley outcrop. The lower eastern part (marked ‘A’) is wellstratified, dipping 24°-26° to the SW. The middle part (marked ‘B’) is a bedded succession with inclinations antithetic to block 'A'. It is separated from the inclined beds of block 'A' by an irregular curved surface marked 'SZ'. Zone ‘C’ shows subhorizontal beds overlying both ‘A’ and ‘B’.
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את מהלך הפרוגרדציה .לקראת שיא הירידה האאוסטטית בסוף הקנומן התיכון ה OMZ -נסוג באיטיות במקביל לירידת גובה פני הים והתקרב למצב סטטי .כתוצאה מכך פסקה הפרוגרדציה והתפתחו תנאי אגרדציה והעמסת-יתר ) (overloadingעל הרמפה העליונה שהובילו לכשל של סדימנטים מידרוניים ולתנועות מאסה לכוון הים הפתוח .השלב האחרון של קריסת מערכת הפרוגרדציה של הקנומן התיכון מיוצג על ידי מישור אי-רציפות מורכב הבנוי מארבע פאזות דיאגנטיות .הפאזות המוקדמות התפתחו בהדרגה מתהליכי צמנטציה תת-ימיים לכוון של דיאגנזה באזור הוודוזי וחשיפה בעוד הפאזה הדיאגנטית המאוחרת קשורה בטרנסגרסיה ,ובחסר השקעה ומינרליזציה תת-ימיים ).(submarine omission קצה-מדף עם מדרון תלול ביותר ) (20-35°התפתח תוך זמן קצר במהלך הקנומן המאוחר כתוצאה משבירה ) (faultingשל הרמפה ההומוקלינלית של הקנומן המאוחר ע"י שבר בכוון מזרח- מערב ,שמיקומו הגיאוגרפי וכוונו חופפים בדיוק את המיקום והאוריינטציה של 'שבר ירקא' ה"צעיר" .התנועה על שבר ירקא הקנומני היתה נורמלית ,עם זריקה לכוון דרום ועם רוטציה של הבלוק התלוי לכוון צפון-מזרח ,והוא מציין קצה מדף קנומני הפונה דרומה .התנועה על השבר התרחשה תוך כדי ירידה אאוסטטית של הקנומן העליון ,שילוב גורמים שהניע תנועת בלוקים סינסדימנטרית לכוון דרום ,ככל הנראה לארכו של מישור ניתוק מורכב ) .(detachmentתנועת הבלוקים הביאה להתפתחותם של לפחות שלושה אגני מתיחה וסינקלינה סינסדימנטרית אחת ) (hangingwall synclineבגליל המערבי מדרום לשבר ירקא ,במורד מידרון הפונה לדרום – דרום- מערב .אגני אלו התמלאו ,חלקם תוך כדי התפתחותם ,בסדימנט ביוקלסטי קלק-ארניטי )חול קרבונטי( ,שמקורו באזור המדף הרדוד שמצפון לשבר ירקא הקנומני ,אזור גיתה–ינוח .קלק- ארניטים אלו הצטברו כחתכים קלינופורמיים ) (clinoformsעבים ותלולים המתעבים במהירות לכוון דרום – דרום-מערב .התללת-יתר ) (over-steepeningשל המידרונות החוליים אל מעל לזווית הקריטית לסדימנטים חוליים גרמה לתנועות מאסה של גרגירים קרבונטיים במורד המדרון. סדימנטים אלו שקעו מחדש באגן הדיסטלי של הכרמל כגוף טורבידיטי משוכב בעל השתרעות נרחבת .עדויות מהכרמל מראות כי באזור דיסטלי זה שקעו בנוסף לקלצי-טורבידיטים גסי-גרגר גם בלוקים שנעו מעל למישורי גזירה ,וכן דבריטים ).(debrites התשתית המבנית שנוצרה תוך זמן קצר בקנומן העליון ,כתוצאה מפעילותו של שבר ירקא בגליל המערבי גרמה לחלוקה של המערכת הקרבונטית לשלושה אזורים) :א( מדף פתוח בגליל מצפון לשבר ירקא ,שמציין את קצה המדף של הקנומן העליון) ,ב( מערכת מידרונית-קלקארניטית עם שיפועים תלולים שהתפתחה על תשתית מיבנית מורכבת ,החל משבר ירקא ולפחות שישה קילומטרים בכוון דרום) ,ג( מערכת אגנית עמוקה באזור הכרמל הבנויה ממשקעים אוטוכטוניים פלגיים וממשקעים אלוכטוניים .נמצא דמיון רב בין חלק מהמאפיינים של מערכת השקעה זו לאלו המאפיינים דלתאות קצה-מדף המתפתחות בסביבות פלוביאטיליות–ימיות סיליסיקלסטיות.
