ORIGIN AND EVOLUTION OF SEDIMENTARY BASINS, THEIR ENERGY AND MINERAL RESOURCES WITH REFERENCE TO INTERNATIONAL ISSUES IN THE MEDITERRANEAN SEA

ORIGIN AND EVOLUTION OF SEDIMENTARY BASINS, THEIR ENERGY AND MINERAL RESOURCES WITH REFERENCE TO INTERNATIONAL ISSUES IN THE MEDITERRANEAN SEA   State of the Art 

  Dr. Mahmoud A. Radi Dar Associate Professor, Marine Geology and Geophysics National Institute of Oceanography and Fisheries 2013 

Contents Subject SUMMARY CHAPTER I: SEDIMENTARY BASINS 1- Sedimentary Basins Definition 2- Origin and Mechanisms of Basins Formation

Earth's Crust Components A- The continental shelf B- The continental slope C- The continental D- Deep-Ocean Trenches E- Plate margin (plate boundary E.1- Convergent boundaries (subduction zones) E.2- Continental crust E.3- Oceanic crust E.4- Triple junctions 3- Sedimentary Basins evolutions Divergent boundaries (ocean ridges) Convergent boundaries Transform boundaries Passive continental margins Active continental margins Foreland basin Systems Pripheral or Pro-foreland basin Retroarc or Retro-foreland basin Dynamic topography 4- Classification of Sedimentary Basins I- Tectonic Basin Classification I.1- Continental or interior sag basins I. 2- Continental graben structures and rift zones form narrow elongate basins bounded by large faults I. 3 - Failed rifts or aulacogens I. 4- Passive margin basins I. 5- Oceanic sag basins or nascent ocean basins I. 6- Basins related to subduction I. 7- Terrane-related basins I. 8- Basins related to collision Retroarc or intramontune basins Pannonian-type basins I. 9- Strike-slip and wrench basins II- Pre-, Syn-, and Post-Depositional Basins 1- Post-depositional basins 2- Syn-depositional basins 3- Pre-depositional basins 5- Basins Morphology 6- Depositional Environments 1- Continental Sediment Environments 1.1- Glacial Environments 1.2- Aeolian Environments 1.3- Rivers and Alluvial Fans 1.4- Lakes and Lacustrine Environments 2- Marine Sediment Environments 2.1- Marine deltas

P.No. 1 9 9 10 12 12 13 13 3 13 14 14 15 15 16 17 18 19 20 21 22 23 23 24 25 25 27 28 28 29 29 30 30 31 33 33 33 34 34 34 35 36 38 39 39 39 40 42 44 45

 I  

2.2- Clastic Coasts and Estuaries 2.3- The beach 2.4- Coastal plains 2.5- Beach barriers 2.6- Shallow Marine Carbonate and Evaporite Environments 2.7- Adjacent sea basins and epicontinental seas 2.8- The shallow seas and continental shelf sediments 2.9- Deep-sea basins CHAPTER II: GEOTHERMAL ENERGY IN THE SEDIMENTARY BASINS 1- Geothermal Gradient 2- Effect of the geothermal energy on hydrocarbon maturation 3- Geothermal energy utilizations 3.1- Hydrothermal Systems - Geothermal Aquifers 3.2- Hot Dry Rocks (HDR) Enhanced Geothermal Systems (EGS) 3.3- Geothermal energy in contemporary balneotherapeutics and Tourism 4- Healing and therapeutic value of geothermal waters 4.1- Therapeutic tourism 4.2- Geothermal Electricity Production around the world CHAPTER III: MINERAL RESOURCES OF THE SEDIMENTARY BASINS I- Organic Mineral Resources I.1- Oil and Natural Gas Resources 1.1- Sedimentary basins and petroleum formation in the Middle East 1.2- Petroleum prospectivity of the principal sedimentary basins on the United Kingdom Continental Shelf 1.3- Prospectivity of the sedimentary basins of Irish Sea I. 2- Coal bearing formations 2.1- Australia 2.2- India II- Inorganic Mineral Resources II.1- Volcanogenic massive sulphides (VMS) II.2- Metaliferous Oxides II.3- Metallic and Gem Minerals in Placer Deposits II.4- Evolution of a Mineralized Geothermal System, Valles Caldera, New Mexico, USA II.5- Mineral Resources of the Western Canada Sedimentary Basin II.6- Mineral Resources of the Australian Sedimentary Basins 6.1- Heavy minerals 6.2- Bauxite 6.3- Sedimentary phosphate deposits 6.4- Other Metals CHAPTER IV: MEDITERRANEAN SEA I- Mediterranean Geosynclinal Belt II- Origin and evolution of Mediterranean geosyncline III- Paleoenvironmental analysis IV- Mediterranean basins IV. 1- Tectonic Settings of Eastern Mediterranean basin IV.2- Tectonic Settings of the Western Mediterranean V- Origin and Tectonic History of Mediterranean Sub-basins V.1- The Levantine Basin V.2- Aegean Sea basin V.3- Adriatic Sea basin V.4- Ionian Sea basin V.5- The Tyrrhenian Sea

46 46 47 48 49 49 50 50 52 54 55 56 57 57 59 59 60 60 62 63 63 63 65 66 66 66 66 66 67 67 67 68 68 69 69 69 69 69 70 70 71 73 74 75 76 79 79 80 82 84 86

  II  

V.6- The Alboran Sea V.7- The Algerian Basin VI- Geothermal Potentials and Uses of the Mediterranean VI. A- Geothermal potentials A.1- Geothermal Resources in Foreland Environments A.2- Thermal Coastal Springs VI. B- Geothermal Uses B.1- Electrical production V- Mineral Resources in the Mediterranean Region V.1- Organic minerals (Oil – Natural Gas – Coal) 1. A- Oil and natural Gas resources A.1- Lavantine basin A.2- Adriatic Sea A.3- Neogene petroleum system at Alboran - Algerian Basins 1.B- Coal Bearing Formations V.2- Inorganic mineral resources CHAPTER V: EGYPT (Genius of the Place) 1- Sedimentary Basins of Egypt 1.1- Nile Delta 1.2- Eastern Desert 1.3- Red Sea Rift Valley 1.4- Western Desert 2- Geothermal Regime of Egyptian Basins 2.1- Geothermal reservoirs in the Hammam Faraun and Hammam Musa regions 3- Mineral Resources 3.1- Organic minerals A- Oil and Gas B- Coal Bearing formations 3.2- Inorganic minerals 2.1- Talc deposits 2.2- Gold, magnetite and zircon 2.3- Platinum-group minerals 2.4- Uranium isotopes 2.5- Phosphate deposits 2.6- Gypsum deposits 2.7- Limestone deposits 2.8- Shale formations References

88 90 94 94 95 95 96 96 97 97 97 97 97 98 98 99 102 102 102 104 105 106 108 110 111 111 111 112 113 113 113 114 115 116 117 118 119 120

  III  

SUMMARY

Sedimentary

basins are regions of the earth of long-term subsidence creating

accommodation space for infilling by sediments. Sedimentary basins are a characteristic feature of the Earth's crust and lithosphere and range in age from Archaean to the present day. Approximately 70% of the Earth's surface is underlain by basins of one type or another. Sedimentary basins occur in diverse geological settings usually associated with plate tectonic activity. The subsidence results from the thinning of underlying crust, sedimentary, volcanic, and tectonic loading, and changes in the thickness or density of adjacent lithosphere. Some of the world's largest basins occur within or on the stable continental and are termed Intracratonic basins. Sedimentary basins are found in a variety of tectonic settings. Subsequently, they are classified structurally in various ways, with a primary classifications distinguishing among basins formed in various plate tectonic regime (divergent, convergent, transform, intraplate), the proximity of the basin to the active plate margins, and whether oceanic, continental or transitional crust underlies the basin. Convergent boundaries create foreland basins through tectonic compression of oceanic and continental crust during lithospheric flexure. Tectonic extension at divergent boundaries where continental rifting is occurring can create a nascent ocean basin leading to either an ocean or the failure of the rift zone. In tectonic strike-slip settings, accommodation spaces occur as transpressional, transtensional or transrotational basins according to the motion of the plates along the fault zone and the local topography pull-apart basins. On oceanic crust, basins are likely to be subducted, while marginal continental basins may be partially preserved, and intracratonic basins have a high probability of preservation. As the sediments are buried, they are subjected to increasing pressure and begin the process of lithification. Such a definition excludes basins whose sedimentary infill is now incorporated in fold belts, but includes those in the stable continental interiors and flanking regions that have escaped the destructive effects of plate subduction and rifling. The largest sediment thicknesses in the geological record are also believed to have occurred at the margins of the continents. New oceanic crust forms at divergent plate boundaries, at the mid-ocean ridge system. This volcanic mountain range winds through the world ocean, forming one of the deepocean’s most conspicuous features. Along the mid-ocean ridge a distinction is made among fast-spreading, slow-spreading, and ultraslow-spreading boundaries. At fast spreading boundaries, plates move apart at 100 to 200 mm per yr. With rapid spreading, hot magma is abundant and lava flows as sheets from a central peak, giving the ridge a narrow tent-like 1   

profile (e.g., the East Pacific Rise). At slow-spreading boundaries, plates move apart at less than 55 mm per yr and the topography is broader, rougher, and features rift valleys (e.g., MidAtlantic Ridge). At ultraslow-spreading boundaries, plates move apart at less than 20 mm per yr and great slabs of mantle rock rise to the seafloor. Where two oceanic plates converge or where an oceanic plate converges with a continental plate, a subduction zone forms. In a subduction zone, the denser oceanic plate slips under the other plate and descends into the mantle. Hence, lithosphere is incorporated into the mantle in subduction zones. Subduction zones produce trenches on the deep ocean floor and are associated with shallow to deep earthquakes and violent volcanic eruptions. A transform plate boundary occurs where adjacent plates slide laterally past one another. Foreland basins are associated with compressional plate boundaries. By way of contrast, smaller basins of the fore-arc, back-arc and strike-slip type develop in response to an extensional compressional or strike-slip stress field along a plate collision zone. Sedimentary basin will not form unless there is an initial depression for the sediments to fill in. Rift-type (divergent type), compressional-type (convergent type) and strike-slip (transform type) basins are often characterized by thick sequences of continental and shallow-water sediments and therefore require substantial tectonic driving forces in order to explain them. The best known of the basin-forming mechanisms is thermal contraction of the oceanic lithosphere as it cools away from a mid-ocean ridge crest. The wide variety of sedimentary basins produce numerous types of sedimentary environments classified into; continental (fluvial, glacial, eolian), lacustrine, and deltaic environments; adjacent sea basins and epi-continental seas of varying salinity; marine depositional areas of normal salinity; transitional environments may be defined between continental and marine environments (include marine deltas, intertidal environments, coastal lagoons, estuaries, and barrier island systems). On the continents, sedimentation might be thought to begin with clastic materials shed from the flanks of mountain ranges. These alluvial fans are characterized by poorly sorted, boulder and gravel dominated, debris flow conglomerates. Fluvial (river) facies include cross-bedded and rippled river sandstones and parallel or cross-bedded floodplain mudstones (siltstones and clay shale). Lacustrine (lake) facies include sands deposited at the mouths of rivers which empty into the lake and along the shoreline as well as muddy facies on the deep lake bottom. Swamps often form in low-lying areas (for example, the area near sea level behind the shore environment) in which parallel layered, organic-rich black shales and coal form. In arid regions with little vegetation and few 2   