תת-ימיים הגובלים בשני אגנים .מגבה דרומי באזור דרום הכרמל ,גבל בצפונו באגן בכוון מזרח- מערב שנמשך ממרכז-צפון הכרמל אל הגליל התחתון .אגן זה גבל בצפונו במגבה של הגליל ,וזה גבל בצפונו באגן נוסף. הרחבה של מסגרת הרצפים של צפון ישראל לחלקים אחרים של צפון הלבנט ,אל מרכז הלבנון ,האנטי-לבנון ,מערב סוריה והפלמירידים מראה שאזור זה בנוי מיחידות טקטונו- סדימנטריות עם depositional strikeבכוון 'מזרח–צפון-מזרח' .המגבה התת-ימי של דרום הכרמל נמשך לכוון צפון-מזרח והתחבר Hamad upliftהחשוף של דרום מערב סוריה .יחידה מבנית זו )דרום הכרמל – (Hamad upliftגבלה מצפונה באגן שנמשך ממרכז-צפון הכרמל ,אל הגליל התחתון ומשם לצפון-מזרח ,אל אזור בלודן של האנטי-לבנון ואל דרום-מערב הפלמירידים )משוחזרים לאחור לפי כמות התנועה הלטרלית המאוחרת לאורך טרנספורם ים-המלח( .אגן זה )כרמל- פלמירידים( גבל בצפונו במגבה שנמשך מהגליל לכוון צפון-מזרח הפלמירידים .מגבה זה )גליל – צפ'- מז' פלמירידים( גבל בצפונו באגן נוסף ,חלק מאגן התתיס ,שנמשך מצפון הגליל וצפונה ,אל תוך לבנון ומערב סוריה .הכוון 'מזרח–צפון-מזרח' ─ 'מערב–דרום-מערב' של היחידות הטקטונו- סדימנטריות הוא דומיננטי החל מאזור דרום הכרמל וצפונה והוא מעיד על שינוי בכוון של 'קו- התמורה' ) (hinge-beltהקרטיקוני באזור צפון ישראל. חלק מהאירועים הסדימנטריים ,הקנומן-טורוניים ,בצפון ישראל ניתנים לקישור עם אירועים גלובליים .אירוע ה maximum flooding -של הקנומן התחתון ניתן לזיהוי ברחבי הפלטפורמה הערבית והינו אירוע העמקה אאוסטטי .לאירוע זה נמצאו מקבילים קרובי זמן באירופה ,כשבאזורים אלו הוא מציין קצבי סדימנטציה נמוכים יחסית ,אוורור של קרקעית דלת- חמצן והתכווצות של ה .OMZ -שני אירועים אאוסטטיים נוספים מגיעים לשיאם באירועי חשיפה והתפתחות גבולות הרצפים של הקנומן התיכון ושל הקנומן העליון. גבול הרצף של הקנומן התיכון מציין קריסה מהירה של רמפה קרבונטית "בריאה" ,דהיינו בעלת יצרנות שילדית יעילה המאפשרת מילוי הנפח הזמין לגובה פני הים .עליה אאוסטטית עוקבת החלה במהלך החלק העליון של הקנומן התיכון ונמשכה אל הקנומן העליון .עליה אאוסטטית זו התרחשה במקביל לתהליך אאוטרופיקציה של גוף המים ,ירידה הדרגתית ביצרנות השילדית והווצרות רמפה הומוקלינלית )מתונת זווית( בקנומן העליון בתנאי עקה .רמפה זו קרסה כתוצאה מירידה אאוסטטית בקנומן העליון וחשיפה ,ועליה אאוסטטית עוקבת וטביעה .