rivers, aeolian (wind deposited - sand dunes) environments may dominate. Aeolian sandstones frequently display large scale (1 to 3 meter) crossbed sets. Deltas form at the mouths of rivers where large volumes of siliclastics are dumped into the ocean (and lakes also). Thick accumulations of sand, silt and mud form in several subenvironments, including stream channels, flood plain, beaches, tidal flats, and sand bars. Farther offshore, at the edge of the continental shelf, is the continental slope and rise, down which gravity flows or turbidites move poorly-sorted sands and muds down into the deep ocean basins. On the deep abyssal plains, far from the influence of turbidite transported continental materials, organic mud or marine oozes are the result of a fine rain of the shells of microorganisms filtering down from near the surface. Geothermal energy indicates that part of the heat within the Earth that can or might be recovered and exploited by mankind. The geothermal or temperature gradient is the rate of increase in temperature per unit depth in the Earth due to the outflow of heat from the centre. The temperature gradient between the centre of the Earth and the outer limits of the atmosphere averages about 1°C per kilometer. Vertical deformations of the lithosphere result from the purely mechanical effects of sediment loading as well as from changes in the ambient temperature field. The temperature anomalies contribute to these deformations not only by setting up body forces but also by creating thermal in plane forces and associated bending units. Temperatures in the model are governed by the effects of vertical and horizontal thermal conduction such that the lithosphere-asthenosphere boundary is defined as a partial melt isotherm or phase change boundary which migrates vertically depending on the transient thermal state. Due to the long-term availability and the large extent of geothermal heat, geothermal energy represents an efficient renewable energy worldwide. Making geothermal heat an effective source for a sustainable supply of energy requires a quantitative reserve and resource assessment. Though immense in its nature, only a fraction of the Earth’s heat can be utilized in practice, its exploitation being limited to areas characterized by favorable hydrogeological conditions for geothermal resources to develop. A proper geothermal exploration involves different stages comprising: (1) a correct localization of potential areas to ascertain the existence of a particular geothermal field; (2)an accurate estimate of the size of the resource to determine the type of geothermal field; and (3) an appropriate identification of the main physical transport processes involved to properly identify geothermal phenomena. This requires an integrated approach involving different disciplines and methodologies including 3   

geological field measurements, laboratory-based investigations as well as mathematical modeling. It is well that the most significant portion of the world’s mineral, energy and water resources is hosted in sedimentary basins. Formation of these resources results from interactions between different coupled processes comprising groundwater flow, mechanical deformation, mass transport and heat transfer and different water–rock interaction mechanisms. Understanding the relative impact of fluid and other heat driving processes on the resulting geothermal field as well as the resulting subsurface flow dynamics is of crucial importance for geothermal energy production. For geothermal exploration it is essential to quantify the above-mentioned processes by interpretation of their characteristic thermal signatures in the subsurface. Direct-use of geothermal energy is one of the oldest, most versatile and also the most common form of utilization of geothermal energy. Now, there are 78 countries having direct utilization of geothermal energy, is a significant increase from the 72 reported in 2005, the 58 reported in 2000, and the 28 reported in 1995. The thermal energy used is 438,071 TJ/year (121,696 GWh/yr). The distribution of thermal energy used by category is approximately 49.0% for ground-source heat pumps, 24.9% for bathing and swimming (including balneology), 14.4% for space heating (of which 85% is for district heating), 5.3% for greenhouses and open ground heating, 2.7% for industrial process heating, 2.6% for aquaculture pond and raceway heating, 0.4% for agricultural drying, 0.5% for snow melting and cooling, and 0.2% for other uses. The diversified geology of various regions and stratigraphic levels within the basins have given rise to a wide variety of minerals, more than 50 different kinds other than oil, gas and coal, that have an existing or potential resource value. The minerals are divided into industrial (or nonmetallic) minerals and metallic minerals. Subsurface fluid flow plays a significant role in many geologic processes and is increasingly being studied in the scale of sedimentary basins and geologic time perspective. Many economic resources such as petroleum and mineral deposits are products of basin scale fluid flow operating over large periods of time. Volcanogenic massive sulphides are major sources of Zn, Cu, Pb, Ag and Au, and significant sources for Co, Sn, Se, Mn, Cd, In, Bi, Te, Ga and Ge. Some also contain significant amounts of As, Sb and Hg. Historically, they account for 27% of Canada's Cu production, 49% of its Zn, 20% of its Pb, 40% of its Ag and 3% of its Au. Marine placer mineral deposits are metallic and Gem minerals found on the continental shelf from the beaches to the outer shelf. 4   

Mediterranean

Sea is one of the largest mobile regions of the earth’s crust,

separating the Eastern European, Siberian, Sino-Korean, and South China platforms from the African-Arabian and Indian platforms. The Mediterranean geosynclinal belt stretches across Eurasia (Europe-Asia), from the Strait of Gibraltar in the west to the Indonesian archipelago, where it joins the Pacific geosynclinal belt. Geologic features in the present-day Mediterranean essentially result from two major processes: the tectonic displacement caused by the subduction of the African plate underneath the Eurasian plate; and the progressive closure of the Mediterranean Sea involving a series of submarine-insular sills. There are three major geomorphical settings within the Mediterranean basin; areas with stable margin characteristics, areas with unstable convergent margin characteristics, and areas with extensional margin (rifting) characteristics. The main division is that of the Western Mediterranean and Eastern Mediterranean: two basins separated by an underwater ridge that crosses the sea from Sicily to the coasts of Tunisia. The Eastern Mediterranean is one of the key regions for the understanding of fundamental tectonic processes, including continental rifting, passive margins, ophiolites, subduction, accretion, collision and post-collisional exhumation. It involves; Levantine Basin, Aegean Sea basin, Adriatic basin and Ionian Sea basin. The western Mediterranean is the younger part of the Mediterranean, being a basin formed from late Oligocene to present. The western Mediterranean consists of a series of sub-basins such as the Alboran Sea, Algerian and Tyrrhenian Sea basins. The exploitable geothermal resources in the Mediterranean are generally related not to conductive systems but to convective ones. This means that the heath is brought near the surface by fluids (mainly waters) flowing vertically from depth toward the surface, so that sufficiently high temperature may be reached by drilling at economical depth. Geothermal resources are suitable for many different types of uses and according to their temperature are commonly divided into two categories, high and low enthalpy. High enthalpy is suitable for electrical generation with conventional cycles, low enthalpy resources are employed for direct uses. The direct use o f geothermal energy is at a relatively advanced stage in European countries compared with other parts of the world. It supplies a wide range of applications and uses due to the versatility and demand for base-load heat demand plus the availability o f the resource. European countries have been pioneers in the exploitation of geothermal resources. European experience and expertise in this sector has been duplicated by other countries world-wide. Different mineral resources were considered in the Mediterranean sub-basins; 5   

petroleum and gas resources and coal bearing formations. At Turkey, The mineral matter of the basins are mainly clay minerals (illite–smectite and kaolinite), plagioclase and quartz in Bolu coal field, clay minerals (illite–smectite, smectite and illite), quartz, calcite, plagioclase and gypsum in Seben coal field, quartz, K-feldspar, plagioclase and clay minerals (kaolinite and illite), dolomite, quartz, clinoptilolite, opal and gypsum. In Western Europe, intermediate- and high sulphidation Pb–Zn–Ag–Au deposits and minor porphyry Cu–Mo mineralization in the Eastern Rhodopes are predominantly hosted by veins in shoshonitic to high-K calc-alkaline volcanic rocks of closely similar age. Base-metal-poor, high-grade gold deposits of low sulphidation character occurring in continental sedimentary rocks of synextensional basins show a close spatial and temporal relation to detachment faulting prior and during metamorphic core complex formation.

Egypt was subdivided into five major morpho-structural units; the Mediterranean Fault Zone, a belt of linear uplifts and half-grabens, the North Sinai Fold Belt “Syrian Arc”, the Suez and Red Sea Graben, and the intracratonic basins of southern Egypt. The Nile DeepSea Fan (NDSF) forms a thick sedimentary wedge covering about 100,000 km2, constructed, for the most part, since the late Miocene by influx of clastic sediments from the Nile River. The present day NDSF covers a segment of an older passive margin thought to have formed during successive rifting episodes in Jurassic and early Cretaceous times, and the total thickness of sediments on the Egyptian margin (including the post-Miocene NDSF) could exceed 9 km. The Eastern Desert of Egypt constitutes the northwestern end of the Nubian segment of the Arabian-Nubian Shield. The ophiolitic rocks of the Arabian-Nubian Shield have suprasubduction geochemical signatures. The supra-subduction signature of the ophiolites in the Eastern Desert led to further debate on whether they were formed in a back-arc setting or in a forearc setting during subduction. The Neoproterozoic ophiolites of the Eastern Desert were formed in a forearc setting based on the depleted nature of the serpentinized mantle rocks. The Red Sea occupies part of a large rift valley in the continental crust of Africa and Arabia. This break in the crust is part of a complex rift system that includes the East African Rift System. To the north, the Red Sea bifurcates into the Gulfs of Suez and Aquaba, with the Sinai Peninsula in between. The Gulf of Suez is a failed intercontinental rift that forms the NW–SE trending continuation of the Red Sea rift system and was initiated during the late Oligocene to Early Miocene by the NE–SW separation of the African and Arabian plates. It extends more than 300 km in length and can be divided into three parts: the northern portion 6   

of the Gulf dips to the SW; the central part dips to the NE; and the southern part dips to the SW. The structure of the Gulf of Suez region is governed by normal faults and tilted blocks, of which the crests represent a major hydrocarbon exploration target. The Western Desert of Egypt consists of a number of sedimentary basins that received a thick succession of Mesozoic sediments. Various geological studies have been carried out dealing with the stratigraphy, facies distribution, and tectonic framework of these sedimentary basins. The sedimentary section in the northern part of the Western Desert can be divided into three sequences based on lithology, namely: the lower clastic unit from Cambrian to pre-Cenomanian, the middle carbonates from Cenomanian to Eocene and the upper clastic unit from Oligocene to Recent. The thermal data at the eastern part of Egypt indicate that the geothermal situation of the Red Sea is more complex and broader than the Gulf of Suez. Observations near to the axial trough of the Red Sea have a mean of 470mWm2 that typical associated with an active spreading center. Whereas a mean of 116mWm2 was recorded near the coast of the northern Red Sea that is appropriate with the estimated values at the Gulf of Suez. Two heat flow provinces were distinguished: 1- the west of Nile-north of Egypt normal province with low heat flow about 46 mWm-2 and reduced heat flow of 20 mWm-2 typical of Precambrian platform tectonic setting and 2- the eastern Egypt tectonically active province with heat flow up to 80-130 mWm-2 including the Gulf of Suez and the northern Red Sea Rift System with reduced heat flow of > 30-40 mWm-2, at the transition between the two provinces. Three distinct oil and gas provinces were well known in Egypt; the Gulf of Suez, the Nile delta and Western Desert. The largest part of the production and reserves drives from prolific area of the Gulf of Suez. Egypt's hydrocarbons are accumulated in formations ranging in age from Carboniferous to Pliocene. The reservoirs are formed essentially by sands and sandstones and to a lesser extent by carbonates. Safa Formation belongs to the upper clastic unit of Middle Jurassic age is the well known coal bearing in Egypt. The thickness of the main coal seams ranges from 130 cm to 2 m and are underlain and overlain by thin black shale beds. Talc deposits occur within mafic, intermediate and felsic volcanic rocks and the talc ore bodies represent a distinct lithological unit within the volcanics. Gold deposits and occurrences located in the Nubian Shield have been known in Egypt since Predynastic times. These are stratabound deposits and non-stratabound deposits hosted in igneous and metamorphic rocks, as well as placer gold deposits. Platinum-group element (PGE) mineralization has been recently reported in podiform chromitites from the late Proterozoic 7   

Pan–African ophiolite of the Eastern Desert of Egypt. Geochemical comparison between the ore and the Nubia sandstone showed that the ore is depleted in the residual elements (Al, Ti, V, and Ni) and enriched in the mobile elements (Fe, Mn, Zn, Ba, and U) which indicates that the Bahariya iron ore is not a lateritic deposit despite the deep weathering in this area. Phosphorite deposits in Egypt, known as the Duwi Formation, are a part of the Middle East to North Africa phosphogenic province of Late Cretaceous to Paleogene age. Phosphatic grains in these deposits are classified into phosphatic mudclasts and phosphatic bioclasts. Gypsum crusts are recorded only capping the Middle Eocene carbonate rocks that are interbedded with thick gypsiferous shale beds in the north central part of Egypt. Thebes Formation forms an extensive carbonate platform on the southern margin of Tethys, outcropping along the Nile Valley and over large areas of the Western Desert of Upper Egypt. The upper Oligocene Wadi Arish Formation is composed of a carbonatedominated succession at Gebel Risan Aneiza (Sinai) with about 77-m-thick. Hagul formation represents Upper Miocene clastic/limestone sequence of about 22 m thick measured near the entrance of Wadi Hagul and Thebes Limestone, the last marine deposit before Red Sea protorifting began in Oligocene times. The Thebes Limestone formation contains several beds, 0.5 to 2 m thick near Qusier city. Carbonaceous shales have a wide distribution on the Egyptian surface and in subsurface sedimentary sequences e.g. in sediments of predominantly Carboniferous, Jurassic, Cretaceous, Paleocene and Eocene age. The carbonaceous and black shales in Egypt gained interest since five decades when the phosphorite deposits were discovered and exploited.

8   

CHAPTER I SEDIMENTARY BASINS 1- Sedimentary Basins Definition Sedimentary basins are the areas in which sediments have accumulated during a particular time period at a significantly greater rate and to a significantly greater thickness than surrounding areas and be preserved for long geological time periods, compare with physiographic basin – a depression in the surface of the land or sea-floor that may or may not be infilled with sediments. In addition, there also exist areas of long-persisting denudation, as well as regions where erosional and depositional processes more or less neutralize each other (creating what is known as non-deposition or omission) (Einsele, 1992). Sedimentary basins of one type or another today cover about 70% of the Earth's surface and contain sediment thicknesses that range from about a kilometer to several 10's km. Some basins are geologically young others have existed for 100's million years. According to Bally (1975) sedimentary basins may defined as realms of subsidence with thicknesses of sediments commonly exceeding one kilometer that are today still preserved in a more or less coherent form. A basin is born from the meeting of a sedimentary deposit and a more or less pronounced concavity in the basement. The Earth's surface exhibits a wide variety of sedimentary basins. Most of them are mobile zones by definition and are encountered at the plate boundaries. However, some of them, particularly the most extensive basins are situated on the plate themselves. These are the cratonic and intracratonic basins (Perrodon, 1983). Cratonic basins are sites of prolonged, broadly distributed but slow subsidence of the continental lithosphere, and are commonly filled with shallow water and terrestrial sedimentary rocks (Allen and Armitage, 2011), while the intracratonic basins are the basins that occur within the continental interiors away from the plate margins that undergo differential subsidence relative to the surrounding area (Busby and Azor, 2012). Allen and Allen (2006) reported that the sedimentary basins are regions of prolonged subsidence of the earth's crust. Sedimentary basins can have numerous different shapes; they may be approximately circular or, more frequently elongate depressions, troughs, or embayments, but often they may have quite irregular boundaries. Even areas without any topographic depression, such as alluvial plains, may act as sediment traps. The size of sedimentary basins is highly variable, though they are usually at least 100 km long and tens of km wide. 9   

2- Origin and Mechanism M ms of Bassins Form mation M Most sedimenntary

stuudies basiins

off

the

agreee

that

intraplaate stresses play criticcal role during

formation.

basin

The

formatioon of sedim mentary bassins by lithosphheric

streetching

reequires

transitioonal stress levels (Clooetingh et

al..,

1989).

mechannisms ultimateely

of relatedd

The

d driving

subsidence to

are

proocesses

within the relativvely rigid, cooled thermal boundary layer of thee Earth (lithospphere). Lithhosphere innvolves Earth's crust c and Asthenospheere (the upper most m part off mantle) (F Fig., 1). Asthenoosphere is the upper part p of the Earrth's mantlee lies beneaath the

 Source e: http://en.wikkipedia.org/wikki/Lithosphere 

lithosphhere and coonsists of paartially molten rock, extennding from a depth of abouut 75 km (446.5 mi) too about 200 km m (124 mi). The lithospphere is compossed of a number n of plates which are a in motioon with resppect to each othher. Sedimentarry basins therefore exist inn a backgroound enviroonment

Fig., (1) Structure of the Earrth and the lithosphere compo ositions continennt and ocean.(S Source: in the http://www.claassroomatsea.neet/general_sciennce/plate_tecton nics/te ctonics_intro.h html)

of platee motion. The T major subdivisionns of basin n settings innclude diveergent, con nvergent, transforrm, and hybbrid; 23 bassin categories occur wiithin these settings (Inngersoll, 198 88). The general characterisstics of the evolution e off many sediimentary baasins, especcially those found at rifted or o sheared continental margins, are a conven ntionally exxplained byy models off crustal subsidennce driven by the therrmal contracction of an anomalouslly hot lithosphere (Steephenson et al., 1989). Activve sedimentt depositionn, at a speciffic rate, occcurs within the overlyin ng basin

10   

and the thermal state of the sedimentary basin is fully coupled with that of the lithosphere. Sleep (1971) proposed that the tectonic subsidence of continental margin basins was caused by crustal thinning after uplift and erosion at the time of continental rifting. The apparent decrease in the widths of some basins through time can be explained by the model if sediment deposition is followed by erosion of the basin and its edges (Watts et al., 1982). The modes of rift continental margins formations can be defined based on the general geometry of the crust and mantle lithosphere during extension (Huismans and Beaumont, 2009); 1) core complex mode, where upper crustal extension is concentrated in a local area concomitant with lower crustal thinning over a wide area; 2) wide rift mode, with uniform crustal and mantle lithosphere thinning over a width greater than the lithospheric thickness; and 3) narrow rift mode, with crust and mantle lithosphere thinning over a narrow area. Narrow rifting is attributed to local weakening factors such as thermal thinning of the lithosphere, local strain weakening of the strong layers in the system, or local magmatism (Buck, 1991; Buck et al., 1999). Three explanations have been provided for wide rifts: 1) a local increase of the integrated strength resulting from replacement of crustal material by stronger mantle lithospheric material and concomitant cooling during lithosphere extension causing extension to migrate to un-thinned weaker areas of lithosphere resulting in a wide rift mode (England, 1983; Houseman and England, 1986); 2) flow of weak lower crust to areas of thinned crust in response to pressure gradients related to surface topography that result in delocalization of deformation (Buck, 1991; Buck et al., 1999); and 3) the degree of brittleductile coupling in systems containing a frictional layer bonded to a viscous layer, where the occurrence of localized or distributed, pure shear modes depends on the coupling between the layers and the lower layer viscosity (Huismans et al., 2005). Core complex modes of extension are understood to result when rapid lower crustal flow removes the crustal thickness variations (Fig., 2) required for mechanisms that would results in a wide rift zone (Buck, 1991). The thermal and mechanical response of the lithosphere to extension which occur during rifting have very great attentions. The first type involves nicking of the lithosphere, so that extension produces thinning of both upper and lower lithosphere over a given horizontal distance (Keen, 1989). The second type of geometry involves offset of the lithosphere along a low angle detachment (Wernicke, 1985). The low angle detachment or shear zone either extends through the entire lithosphere or only through the upper lithosphere. In latter case, the motion along the shear zone may be transferred to the lower lithosphere (Keen, 1989). 11   

According to Huismans and Beaumont (2009), the predicted rift modes belong to three fundamental types: 1) narrow, asymmetric rifting in which the geometry of both the upper and lower lithosphere is approximately asymmetric; 2) narrow, asymmetric, upper lithosphere rifting

concomitant

with

narrow,

symmetric,

lower

lithosphere extension; and

3)

wide,

symmetric,

crustal

rifting

concomitant

with narrow, mantle lithosphere extension. Watts et al., (1982) concluded

that,

Fig., (2) Schematic diagram illustrates the basic concepts of plate tectonic theory. Continental crust = orange; oceanic crust = green. (Source: http://www.classroomatsea.net/general_science/plate_tectonics/tectonics_intro.html

the

dominant mechanisms affecting basin subsidence are thermal contraction following heating and thinning of the lithosphere at the time of their formation, and sedimentary loading. Thermal contraction controls the overall shape of basin that is available for sedimentation, whereas sedimentary loading is the main control on the stratigraphy of a basin. They added, the flexural strength increases with time after basin initiation as the lithosphere cools. In the oceans the flexural strength of the lithosphere increases with age away from a mid-ocean ridge crest, while in the continents, the flexural strength appears to increase with age after a thermal event. The role of flexure varies as a function of both time and position during the evolution of a basin and is an important factor to consider in 'back stripping' sedimentary loads through geological time (Busby and Azor, 2012). Earth's Crust Components (Fig., 3) A- The continental shelf is a very gently sloping, submerged, extension of the continental land mass extending from the shoreline toward the deep-ocean basin. The continental shelf is relatively featureless, although some areas contain glacial deposits. The main features are long valleys running from the coastline into deeper waters. These are seaward extensions of river valleys which (along with the rest of the continental shelf) were flooded in the last Ice Age. Along some coasts the continental shelf is almost nonexistent, while at others it may extend seaward as far as 1500 km. On average, the continental shelf is 80 km wide and has a depth of 130m at the seaward edge. Although the continental shelves constitute only 7.5% of 12   

the total ocean area, they have economic importance due to their large reserves of petroleum and natural gas, as well as being home to many fishing grounds (Busby and Azor, 2012). B- The continental slope is the region of the outer edge of a continent between the generally shallow continental shelf and the deep-ocean floor. It marks the boundary between the continental crust and the oceanic crust. The

Fig., (3) Schematic diagram shows continental shelf, slope, trench and continental rise. (Source: http://ehsgeowiki.wikispaces.com/Ocean +Trench).

angle of inclination of the continental slope averages 5 degrees, although in places it may exceed 25 degrees. The continental slope is a relatively narrow feature, averaging about 20 km. in width (Busby and Azor, 2012). C- The continental rise is the gently sloping surface located at the base of a continental slope, beyond which is the abyssal plains of the deep ocean basin. The average inclination of the rise is only .3%; however, the width of the continental rise may extend for hundreds of kilometers into the deep-ocean basin. The continental rise consists of a thick accumulation of sediment that moved downslope from the continental shelf to the deep-ocean floor (Busby and Azor, 2012). D- Deep-Ocean Trenches are long and narrow, and are the deepest segment of the ocean. Some deep-ocean trenches reach depths greater than 11000m. Most are located in the Pacific Ocean. Earthquakes and volcanic activity are common in these regions. Hence, volcanic mountains often parallel trenches (Busby and Azor, 2012). E- Plate margin (plate boundary The boundary of one of the plates that form the upper layer (the lithosphere) and together cover the surface of the Earth. Earthquakes occur along rather narrow belts, and these belts mark boundaries between lithospheric plates. There are four types of seismic boundaries, distinguished by their epicenter distributions and geologic characteristics: ocean ridges, subduction zones, transform faults, and collisional zones (Condie, 2003). Seven major plates are recognized: the Eurasian, Antarctic, North American, South American, Pacific, African and Australian plates. Both plate theory and first-motion 13   

studies at plate boundaries indicate that plates are produced at ocean ridges, consumed at subduction zones, and slide past each other along transform faults. Plate boundaries are dynamic features, not only migrating about the Earth's surface, but changing from one type of boundary to another. In addition, new plate boundaries can be created in response to changes in stress regimes in the lithosphere. Also, plate boundaries disappear as two plates become part of the same plate, for instance after a continent-continent collision (Condie, 2003). Plate margins are of three main types: (a) constructive margins where newly created lithosphere is being added to plates which are moving apart at oceanic ridges; (b) convergent margins which can be either destructive margins, where one plate is carried down into the mantle, beneath the bordering plate, at a subduction zone, or a collision zone, where two island arcs or continents, or an arc and a continent, are colliding; or (c) conservative margins, where two plates are moving in opposite directions to each other along a transform fault. All three margins are seismically active, with volcanic activity at constructive and destructive margins. Some plate margins exhibit features of more than one of the three main types and are known as combined plate margins (Kimura et al., 2012). E.1- Convergent boundaries (subduction zones) a convergent plate boundary where one plate subducts beneath the other, usually because it is denser (Fig., 4). The western coast of South America is roughly coincident with a subduction zone in which a plate consisting of ocean floor is

subducting

continental America

mass

beneath of

(Gutscher,

the South

2002).