אירוע טקטוני מקומי הקשור בהתרוממות ובחשיפה של הגליל הביא להתפתחות גבול רצף נוסף של הקנומן העליון ביותר. תהליך האאוטרופיקציה של גוף המים שהחל בחלק המאוחר של הקנומן התיכון התגבר במהלך הקנומן העליון והגיע לשיאו בטורון התחתון ,תקופה שהתאפיינה בחסר קיצוני בחמצן בגוף המים .שיאו הטורוני של תהליך זה מקביל בזמן לאירוע האנוקסיה האוקיאני השני של הקרטיקון ).(Oceanic Anoxic Event-2 אירוע הקריסה של הרמפה הקרבונטית של הקנומן התיכון קשור בהתמוטטות מכנית, בחשיפה אטמוספירית ,ובחסר סדימנטציה ) (omissionבמצב של עליה בגובה פני הים .ההתמוטטות המכנית של הרמפה באה לביטוי בתופעות של תנועות-מאסה במורד המידרון הכוללות גלישת בלוקים עם מישורי גזירה
פלנריים בבסיסם ) ,(translational slidesצלקות גלישה ,ברקציות
הקשורות ב ,debris flow-וקלצי-טורבידיטים .ההתמוטטות המכנית של הרמפה מיוחסת לשני גורמים הקשורים ישירות בירידה האאוסטטית של הקנומן התיכון) :א( התפתחות כשל כתוצאה מעומד הידרוסטטי נמוך ,בריחת מי נקבים וליקוויפקציה )) ;(liquefactionב( המצב הפליאו- אוקיאנוגרפי היחודי בו היתה נתונה המערכת הסדימנטרית של הקנומן התיכון בו ה OMZ-הגביל
תקציר מטרתה של עבודה זו היא לשחזר את השלבים בהתפתחות הדינמית של המערכת הסדימנטרית הקרבונטית מגיל קנומן וטורון בצפון ישראל ובצפון הפלטפורמה הערבית ,ולפענח את האופן בו הושפעה המערכת מאירועים אאוסטיים ,פליאו-אוקיאנוגרפיים גלובליים ,וטקטוניים. מתוך ניתוח סדימנטולוגי ומיקרופציאלי של חתכים עמודיים וחתכי רוחב בגליל ובכרמל נתגלו עשרים יחידות פציאליות קנומן-טורוניות הנעות בטווח שבין יחידות אגניות אוטוכטוניות או אלוכטוניות ,יחידות פציאליות של הרמפה החיצונית ,רמפה תיכונה ,רמפה פנימית ,וכן יחידות בעלות מאפיינים דיאגנטיים המעידים על חשיפה אטמוספירית .יחידות הפציאס מאורגנות אנכית במגוון צירופים 12 ,במספר ,הבונים יחידות מחזוריות בסיסיות ,בלתי ניתנות לחלוקה נוספת ,דהיינו מחזורים מסדר גבוה ) .(UC cyclesעוביים של מחזורים אלו ,ברב המיקרים ,נע מעשרות סנטימטרים עד למטרים בודדים .מחזורים אלו ניתנים לחלוקה לשתי קבוצות :קבוצה ראשונה קשורה בהצטמצמות הדרגתית של נפח ההצטברות הזמין ) ,(accommodation spaceדהיינו מחזורים המרדידים כלפי מעלה; קבוצה שניה קשורה בהתרחבות הדרגתית של נפח ההצטברות הזמין ,דהיינו מחזורים המעמיקים כלפי מעלה .חלק מהמחזורים הבסיסיים מרכיבים שמונה טיפוסי מחזורים מורכבים ,מסדר נמוך ) .(CC cyclesמחזורים אלו הם אבני הבניין של החתך הקנומן-טורוני של צפון ישראל ,והם מהווים את הבסיס לקורלציה ולניתוחים סטרטיגרפיים, סדימנטולוגיים ,פליאואקולוגיים וטקטוניים מתקדמים יותר המוצגים בעבודה זו. חלק מהמחזורים ,לרב מחזורי העמקה ,קשורים בתנאי עקה ומחסור בחמצן על קרקעית תת-ימית ו/או בגוף המים .