Convergent plate boundaries are defined hypocenters

by that

earthquake lie

in

Fig., (4) Subduction zone. Source: http://geology.com/nsta/

divergent-plate-boundaries.shtml

an

approximate plane and dip beneath arc systems (Condie, 2003). E.2- Continental crust is the layer of granitic, sedimentary and metamorphic rocks which form the continents and the areas of shallow seabed close to their shores, known as continental shelves. It is less dense than the material of the Earth's mantle and thus "floats" on top of it. Continental crust is also less dense than oceanic crust, though it is considerably

14   

thicker; mostly 35 to 40 km versus the average oceanic thickness of around 7-10 km. About 40% of the Earth's surface is now underlain by continental crust (Gutscher, 2002). E.3- Oceanic crust is the outermost layer of Earth’s lithosphere that is found under the oceans and formed at spreading centers on oceanic ridges. The oceanic crust is about 6 km (4 miles) thick. It is composed of several layers, not including the overlying sediment. The topmost layer, about 500 meters (1,650 feet) thick, includes lavas made of basalt. Oceanic crust differs from continental crust in several ways: it is thinner, denser, younger, of different chemical composition, and formed above the subduction zones (Gutscher, 2002). E.4- Triple junctions are points where three plates meet. Such junctions are a necessary consequence of rigid plates on a sphere, since this is the common way a plate boundary can end. There are sixteen possible combinations of ridge, trench, and transform-fault triple junctions, of which only six are common. Triple junctions are classified as stable or unstable, depending on whether they preserve their geometry as they evolve. It is important to understand evolutionary changes in triple junctions, because changes in their configuration can produce changes that superficially resemble changes in plate motions. Triple junction evolution is controlled by the lengths of transform faults, spreading velocities, and the availability of magma (Condie, 2003).

15   

3- Sedimentary Basins evolutions: Sedimentary basins are dominated during their evolution by epeirogenic or vertical movements of the Earth's crust. Epeirogenic setting is the formation and submergence of continents by broad relatively slow displacements of the earth's crust. Epeirogenic movement can be permanent or transient. Transient uplift can occur over a thermal anomaly due to convicting anomalously hot mantle, and disappears when convection wanes. Permanent uplift can occur when igneous material is injected into the crust, and circular or elliptical structural uplift (that is, without folding) over a large radius (tens to

Fig., (5) Triple junction. Source: http://www.geosophia.co.in/more_article %202.php

thousands of km) is one characteristic of a mantle plume. Although an individual basin may change its tectonic setting during its evolution, most basins can be classified as occurring in either a rifted or an orogenic setting (Sloss and Speed, I974; Dickinson and Yarborough 1976; Bally and Snelson I980). Rifted basins are associated with divergent plate boundaries where extension is dominant, for example, the U.S. Atlantic margin basins (Baltimore Canyon Trough, South Carolina Trough), which are located on transitional crust between ocean and continent, and possibly the North Sea Basin, which occurs on pre-Mesozoic continental crust (). Orogenic basins (Bally and Snelson 1980), on the other hand, are associated with convergent plate boundaries where compression is dominant, for example fore-arc basins (Cook Inlet, Alaska) and foreland basins (Appalachian, Alberta and Ganges) (Watts et al., 1982). Orogenic setting is the variety of processes that occur during mountain-building, including: distinctive patterns of deposition, deformation, metamorphism, intrusions, volcanic activity, oceanic trenches and seismic activity. Lithospheric flexure (also called regional isostasy) is the process by which the lithosphere bends under the action of forces as the weight of a growing orogen or the changes in ice thickness related to (de)glaciations. Sedimentary basins subside primarily owing to the following processes: attenuation of crust as a result of stretching and erosion, contraction of lithosphere during cooling, depression of lithosphere by sedimentary and tectonic loads and the vertical crustal movements. Phase changes occur beneath the lithosphere in the upper mantle, such as

16   

localized cooling followed by contraction which will create a superficial depression (later on it will be filled up by sediments). Conversely, lithosphere may locally heat up and expand causing the continental crust to dome. Erosion follows and creates a hollow for sediments to fill in (Einsele, 1992). The first two processes dominate in most divergent settings, whereas the third process dominates in most convergent settings. Intraplate, transform, and hybrid settings experience complex combinations of processes. Several basin types have low preservation potential, as predicted by their susceptibilities to erosion and uplift during orogeny and as confirmed by their scarcity in the very ancient record. The relative tectonic motion produces deformation concentrated along plate boundaries which are of three basic types (Einsele, 1992): • Divergent boundaries. • Convergent boundaries. • Transform boundaries. Divergent boundaries (ocean ridges) Ocean ridges are accretionary plate boundaries where new lithosphere

is

formed

Fig., (6a) Divergent tectonic motion Source: http://geology. com/ nsta/divergent-plate-boundaries.shtml

from

upwelling mantle as the plates on both sides of ridges grow in area and move away from the axis of the ridge, new ocean ridges

formed

beneath

supercontinents, and thus as new

oceanic

produced

at

lithosphere a

ridge

is the

supercontinent splits and moves apart on each of the ridge flanks. Divergent boundaries occur when plate are rifted apart and

begin

to

move

Fig., (6b) Divergent tectonic motion. (Source: http://www. indiana.edu/~geol116/week7/week7.htm)

apart,

creating large expanses of oceanic crust. Crust is created in this type of boundaries form where new oceanic lithosphere is formed and plates diverge. These occur at the mid-ocean 17   

ridges (Fig., 6a,b). At fast spreading boundaries, plates move apart at 100 to 200 mm/yr. With rapid spreading, hot magma is abundant and lava flows as sheets from a central peak, giving the ridge a narrow tent-like profile (e.g., the East Pacific Rise). At slow-spreading boundaries, plates move apart at less than 55 mm/yr and the topography is broader, rougher, and features rift valleys (e.g., Mid-Atlantic Ridge). At ultraslow-spreading boundaries, plates move apart at less than 20 mm/yr and great slabs of mantle rock rise to the seafloor. The median valley of ocean ridges varies in geological character due to the changing importance of tectonic extension and volcanism. In the northern part of the Mid-Atlantic ridge, stretching and thinning of the crust dominate in one section, while volcanism dominates in another. Where tectonic thinning is important, faulting has exposed gabbros and serpentinites from deeper crustal levels (Condie, 2003). The axial topography of fast- and slow-spreading ridges varies considerably. A deep axial valley with flanking mountains characterizes slow-spreading ridges, while relatively low relief, and in some instances a topographic high, characterize fastspreading ridges. Model studies sag as oceanic lithosphere thickens with distance from a ridge axis, horizontal extensional stresses can produce the axial topography found on slowspreading

ridges.

In

fast-

spreading ridges, however, the calculated stresses are too small to result in appreciable relief. The axis of ocean ridges is not continuous, but may be offset by several tens to hundreds of kilometers by transform faults Source: http://geology.com/nsta/convergent-plate-boundaries.shtml

(Condie, 2003). . Convergent boundaries form where plates converge. One plate

is

beneath

usually the

convergent

subducted

other plate

at

a

boundary.

Convergent boundaries may be of different types, depending on the

types

of

lithosphere

involved. These results in a wide

Fig., (7) Convergent Tectonic motion between continent and ocean plate. Source: http://rainforestgirl.edu.glogster.com/plate-boundaries/

diversity of basin types formed at convergent boundaries. There are two types of convergent 18   

boundaries. Collision- Two plates with continental crust collide and create mountains. Subduction- Two plates with oceanic crust or one plate with continental crust and one with oceanic crust move together. The oceanic (older, denser, colder) subducts, or sinks under the other plate. This creates coastal mountains (Fig., 7). At convergent boundaries oceanic lithosphere is always destroyed by descending into a subduction zone. This is because oceanic rock is heavy, compared to the continents, and sinks easily. Because oceanic lithosphere is created and destroyed so easily ocean basins are young; the oldest we have is only about 200 million years old. Continents, on the other hand, composed of light weight rock never subducts. Thus, continental rock once formed is more or less permanent;

the

oldest

continental

fragment is 3.9 billion years old, virtually as old as the earth itself (Einsele, 1992). Transform

are

boundaries

the

boundaries between two plates that are sliding against each other horizontally. Neither plate is destroyed in this process at the boundary. Another name for this boundary type is a transform fault. Most transform faults occur on the ocean floor but they are a few major faults that are located on continental

plates.

These

can

be

complex and are associated with a variety of basin types (Fig. 8). These offsets may have developed at the time spreading

began

inhomogeneous lithosphere.

and

fracturing

Transform

reflect of

the

faults,

like

ocean ridges, are characterized by shallow earthquakes (< 50 km deep).

Fig., (8) Transform motion, the plates sliding against each other. Sources: 1- Oceanic Transform Boundaryhttp://www.kidsgeo.com/images/transform-boundary.jpg. 2Continental Transform Boundary, ttp://www. visionlearning.com/library/module_viewer.php?mid=66 3- http://pubs.usgs.gov/fs/1999/fs110-99/

At transform boundaries two plates just slide past one another horizontally, and quietly 19   

compared to convergent and divergent plate boundaries. Most of these are found in the ocean basins, but the San Andreas Fault in California and Mexico is an example coming on land (Einsele, 1992). Many basins form at the continental margins. Continental margins are described either as passive (Fig., 9a), where the boundary between oceanic and continental lithosphere is not a plate boundary (as around most of the present day Atlantic Ocean), or active where the ocean-continent boundary is a plate boundary associated with subduction (as around most of the present day Pacific Ocean). Passive continental margins occur away from plate tectonic boundaries along the edges of opening ocean basins like the Atlantic basin. These margins are characterized by minimal tectonic and

igneous

activity

(Condie,

2003), and generally consist of a gently sloping shelf, a slope and a rise. Passive margins are found

Fig., (9a) Active and passive continental margins. Source: http://sio.ucsd.edu/png/science/

along most of the coastal areas that surround the Atlantic Ocean, such as the east coasts of North and South America, Western Europe and Africa. Passive margins are not associated with plate boundaries and therefore experience minimal volcanism and few earthquakes. Passive continental margins (Fig., 9b,c,d) are comprised of three main

Fig., (9b) passive continental margins. Source: http://www.earth.northwestern.edu/people/seth/202/lectures/Platetect /Continentalevl/passive.htm

features: the continental shelf, the continental slope and the continental rise. Depositional systems in cratonic and passive margin basins vary depending on the relative roles of fluvial, aeolian, deltaic, wave, storm and tidal processes. Spatial and temporal distribution of sediments is controlled by regional uplift, the amount of continent covered by shallow seas, and climate (Klein, 1982).

20   

Active

continental

are

margins

usually narrow and consist of highly deformed

sediments.

They

occur

where oceanic lithosphere is being subducted beneath the margin of a continent.

In

active

continental

margins, the continental slope and the continental wall of the trench are essentially the same feature. An active continental margin is found where either a subduction zone or a transform fault coincides with continent-ocean interface. Examples are the Andean and Japan continental-margin arc systems and the San Andreas transform fault in California (Condie, 2003). The oceanic

Fig., (9c) Active Continental margins. Source: http://www. geology.ohio-state.edu/~vonfrese/gs100/lect21/index.html

lithosphere is being subducted beneath the edge of a continent. The sediments from the ocean floor and pieces of the oceanic crust are scraped from the descending oceanic plate and plastered against the edge of the overriding continent. This area of highly deformed sediment is called an accretionary wedge.

Some

active

continental

Fig. (9d) Active continental margin. Source: http://elearning.stkc.go.th/lms/html/earth_science/LOcanada6/60 4/7_en.htm.

margins do not have an accretionary wedge, indicating the ocean sediments are being carried directly into the mantle. Here the continental margin is very narrow and the trench may lie only 50km offshore (Einsele, 1992).

Fig. (9e) Isostatic changes in the crustal/lithosphere thickness. Source: http://www. earth.northwestern .edu/people/ seth/ 202/ lectures/Platetect/Continentalevl/Image130.gif

The formative mechanisms of sedimentary basins fall into a small number of categories, although all mechanisms may operate during the evolution of a basin as documented in Allen and Allen (2006): 21   

•

Isostatic consequence of changes in crustal/lithosheric thickness, such as caused mechanically by lithospheric stretching, or purely thermally, as in cooling and subsidence of oceanic lithosphere as it moves away from oceanic spreading centers;

•

Loading (and unloading) of the lithosphere causes a deflection or flexural deformation and therefore subsidence (and uplifting) as in foreland basins (Fig., 9e).