למספר טיפוסים של מחזורי הרדדה מסדר נמוך ,משמעות פליאואקולוגית וגיאומטרית בולטת CC1 :הינו מחזור מורכב בעל דפוסי הצטברות אנכיים האפיינים לחלקים נרחבים של תצורת סחנין בגליל ולתצורת זיכרון של דרום הכרמל .במחזור מטיפוס זה ,יחידה פציאלית תיכונה שמייצגת יצרנות שילדית גבוהה ) .(carbonate factoryיחידה זו הניעה פרוגרדציה שהוגבלה ע"י גג ה .Oxygen Minimum Zone (OMZ) -יצרנות שלדית מהירה על הרמפה התיכונה לעומת חוסר יצרנות באגן העמוק הביאו להתפתחות פרופיל של distally-steepened
CC2 .rampהינו מחזור הרדדה מורכב בעל דפוסי הצטברות אנכיים האפייניים לרב תצורת ינוח של הגליל המערבי ולחלקים של תצורת בענה בגליל .ההצטברות של מחזור זה היתה בשווי משקל עם בסיס הגלים והיא מבטאת עודף נוטריינטים בגוף המים ויצרנות שילדית נמוכה על רמפה נמוכת זווית ) CC3 .(homoclinal rampהינו מחזור מורכב בעל דפוסי הצטברות אנכיים האפייניים לתצורת בענה )קישק( של הגליל המערבי .דפוס ההצטברות של המחזור מבטא מילוי נפח ההצטברות הזמין לגבה פני הים ,אך יצרנות שילדית רציפה בדומה ל CC1לא זוהתה .תכונותיו הפציאליות של המחזור מעידות על מילוי נפח האגן בעיקר כתוצאה מהסעת גרגירים קרבונטיים מהרמפה הפנימית אל ה.shoreface - זיהויים של הדפוסים המחזוריים ושל משטחי אי-רציפות )בעיקר sequence boundariesו- (maximum-flooding surfacesוהשימוש בנתונים ביוסטרטיגרפיים )בעיקר מהספרות( איפשרו הרכבת מודל קורלציית מחזורים .המודל מראה כי החתך הקנומן-טורוני בצפון ישראל בנוי משלושה רצפים ) (sequencesקנומניים ומרצף טורוני אחד ,ושלכל אחת ממערכות ההשקעה הגנטיות מסדר נמוך יותר ) (system-tractsיש משמעות מרחבית ,אופי פציאלי וגיאומטריה יחודיים. האינטגרציה של שינויי העובי עם הקונפיגורציה הפציאלית בכל אחת ממערכות ההשקעה הגנטיות ) (system-tractsמראה כי לכוון שינויי הפציאס בצפון ישראל רכיב צפון-דרום דומיננטי וכי יחידות גנטיות אלו מידקקות לכוון הגליל .המסגרת הסדימנטרית-מבנית המקומית בנויה משני מגבהים
משרד התשתיות הלאומיות
המכון הגיאולוגי
התפתחות גנטית ופליאואקולוגית של המערכת הסדימנטרית הקרבונטית מגיל קרטיקון תיכון בצפון ישראל והלבנט
רן פרנק
עבודת זו הוגשה כחיבור לקבלת תואר "דוקטור לפילוסופיה" לסנאט אוניברסיטת בן-גוריון ,באר-שבע. העבודה נעשתה בהדרכתם של: דר' בנימין בוכבינדר ,המכון הגיאולוגי ,ירושלים. פרופ' חיים בניימיני ,אוניברסיטת בן-גוריון ,באר שבע.
דוח מס' GSI/03/2010
ירושלים ,שבט תש"ע ,פברואר 2010