Foreland basin System A

foreland

basin

system

is

defined as: (a) an elongate region of potential sediment accommodation that forms on continental crust between a contractional orogenic belt and the adjacent craton, mainly in response to geodynamic

processes

related

to

subduction and the resulting peripheral or retroarc fold-thrust belt (Fig., 10a); (b) it consists of four discrete depozones, referred to as the wedge-top, foredeep, forebulge and back-bulge depozones –

Fig., (10a) Foreland basin development. Source: http://ww w.searchanddiscovery.com/documents/2009/50203smith/ images/fig27.htm .

which of these depozones a sediment particle occupies depends on its location at the time of deposition, rather than its ultimate geometric relationship with the thrust belt; (c) the longitudinal dimension of the foreland basin system is roughly equal to the length of the fold-thrust belt, and does not include sediment that spills into remnant ocean basins or continental rifts (impactogens) (DeCelles and Giles 1996). The generally accepted definition

Fig. , (10b) Schematic map view of a ‘typical’ foreland basin, bounded longitudinally by a pair of marginal ocean basins. After DeCelles and Giles (1996)

of a foreland basin attributes sediment accommodation solely to flexural subsidence driven by the topographic load of the thrust belt and sediment loads in the foreland basin. Equally or more important in some foreland basin 22   

systems are the effects of subduction loads (in peripheral systems) and far-field subsidence in response to viscous coupling between subducted slabs and mantle–wedge material beneath the outboard part of the overlying continent (in retroarc systems). Pripheral or Pro-foreland basin occurs on the plate (Fig., 11) that is subducted or underthrust during plate collision (i.e. the outer arc of the orogen). Pro-foreland basins

are

characterised

by:

(1)

Accelerating tectonic subsidence driven primarily by the translation of the basin fill towards

the

mountain

belt

at

 

the

convergence rate. (2) Stratigraphic onlap onto the cratonic margin at a rate at least equal to the plate convergence rate. (3) A basin infill that records the most recent development of the mountain belt with a

 

preserved interval determined by the width of the basin divided by the convergence rate (Naylor and Sinclair, 2008). Retroarc or Retro-foreland basin occurs on the plate that overrides during plate convergence or collision (i.e. situated behind the magmatic arc (Fig., 11) that is linked with the subduction of oceanic lithosphere). Retro -foreland basins are relatively stable, are not translated into the mountain achieved,

belt and

once

steady-

are

state

is

consequently

Fig., (11) Pripheral (Pro-foeland) and retroarc (RetroForland) loading and unloading subduction (after, De Celles and Giles, 1996; (Naylor and Sinclair, 2008).

characterised by: (1) A constant tectonic subsidence rate during growth of the thrust wedge, with zero tectonic subsidence during the steady- state phase (i.e. ongoing accretion-erosion, but constant load). (2) Relatively little stratigraphic onlap driven only by the growth of the

23   

retro-wedge. (3) A basin fill that records the entire growth phase of the mountain belt, but only a condensed representation of steady state conditions (Naylor and Sinclair, 2008). Dynamic topography A

substantial

portion

of

Earth's topography is known to be

caused

by

the

viscous

coupling of mantle flow to the lithosphere

but

the

contributions

of

asthenospheric

flow

deeper

relative shallow

flow

versus remains

controversial (Fig., 12). The motions of continents relative to large-scale patterns of mantle convection can contribute to the creation

and

destruction

of

sediment accommodation space due

to

transient,

displacement

of

dynamic

the

surface

topography, usually referred to as

dynamic

Fig., (12) Subsidence/uplift dynamic topography structure due to viscous flow of mantle (After Smith et al., 2009). 

topography

(Lithgow-Bertelloni and Gurnis, 1997; Gurnis et al., 1998). Previously, the anomalous depth has been attributed to asthenospheric flow and the coupling of the shallow mantle. It is well established that mantle convection imparts an influence on surface plate dynamics and the surface expression of such deep earth processes is manifested in large-scale and non-isostatic vertical motions also termed “dynamic topography” (Shephard et al., 2012). Large-scale, mantle-driven dynamic topography can be approximated by the time-dependent vertical shifts and tilts of a plane, computed from the displacement needed to reconcile the interpreted pattern of marine incursion with a predicted topography in the presence of global sea level variations (DiCaprio et al., 2009, Heine et al., 2010).

24   

4- Classification of Sedimentary Basins   Many sedimentologists therefore prefer a classification scheme based mainly on criteria which can be recognized in the field, i.e., the facies concept and the definition of the depositional environment (fluvial sediments, shelf deposits etc.). A further approach is the subdivision of sediments into important lithologic groups, such as siliciclastic sediments of various granulometries and composition, carbonate rocks, evaporites, etc. Having established the facies, succession, and geometries of such lithologic groups, one can proceed to define the tectonic nature of the basin investigated (Einsele, 1992). Sedimentary basins have been classified principally in terms of the type of lithospheric substratum (continental, oceanic, transitional), the position with respect to a plate boundary (interplate, intraplate) and the type of plate margin (divergent, convergent, transform) closest to the basin (Gutscher, 2002). Allen and Allen (2006) classified the sedimentary basins principally in terms of the type of lithospheric substratum (i.e., continental, oceanic, transitional), their positions with respect to the plate boundary (intracratonic, plate margin), and type of plate motion nearest to the basin (divergent, convergent, transform). According to Einsele (1992), there are three essential types of sedimentary basins can be recognized: (1) Active sedimentary basins that are still accumulating sediments. (2) Inactive, but little deformed sedimentary basins showing more or less their original shape and sedimentary fill. (3) Strongly deformed and incomplete former sedimentary basins, where the original fill has been partly lost to erosion, for example in a mountain belt. The regional deposition of sediments, non-deposition, or denudation of older rocks are controlled mainly by tectonic movements. Most of the recent attempts to classify sedimentary basins have been based on global and regional tectonic concepts (Einsele, 1992). Subsequently, the characteristics of sediments filling a basin of a certain tectonic type are predominantly controlled by other factors and can be extremely variable. In addition to tectonic movements in the basinal area itself, sedimentary processes and facies are controlled by the paleogeography of the regions around the basin (peri-basin morphology and climate, rock types and tectonic activity in the source area), the depositional environment, the evolution of sediment-producing organisms, etc. I- Tectonic Basin Classification Basin-generating tectonics is the most important prerequisite for the accumulation of sediments. Therefore, a tectonic basin classification system is of the most important

25   

sedimentation basins in the Earth's crust. Such a basin classification must be in accordance with the modern concept of global plate tectonics and hence will differ from older classifications and terminology (Einsele 1992). In recent years, several authors have summarized the interaction between plate tectonics and sedimentation processes and proposed the basin classification systems and basically identical them (Dickinson in Dickinson and Einsele (1992) described the classification by Kingston et al. (1983) and Mitchell and Reading (1986), but with some minor modifications. According to them, the different types of sedimentary basins can be grouped into seven categories, which in turn may be subdivided into two to four special basin types as in Table (1): Table (1) Tectonic basin classification (After Kingston et al. 1983; Mitchell and Reading 1986): No

Basin category

Special basin type

1

Continenetal or interior sag basins

2

Continenetal or interior fracture basins

Epicontinental basins, infracratonic basins. Graben structures, rift valley, rift zones, aulacogens

3

Basins on passitve continenetal margins, margin sag basins

4

Ocean sag basins

5

Basins subduction

6

7

related

Basins related collision

Strike-slip/wrench basins

Underlying crust Continenetal

Style of tectonics Divergence

Continenetal

Divergence

Tensional-rifted basins, tension sheared basins, sunk margin basins

Transitional

Divergence + shear

Nascent ocean basin (growing oceanic basin) Deep sea trenches,

Oceanic

Divergence

Oceanic

Convergence

Forearc basins, backarc basins, interarc basins Remnant basins

Transitional oceanic

Dominantly divergence

Oceanic

Convergence

Forland basins (periheral), retroarc basins (intramontane), brocken foreland basins,

Continenetal

Crustal flexuring, local convergace or transform motions

Terranerelated basins Pull-apart basins (transtensional) and transpressional basins

Oceanic Continenetal and/or Oceanic

Basin Characteristic Large areas, subsidence

slow

Relatively narrow basins, foult bounded, rapid subsidence during early rifting. Asymmetric basins partly outobuidling of sediments, moderate to low subsidence during later stages Large asymmetric, slow subsidence Partly asymmetric, greatly varying depth and subsidence Activated subsidence due to rapid sedimentary loading. Asymmetric basins, tend to increasing subsidence, uplift and subsidence Similar to backarc basins

Transform motions ± divergance or convergance

Relatively small, elongate, rapid subsidence

26   

I.1- Continental or interior sag basins According

to

Einsele

(1992), basins on continental crust are

commonly

divergent

plate

generated

by

motions

and

resulting extensional structures and thermal effects. In the case of large interior sag basins, however, major fault

systems

forming

the

boundaries of the depositional area or a central rift zone may be absent.

Subsidence

predominantly

in

occurs

response

Fig., (13). Conceptual diagram of basins in the rift–drift suite, associated with continental extension, modified from Allen and Allen (2006). Cratonic basins are viewed as basins whose primary mechanism for subsidence is low strain rate stretching.

to

moderate crustal thinning or to a slightly higher density of the underlying crust in comparison to neighboring areas (Fig., 13). In addition, slow thermal decay after a heating event and sedimentary loading can promote and maintain further subsidence for a long time. Alternatively,

it

suggested

that

was

recently long-term

subsidence of intracratonic basins may be related to a decrease of the mantle heat flow above a "cold spot", i.e., to abnormal cooling (Ziegler 1989). In general, rates of

Fig., (14) Models of strain geometry; a- pure shear geometery, bsimple shear and c- hybrid model of simple shear.

subsidence are low in this geodynamic setting. Intracontinental sags, rifts, failed rifts and passive continental margins fall within an evolutionary suite of basins unified by the process of lithospheric stretching. Rifts are areas of crustal thinning, demonstrated by the shallow depth of Moho, high surface heat flows, volcanic activity, seismic activity with predominantly extensional focal mechanism solutions, negative Bouguer gravity anomalies and commonly elevated rift margin topography (Allen 27   

and Allen 2006). The nature of the fault system and associated sedimentary basins within extending continental lithosphere depends on the initial crustal structure and geotherm, strain rate and total amount of strain (Fig., 13). Discrete, localized continental rifts appear to form on normal thickness crust and extend slowly over long period of time. At higher strain rates, localized rifts may evolve into passive margins. Passive continental margins are in general seismically inactive, and tectonics are dominated by gravity driven collapse, halokinesis and growth faulting. Passive continental margins can be divided into two types: i) volcanic margins are characterized by extensive extrusive basalts and igneous underplating and significant surface uplift at the time of breakup and ii) nonvolcanic margins lack evidence for strong thermal activity, and consist of extensive sediment traps overlying a strongly rifted basement Allen and Allen 2006). I. 2- Continental graben structures and rift zones form narrow elongate basins bounded by large faults Their cross sections may be symmetric or asymmetric (e.g., halfgrabens) (Fig., 15). If the underlying mantle is relatively hot, the lithosphere may expand and show updoming prior to or

Fig., (15) Continental graben structures, rift zones and Failed rifts and aulacogens. (After Dickinson and Yarborough 1976; Kingston et al. 1983; Mitchell and Reading 1986).

during the incipient phase of rifting substantial thinning of the crust by attenuation, which is often accompanied by the upstreaming of basaltic magma, thus forming transitional crust, causes rapid subsidence in the rift zone. Subsequent thermal contraction due to cooling and high sedimentary loading enable continuing subsidence and therefore the deposition of thick sedimentary infillings (Einsele 1992). I. 3 - Failed rifts or aulacogens If divergent plate motion comes to an end before the moving blocks are separated by accretion of new oceanic crust, the rift zone is referred to as "failed". A certain type of such failed rifts is an aulacogen. Aulacogens represent the failed arm of a triple junction of a rift zone (Fig., 15), where two arms continue their development to form an oceanic basin. Aulacogen floors consist of oceanic or transitional crust and allow the deposition of thick sedimentary sequences over relatively long time periods. Basins similar to aulacogens may also be initiated during the closure of an ocean and during orogenies (Einsele 1992).

28   

I. 4- Passive margin basins. The initial stage of a true oceanic basin setting (or a proto-oceanic rift system) is established when two divergent continents separate and new oceanic crust forms in the intervening space. This does not necessarily mean that such a basin type fills with oceanic sediments, but it does imply that the central basin floor lies at least 2 to 3.km below sea level (Fig., 16). When such a basin widens

due

to

continued

divergent plate motions and accretion

of

oceanic

crust

(drifting stage), its infilling with sediments lags more and more behind ocean spreading. Consequently, the sediments are deposited predominantly at

Fig., (16) Passive margin basins and the oceanic basin plain and faultbounded basin. (After Dickinson and Yarborough 1976; Kingston et al. 1983; Mitchell and Reading 1986).

the two continental margins of the growing ocean basin. The marginal "basins" developing on top of thinned continental crust are commonly not bordered by morphological highs and represent asymmetric depositional areas. Their underlying crust increasingly thins seaward; hence subsidence tends to become greater and faster in this direction. Here, sediments commonly build up in the form of a prism. Some of these marginal basins may be affected and bordered by transform motions (tension-sheared basins). In a sediment starved environment, subsided transitional crust can create deep plateaus (sunk basins). In general, subsidence of these marginal basins tends to decrease with passing time, unless it is reactivated by heavy sediment loads (Einsele 1992). I. 5- Oceanic sag basins or nascent ocean basins These types of basins occupy the area between a mid-oceanic ridge, including its rise, and the outer edge of the transitional crust along a passive continental margin. accumulate

They

commonly

deep-sea

fan

or

Fig., (17) Oceanic sag basins or nascent ocean basins (After Dickinson and Yarborough 1976; Kingston et al. 1983; Mitchell and Reading 1986)

basin plain sediments. Due to the advanced cooling of the aging oceanic crust, subsidence is usually low, unless it is activated by thick sedimentary loading near the continental margin. Fault-bounded basins of 29   

limited extent are common in conjunction with the growth of mid-oceanic ridges (Fig., 17) (Einsele 1992). I. 6- Basins related to subduction. Another group of basins is dominated by convergent plate motions and orogenic deformation. Basins related to

the

development

of

subduction complexes along island

arcs

or

active

continental margins include deep-sea trenches, forearc

Fig., (18) Basins related to subduction (After Allen and Allen, 2006).

basins, backarc basins (Fig., 18), and smaller slope basins and intra-arc basins. Deep-sea trench floors are composed of descending oceanic crust. Therefore, some of them represent the deepest elongate basins present on the globe. In areas of very high sediment influx from the neighboring continent, however, they are for the most part filled up and morphologically resemble a continental rise. Deep-sea trenches commonly do not subside as do many other basin types. In fact, they tend to maintain their depth which is controlled mainly by the subduction mechanism, as well as by the volume and geometry of the accretionary sediment wedge on their landward side. Forearc basins occur between the trench slope break of the accretionary wedge and the magmatic front of the arc. The substratum beneath the center of such basins usually consists of transitional or trapped oceanic crust older than the magmatic arc and the accretionary subduction complex. Rates of subsidence and sedimentation tend to vary, but may frequently be high. Subsequent deformation of the sedimentary fill is not as intensive as in the accretionary wedge (Einsele 1992). Backarc or interarc basins form by rifting and ocean spreading either landward of an island arc, or between two island arcs which originate from the splitting apart of an older arc system (Allen and Allen, 2006). The evolution of these basins resembles that of normal ocean basins between divergent plate motions. Their sedimentary fill frequently reflects magmatic activity in the arc region. I. 7- Terrane-related basins They are situated between micro-continents consisting at least in part of continental crust (Nur and Ben-Avraham 1983) and larger continental blocks. The substratum of these basins is

30   

usually oceanic crust. They may be bordered by a subduction zone and thus be associated with either basins related to subduction or collision. I. 8- Basins related to collision. Partial collision of continents with irregular shapes and boundaries which do not fit each other leads to zones of crustal overthrusting

and,

along

strike, to areas where one or more

oceanic

basins

of

reduced size still persist. These remnant basins (Fig. 19) tend to collect large volumes of sediment from nearby rising areas and to undergo

Fig. (19) Tectonic basin classification, Subduction and collision-related basins (remnant basin) (After Allen and Allen, 2006).

substantial

synsedimentary deformation (convergence,

also

often

accompanied by strike-slip motions). Foreland basins, and peripheral basins in front of a foldthrust belt, are formed by depressing and flexuring

the

continental

Fig. (20) Tectonic basin classification. Collision-related basins and strikesliplwrench basins. (After Allen and Allen, 2006).

crust ("A-subduction", after Ampferer, under

the

Alpine-type) load

of

the

overthrust mountain belt. The extension

of

these

asymmetric basins tends to

Fig., (21) Retroarc or intramontune basins (After Allen and Allen, 2006).

increase with time, but a resulting large influx of clastic sediments from the rising mountain range of ten keeps pace with subsidence. As a result of the collision of two continental crusts, the overriding plate may be affected by "continental escape", leading to extensional graben structures or rifts perpendicular to the strike of the fold-thrust belt (Figs., 20, 21). Three stages of arc-continent collision were recognized (Escalona and Mann, 2011): 31   

Stage one of arc-continent collision Initial collision is characterized by overthrusting of the south- and southeastwardfacing Caribbean arc and forearc terranes onto the northward-subducting Mesozoic passive margin of northern South America. Northward flexure of the South American craton produces a foreland basin between the thrust front and the downward-flexed continental crust that is initially filled by clastic sediments shed both from the colliding arc and cratonic areas to the south. As the collision extends eastward towards Trinidad, this same process continues with progressively younger foreland basins formed to the east. On the overthrusting Caribbean arc and forearc terranes, north-south rifting adjacent to the collision zone initiates and is controlled by forward momentum of southward-thrusting arc terranes combined with slab pull of the underlying and subducting, north-dipping South American slab. Uplift of fold-thrust belts arc-continent suture induces rerouting of large continental drainages parallel to the collisional zone and to the axis of the foreland basins. Stage two This late stage of arc-continent collision is characterized by termination of deformation in one segment of the fold-thrust belt as convergent deformation shifts eastward. Rebound of the collisional belt is produced as the north-dipping subducted oceanic crust breaks off from the passive margin, inducing inversion of preexisting normal faults as arc-continent convergence reaches a maximum. Strain partitioning also begins to play an important role as oblique convergence continues, accommodating deformation by the formation of parallel, strike-slip fault zones and backthrusting (southward subduction of the Caribbean plate beneath the South Caribbean deformed belt). As subsidence slows in the foreland basins, sedimentation transitions from a marine underfilled basin to an overfilled continental basin. Offshore, sedimentation is mostly marine, sourced by the collided Caribbean terranes, localized islands and carbonate deposition. Stage three This final stage of arc-continent collision is characterized by: 1) complete slab breakoff of the northward-dipping South American slab; 2) east-west extension of the Caribbean arc as it elongates parallel to its strike forming oblique normal faults that produce deep rift and halfgrabens; 3) continued strain partitioning (strike-slip faulting and folding). The subsidence pattern in the Caribbean basins is more complex than interpreted before, showing a succession of extensional and inversion events. The three tectonic stages closely control the structural styles and traps, source rock distribution, and stratigraphic traps for the abundant hydrocarbon resources of the on- and offshore areas of Venezuela and Trinidad. 32   

Retroarc or intramontune basins occur in the hinterland of an arc orogen ("B-subduction" zone). They may affect relatively large areas on continental crust. Limited subsidence appears to be caused mainly by tectonic loading in a backarc fold-thrust belt. Pannonian-type basins originate from postorogenic divergence between two fold-thrust zones (Fig. 22a). They are usually associated with an A-subduction zone and are floored by thinning continental or transitional crust. During crustal collision, some foreland (and retroarc) basins can get broken up into

separate

smaller

blocks,

whereby strike-slip motions may also play a role (Fig. 22c). Some of the blocks are affected by uplift,

others

by

subsidence,

forming basinal depressions. The mechanics of such tilted block  

basins were studied, for example, in the Wyoming Province of the Rocky

Mountain

foreland

(McQueen and Beaumont 1989). So-called

Chinese-type

basins

(Bally and Snelson 1980) result from

block

faulting

in

the

hinterland of a continent-continent collision. They are not directly associated with an A-subduction margin, but it appears unnecessary

Fig. (22) Tectonic basin classification: Collision-related basins and strike-slip/wrench basins. (Allen and Allen, 2006). 

to classify them as a special new basin type (Hsii 1989). I. 9- Strike-slip and wrench basins Transform motions may be associated either with a tensional component (transtensional) or with a compressional component (transpressional). Transtensional fault systems locally cause crustal thinning and therefore create narrow (Fig. 22b), elongate pull-apart basins. If they evolve on continental crust, continuing transform motion may lead to crustal separation perpendicular to the transform faults and initiate accretion of new oceanic crust in limited spreading centers. 33   

Until this development occurs, the rate of subsidence is usually high. Transpressional systems generate wrench basins of limited size and endurance. Their compressional component can be inferred from wrench faults and fold belts of limited extent (Fig. 22c). In order to identify these various basin categories, one must know the nature of the underlying crust as well as the type of former plate movement involved during basin formation, i.e., divergence or convergence. Even in the case of transform movement, either some divergence or convergence must take place. Small angles of convergence show up as wrenching or fold belts, and small angles of divergence appear as normal faulting or sagging. One should bear in mind that all these basin types represent proto-types of tectonically controlled basins. They offer a starting point for the study and evaluation of basins, but there are no type basins which can be used as a complete model for any other basin (Burchfiel and Royden 1988). Even within a single broad tectonic setting, the development of smaller individual basins may display great variation. As soon as basins are analyzed in greater detail, the broad tectonic basin classification listed above becomes less useful. In addition, over long time periods, a sedimentary basin may evolve from one basin type into another (polyhistory basins) and thus exhibit a complex tectonic and depositional history (Allen and Allen, 2006).

II- Pre-, Syn-, and Post-Depositional Basins Principally, tectonic movements and sedimentary processes can interact in three different ways. These are used to distinguish between different types of sedimentary basins (Selley 1985a): 1- Post-depositional basins. The deposition of sediments largely predates tectonic movements forming a basin structure. Hence, there is no or little relationship between the transport, distribution, and facies of these sediments and the later evolved basin structure (Fig. 23a). However, some relationship between the syndepositional subsidence phase and the subsequent basin-forming process cannot be excluded. 2- Syn-depositional basins. Sediment accumulation is affected by syn-depositional tectonic movements, e.g., differential subsidence (Fig. 23b). If the sedimentation rate is always high enough to compensate for subsidence, the direction of transport and the sedimentary facies remain unchanged, but the thickness of the sediment in certain time slices varies. In (Fig. 23b), the sediment thickness increases toward the center of the basin. In this case, the basin structure is syn-depositional, but there was hardly a syn-depositional morphological basin controlling the sedimentary facies of the basin. If sedimentation is too slow to fill up the

34   

subsiding

area,

morphological

a

basin

will

develop. Then, the distribution and facies of the succeeding sediments will be affected by the

morphology

of

the

deepening basin. 3-

Pre-depositional

Rapid

tectonic

basins.

movements

predate significant sediment accumulation

and

create

a

morphological basin, which is filled later by post-tectonic sediments. The water depth in the basin decreases with time, although

some

syn-

depositional subsidence due to sediment loading is likely. Sediment transport as well as vertical

and

lateral

facies

development is substantially influenced

by

the

basin

morphology. Of course, there are transitions between these

Fig. (23) Post depositional basin created by tectonic after the deposition of sheet-lie fluvial and lake sediments; younger syn-tectonic basin fill is removed by subsequent erosion. b Syn-depositional tectonic movements control varying thicknesses of fluvial and shallow marine sediments and generate a basin-fill structure, although a morphological basin barely existed. c Rapid, pre-movements depositional tectonics creates a deep morphological basin which is later filled up by post-tectonic sediments. The geometry of the former basin can be derived from transport directions and facies distribution (Allen and Allen, 2006). 

simplified basin types and certain basins may show a complex history and therefore contain pretectonic as well as syntectonic or post-tectonic sediments (Allen and Allen, 2006).

35   

5- Basins Morphology The geometry of an ultimate basin fill is controlled mainly by basin-forming tectonic processes, but the morphology of a basin defined by the sediment surface is the product of the interplay between tectonic movements and sedimentation. Therefore a purely tectonic classification of sedimentary basins is not sufficient for characterizing depositional areas. It is true that a sedimentary basin in a particular tectonic setting also often undergoes a specific developmental or subsidence history, but its morphology, including water depth, may be controlled largely by other factors, such as varying influx and distribution of sediment from terrigenous sources (Allen and Allen, 2006). For example, a fluvial depositional system can develop and persist for considerable time on top of subsiding crust in various tectonic settings (Miall 1981). Fluvial deposits are known graben

from

continental

structures,

passive

continental margins, foreland basins, forearc and backarc basins, pull-apart basins, etc. Fluvial sediments (Fig. 24) accumulate as long as rivers reach the depositional area

Fig. (24) Base level of erosion, hydrodynamic regime in the sea, and gravity mass movements as limiting factors controlling upbuilding and outbuilding of sediments modified by sea level changes. (Allen and Allen, 2006). 

and supply enough material to keep the subsiding basin filled. Although the basin-forming processes and subsidence histories of these examples differ fundamentally from each other, the sedimentary facies of their basin fills display no or only minor differences. In order to distinguish between these varying tectonic settings, one has to take into account the geometry of the entire basin fill, as well as vertical and lateral facies changes over long distances, including paleocurrent directions and other criteria. Syndepositional tectonic movements manifested by variations in thickness, small disconformities, or faults dying out upward (Fig., 25) may indicate the nature of the tectonic processes involved. The erosional base level and sediment distribution within a basin are additional important factors modifying basin morphology and thus the development of special sedimentary facies (Fig., 25). In a fluvial environment, sediments cannot accumulate higher than the base level and gradient of the stream. If there is more influx of material into the depositional system than necessary for compensation of subsidence, the sediment surplus will 36   

be carried farther downslope into lakes or the sea. This signifies that the level up to which a basin can be filled with sediments may depend

on

the

geographic

position of the basin in relation to the erosional base. The morphology of waterfilled basins may significantly change as a result of depositional processes. Lakes and low-energy basins

frequently

show

a

prograding deltaic facies, causing pronounced

basinward

outbuilding

of

Consequently,

sediment. the

areal

distribution of the finer-grained sediment in the deeper basin portions decreases with time, although the initial, tectonically controlled

basin

configuration

persists. By contrast, high-energy

Fig. (25). Overview of depositional environments, based primarily on basin morphology and peri-basin characteristics. All basins, particularly those on land (a) or adjacent to continents (b and c), are strongly affected by variations in terrigenous input under differing conditions of climate and relief. d Various marine basins (Allen and Allen, 2006). 

basins are little influenced by sediment outbuilding (Fig., 25). For example, terrigenous sediments transported into highenergy shelf seas tend to be reworked and swept into deeper water by wave action and bottom currents, except for some local seaward migration of the shoreline. Even on deep submarine slopes and in the deep sea, there is no general outbuilding or upbuilding of sediments, because gravity mass movements and deep bottom currents redistribute large quantities of material. These few examples demonstrate that the most appropriate classification scheme for sedimentary basins depends primarily on the objectives of the study. If tectonic structure and evolution of a region are the main topics, then basin fill geometry and subsidence history derived from the thickness of stratigraphic units are of primary importance. If, on the other hand, the depositional environment, sedimentary facies, and paleogeographic reconstructions are of primary interest, then the basin classification used should not be strictly tectonic. Such 37   

a classification should also take into account changes in basin morphology caused by depositional processes, the chemical and hydrodynamic regimes of the basin, and peri-basin characteristics such as the size and nature of the drainage areas on nearby land. The surface of recent sediments on land and under water can be well observed, but in many cases, for example in fluvial environments, such temporary surfaces are rarely preserved in the sedimentary record. By contrast, indurated beds alternating with weaker material frequently show excellently preserved lower and upper bedding planes with trace fossils, various marks, and imbrication phenomena which are difficult to observe in soft sediments. Diagenesis may, however, also obscure primary bedding features.

6- Depositional Environments On the surface of our present-day globe, on land and below the sea, hundreds of depositional areas are known which meet

the

definition

of

sedimentary basins. In the various

types

sedimentary

basins

predominantly

of are

classified

according

to

their

depositional

environment

Fig., (26) Types of continental Sedimentary Environments.

and basin morphology (Fig., 26) including: 1- Continental (fluvial, glacial, eolian), lacustrine, and deltaic environments. 2- Adjacent sea basins and epicontinental seas of varying salinity. 3- Marine depositional areas of normal salinity. 4- Transitional environments may be defined between continental and marine environments. This group includes marine deltas, intertidal environments, coastal lagoons, estuaries, and barrier island systems.

38   

1- Continental sediment environments: 1.1- Glacial Environments: The continental glacial deposits (Fig., 27) generally have a low preservation potential in the long term and are rarely incorporated into the stratigraphic record. Glacial processes which bring sediment into the marine environment generate deposits that have a much higher chance of long-term preservation, and recognition of the characteristics sediments

of can

these provide

important

clues

climates

(Nichols,

2009).

deposits

are

Glacial

about

past

compositionally immature and tills are typically composed of detritus that simply represents broken

up

and

bedrock

from

glacier.

Reworked

powdered

beneath

the

Fig., (27) Glacial landforms and glacial deposits in continental glaciated areas After (Nichols, 2009). 

glacial

deposits on outwash plains may show a slightly higher compositional and textural maturity. There is a paucity of clay minerals in the fine-grained fraction because of the absence of chemical weathering processes in cold regions. Continental glacial deposits have a relatively low preservation potential in the stratigraphic record, but erosion by ice in mountainous areas is an important process in supplying detritus to other depositional environments. Glaciomarine deposits are more commonly preserved, including dropstones which may provide a record of periods of glaciation in the past (Nichols, 2009). Quaternary valley and piedmont glaciers form distinctive moraines but are largely confined to upland areas that are presently undergoing erosion. Of more interest from the point of view of the stratigraphic record are the tills formed in lowland continental areas and in marine environments as these are much more likely to lie in regions of net accumulation in a sedimentary basin. The volume of material deposited by ice sheets and ice shelves is also considerably greater than that associated with upland glaciations (Nichols, 2009). 1.2- Aeolian Environments: Aeolian sedimentary processes are those involving transport and deposition of material by the wind (Fig., 28). The whole of the surface of the globe is affected by the wind to varying degrees, but aeolian deposits are only dominant in a relatively restricted range of settings. The most obvious aeolian environments are the large sandy deserts in hot, dry areas of continents, but there are significant accumulations of wind-borne 39   

material associated with sandy beaches and periglacial sand flats. Sands deposited in these desert areas are characteristically both

compositionally

mineralogically

and

mature

with

large-scale cross-bedding formed by

the

migration

of

dune

bedforms. Oxidising conditions in

deserts

preservation

of

preclude

the

much

fossil

material,

and

sediments

typically

red–yellow

are

colours

(Nichols, 2009). Aeolian dust

Fig. (28) Depositional environments in arid regions: coarse material is deposited on alluvial fans, sand accumulates to form aeolian dunes and occasional rainfall feeds ephemeral lakes where mud and evaporite minerals are deposited (Nichols, 2009). 

deposits are deposits of Quaternary age in Eastern Europe, North America and China that are interpreted as accumulations of wind-blown dust (Pye 1987). These deposits, known as loess, locally occur in beds several meters thick made up predominantly of well-sorted silt-sized material, with little clay or sand-sized material present. Associated facies in arid regions are mud and evaporites deposited in ephemeral lakes and poorly sorted fluvial and alluvial fan deposits. Aeolian deposits are less common outside of desert environments, occurring as local sandy facies associated with beaches and glaciers, and as dust distributed over large distances into many different environments, but, apart from Quaternary loess, rarely in significant quantities (Nichols, 2009). 1.3- Rivers and Alluvial Fans: Rivers are an important feature of most landscapes, acting as the principal mechanism for the transport of weathered debris away from upland areas and carrying it to lakes and seas, where much of the clastic sediment is deposited. River systems can also be depositional, accumulating sediment within channels and on floodplains. The grain size and the sedimentary structures in the river channel deposits are determined by the supply of detritus, the gradient of the river, the total discharge and seasonal variations in flow. Overbank deposition consists mainly of finer-grained sediment, and organic activity on alluvial plains contributes to the formation of soils, which can be recognized in the stratigraphic record as palaeosols. Water flows over the land surface also occur as unconfined sheet floods and debris flows that form alluvial fans at the edges of alluvial plains. Fluvial and alluvial deposits in the stratigraphic record provide evidence of tectonic activity and indications of the palaeoclimate at the time of deposition (Nichols, 2009). 40   

The fluvial environment is controlled by its erosional base level as well as by the sediment supply from more elevated regions sufficient to compensate for subsidence in different tectonic settings (Einsele, 1992). Fluvial environments (Fig., 29a) are characterised by flow and deposition in river channels and associated overbank sedimentation (Nichols, 2009).

Under

circumstances,

these

the

river

gradient and thus a more or less

constant

average

net

transport direction can be maintained for rather long time periods. A topographic depression,

i.e.,

a

syndepositional morphological basin can only develop when fluvial transport lags behind basin

subsidence

(Einsele,

1992).

Fig. (29a). The geomorphological zones in alluvial and fluvial systems: in general braided rivers tend to occur in more proximal areas and meandering rivers occur further downstream (Nichols, 2009).

Three

geomorphological zones can be recognized within fluvial and alluvial systems (Einsele, 1992). In the erosional zone the

streams

downcutting,

are

actively removing

bedrock from the valley floor and from the valley sides via downslope

movement

of

Fig. (29b). Types of alluvial fan: debris-flow dominated, sheetflood and stream-channel types – mixtures of these processes can occur on a single fan (Nichols, 2009). 

material into the stream bed. In the transfer zone, the gradient is lower, streams and rivers are not actively eroding, but nor is this a site of deposition. The lower part of the system is the depositional zone, where sediment is deposited in the river channels and on the floodplains of a fluvial system or on the surface of an alluvial fan. These three components are not present in all systems: some may be wholly erosional as far as the sea or a lake, and others may not display a transfer zone. The erosional part of a fluvial system contributes a substantial proportion of the clastic sediment provided for deposition in other sedimentary environments. In the stratigraphic record the 41   

channel fills are represented by lenticular to sheet-like bodies with scoured bases and channel margins, although these margins are not always seen. The deposits of gravelly braided rivers are characterised by crossbedded conglomerate representing deposition on channel bars (Nichols, 2009). Both sandy braided river and meandering river deposits typically consist of finingupward successions from a sharp scoured base through beds of trough and planar crossbedded, laminated and cross-laminated sandstone. Lateral accretion surfaces characterize meandering rivers that are also often associated with a relatively high proportion of overbank facies. Floodplain deposits are mainly alternating thin sandstone sheets and mudstones with palaeosols; small lenticular bodies of sandstone may represent crevasse splay deposition. Palaeocurrent data from within channel deposits are unidirectional, with a wider spread about the mean in meandering river deposits; palaeocurrents in overbank facies are highly variable. Alluvial fans are cones of detritus that form at a break in slope at the edge of an alluvial plain (Fig., 29b). They are formed by deposition from a flow of water and sediment coming from an erosional realm adjacent to the basin. The term alluvial fan has been used in geological and geographical literature to describe a wide variety of deposits with an approximately conical shape, including deltas and large distributary river systems. Alluvial fans form where there is a distinct break in topography between the high ground of the drainage basin and the flatter sedimentary basin floor (Einsele, 1992). Alluvial fan deposits are located near to the margins of sedimentary basins and are limited in lateral extent to a few kilometers from the margin. The facies are dominantly conglomerates, and may include matrix-supported fabrics deposited by debris flows, well-stratified gravels and sands deposited by sheetflood processes and in channels that migrate laterally across the fan surface. Alluvial and fluvial deposits will interfinger with lacustrine and/or aeolian facies, depending on the palaeoclimate, and many (but not all) river systems feed into marine environments via coasts, estuaries and deltas. 1.4- Lakes and Lacustrine Environments Lakes are an inland body of water. Although some modern lakes may be referred to as ‘inland seas’, it is useful to draw a distinction between water bodies that have some exchange of water with the open ocean (as lagoons) and those that do not, which are true lakes. Lakes form where there is a supply of water to a topographic low on the land surface. They are fed mainly by rivers and lose water by flow out into a river and/or evaporation from the surface. Lakes form where there is a depression on the land surface which is bounded by a sill such 42   

that water accumulating in the depression is retained. Lakes are typically fed by one or more streams that supply water and sediment from the

surrounding

hinterland.

Groundwater may also feed water into a lake. Sand and mud are the most common components of lake deposits. The amount of sediment accumulated in lakes is small compared with marine basins, but they may be locally significant, resulting in strata hundreds of meters thick and covering hundreds

Fig. (30). Hydrological regimes of lakes (Nichols, 2009). 

to thousands of square kilometers. The balance between inflow and outflow and the rate at which evaporation occurs control the level of water in the lake and the water chemistry. Under conditions of high inflow the water level in the lake may be constant, governed by the spill point of the outflow, and the water remains fresh. Low water input coupled with high evaporation rates in an enclosed basin results in the concentration of dissolved ions, which may be precipitated as evaporites in a perennial saline lake or when an ephemeral lake dries out. Lakes are therefore very sensitive to climate and climate change. Many of the processes that occur in seas also occur in lakes: deltas form where rivers enter the lake, beaches form along the margins, density currents flow down to the water bottom and waves act on the surface. There are, however, important differences with marine settings: the fauna and flora are distinct, the chemistry of lake waters varies from lake to lake and certain physical processes of temperature and density stratification are unique to lacustrine environments (Nichols, 2009). In lacustrine sedimentation, terrigenous materials entering the basin may come either from one or several nearby sources, or, solely or in addition, from a distant source (Fig., 30). Consequently, deposition will be either texturally immature or markedly mature and display either a fairly uniform or complex composition. In addition, the climate in the source area(s) exerts a strong influence. Where sediment accumulation cannot compensate for subsidence, long persisting, deepening lakes or shallow seas evolve (Einsele, 1992). Other characteristics of fluvial and alluvial facies include an absence of marine fauna, the presence of land plant 43   

fossils, trace fossils and palaeosol profiles in alluvial plain deposits (Nichols, 2009). Limestones, evaporites and organic material are of lacustrine deposits as well as plants and animals living in a lake may be preserved as fossils in lacustrine deposits, and concentrations of organic materials can form beds of coal. The characteristics of the deposits of lacustrine environments are controlled by factors that control the depth and size of the basin (which are largely determined by the tectonic setting), the sediment supply to the lake (which is a function of a combination of tectonics and climatic controls on relief and weathering) and the balance between water supply and loss through evaporation (which is principally related to the climate). If the climate is humid a lake will be hydrologically open, with water flowing both in and out of it. Such lakes can be considered to be overfilled (Bohacs et al. 2000, 2003), and their deposits are characterised by accumulation both at the margins, where sediment is supplied to deltas and beaches, and in the deep water from suspension and turbidity currents. The lake level remains constant, so there is no evidence of fluctuations in water depth under these conditions (Nichols, 2009). The majority of large modern lakes are freshwater lakes; they occur at latitudes ranging from the Equator to the Polar Regions (Bohacs et al. 2003) and include some of the largest and deepest in the world today. Lacustrine deposits from lakes of similar scales are known from the stratigraphic record, mainly from Devonian through to Neogene strata (Nichols, 2009). Saline lakes are perennial, supplied by rivers containing dissolved ions weathered from bedrock and in a climatic setting where there are relatively high rates of evaporation. The salinity may vary from 5 g L-1 of solutes, which is brackish water, to saline, close to the concentration of salts in marine waters, to hypersaline waters, which have values well in excess of the concentrations in seawater. From a sedimentological point of view, brackish water lakes are similar to freshwater lakes because it is the high concentrations of salts that provide saline lakes with their distinctive character (Nichols, 2009).

2- Marine Sediment Environments The physical processes of tides, waves and storms in the marine realm define regions bounded by water depth changes. The beach foreshore is the highest energy depositional environment where waves break and tides regularly expose and cover the sea bed. At this interface between the land and sea storms can periodically inundate low-lying coastal plains with seawater. Across the submerged shelf, waves, storms and tidal currents affect the sea bed to different depths, varying according to the range of the tides, the fetch of the waves and the 44   

intensity of the storms. Sedimentary structures can be used as indicators of the effects of tidal currents, waves in shallow water and storms in the offshore transition zone. Further clues about the environment of deposition are available from body fossils and trace fossils found in shelf sediments (Nichols, 2009). 2.1- Marine deltas represent a transitional, highly variable depositional environment between continental

and

marine

conditions. A delta can be defined

as

shoreline

a

‘discrete

protuberance

formed at a point where a river enters the ocean or other body of water’ (Fig. 31) (Janok

et

al.,

2003;

Bhattacharya 2006), and as such it is formed where sediment brought down by the river builds out as a body into the lake or sea (Nichols,

Fig. (31) The forms of modern deltas: (a) the Nile delta, the ‘original’ delta, (b) the Mississippi delta, a river-dominated delta, (c) the Rhone delta, a wave-dominated delta, (d) the Ganges delta, a tide-dominated delta (Nichols, 2009). 

2009). In marine settings the interaction of subaerial processes with wave and tide action results in complex sedimentary environments that vary in form and deposition according to the relative importance of a range of factors (Nichols, 2009). The subaerial part of such a delta is controlled by fluvial and possibly lacustrine processes, whereas its coastal and subaqueous regions are dominated by the hydrodynamic and chemical properties of the sea (Einsele, 1992). Delta form and facies are influenced by the size and discharge of the rivers, the energy associated with waves, tidal currents and longshore drift, the grain size of the sediment supplied and the depth of the water. They are almost exclusively sites of clastic deposition ranging from fine muds to coarse gravels. Deposits formed in deltaic environments are important in the stratigraphic record as sites for the formation and accumulation of fossil fuels. Large terrigenous sediment supply causes prograding of the deltaic complex toward the sea; high sedimentation rates and subsidence enhanced by the sediment load enable the formation of thick, widely extended deltaic sequences. Marine delta complexes provide a particularly good example of depositional environments which are controlled predominantly by exogenic factors (Einsele, 1992). 45   

2.2- Clastic Coasts and Estuaries Coasts are the areas of interface between the land and the sea, and the coastal environment can comprise a variety of zones, including coastal plains, beaches, barriers and lagoons

(Fig.,

32).

The

shoreline is the actual margin between the land and the sea. Coastlines can be divided into two general categories on

the

basis

of

their

morphology, wave energy and sediment budget. The morphology of coastlines is very variable, ranging from

Fig. (32) Reflective coasts are usually erosional with steep beaches and a narrow surf zone. Dissipative coasts may be depositional, with sand deposited on a gently sloping foreshore (Nichols, 2009). 

cliffs of bedrock to gravelly or sandy beaches to lower energy settings where there are lagoons or tidal mudflats. Wave and tidal processes exert a strong control on the morphology of coastlines and the distribution of different depositional facies. Wave-dominated coasts have well-developed constructional beaches that may either fringe the coastal plain or form a barrier behind which lies a protected lagoon. Barrier systems are less well developed where there is a larger tidal range and the deposits of intertidal settings, such as tidal mudflats, become important. Estuaries are coastal features where water and sediment are supplied by a river, but, unlike deltas, the deposition is confined to a drowned river valley (Nichols, 2009). Erosional coastlines typically have relatively steep gradients where a lot of the wave energy is reflected back into the sea from the shoreline: both bedrock and loose material may be removed from the coast and redistributed by wave, tide and current processes. At depositional coastlines the gradient is normally relatively gentle and a lot of the wave energy is dissipated in shallow water: provided that there is a supply of sediment, these dissipative coasts can be sites of accumulation of sediment (Woodroffe, 2002). 2.3- The beach is the area washed by waves breaking on the coast. The seaward part of the beach is the foreshore, which is a flat surface where waves go back and forth and which is gently dipping towards the sea (Fig., 33a). Where wave energy is sufficiently strong, sandy and gravelly material may be continuously reworked on the foreshore, abrading clasts of all sizes to a high degree of roundness, and effectively sorting sediment into different sizes 46   

(Nichols, 2009). Sandy sediment is deposited in layers parallel to the slope of the foreshore, dipping offshore at only a few degrees to the horizontal (much less than the angle of repose). This low-angle stratification of well-sorted, well-rounded sediment is particularly characteristic of wave-dominated sandy beach environments (Clifton, 2006). Grains are typically compositionally mature as well as texturally mature because the continued abrasion in the beach swash zone tends to break down the weaker clasts (Nichols, 2009). 2.4- Coastal plains are lowlying areas adjacent to seas. They are part of the continental environment where there are fluvial,

alluvial

or

aeolian

processes of sedimentation and pedogenic modification. Coastal plains are influenced by the adjacent

marine

environment

Fig. (33a) Morphological features of a beach comprising a beach foreshore and backshore separated by a berm; beach dune ridges are aeolian deposits formed of sand reworked from the beach (Nichols, 2009). 

when storm surges result in extensive flooding by seawater. A deposit related to storm flooding can be recognised by features such as the presence of bioclastic debris of a marine fauna amongst deposits that are otherwise wholly continental in

Fig. (33b) A wave-dominated coastline with a coastal plain bordered by a sandy beach: chenier ridges are relics of former beach strand plains (Nichols, 2009).

character (Fig., 33b, c). Sandy coastlines where an extensive area of beach deposits lies directly adjacent to the coastal plain are known as strand plains. Along coasts supplied

Fig. (33c) Morphological features of a coastline influenced by wave processes and tidal currents (Nichols, 2009). 

with sediment, beach ridges create strand plains that form sediment bodies tens to hundreds of meters across and tens to hundreds of kilometers long and progradation of strand plains can produce extensive 47   

sandstone bodies. The strand plain is composed of the sediment deposited on the foreshore and backshore region. The backshore area merges into the coastal plain and may show evidence of subaerial conditions such as the formation of aeolian dunes and plant colonization (Nichols, 2009). 2.5- Beach barriers are composed of sand and/or gravel material and are largely built up by wave action. They may be partially attached to the land, forming a beach spit, or wholly attached as a welded barrier that completely encloses a lagoon, or can be isolated as a barrier island in front of a lagoon (Fig., 34). In practice, the distinction between these three forms can be difficult to identify in ancient successions and their sedimentological characteristics are very similar. Barriers range in size from less than 100m wide to several kilometers and their length ranges from a few hundred meters to many tens of kilometers (Davis and Fitzgerald 2004). The largest tend to form along the open coasts of large oceans where the wave energy is high and the tidal range is small (Nichols, 2009). Lagoons are coastal bodies of water that have very limited connection to the open ocean. Seawater

reaches

a

lagoon

directly through a channel to the

Fig. (34). Distribution of depositional settings in a wave-dominated estuary (Nichols, 2009). 

sea or via seepage through a barrier; fresh water is supplied by rainfall or by surface run-off from the adjacent coastal plain. If a lagoon is fed by a river it would be considered to be part of an estuary system. They are typically very shallow, reaching only a few meters in depth (Nichols, 2009). An estuary is the marine-influenced portion of a drowned valley (Dalrymple et al. 1992). A drowned valley is the seaward portion of a river valley that becomes flooded with seawater when there is a relative rise in sea level. They are regions of mixing of fresh and seawater. Sediment supply to the estuary is from both river and marine sources, and the processes that transport and deposit this sediment are a combination of river and wave and/or tidal processes. An estuary is different from a delta because in an estuary all the sedimentation occurs within the drowned valley, whereas deltas are progradational bodies of sediment that build out into the marine environment. A stretch of river near the mouth that does not have a marine influence would not be considered to be an estuary (Nichols, 2009).

48   

2.6- Shallow Marine Carbonate and Evaporite Environments: Limestones are common and widespread sedimentary rocks that are mainly formed in shallow marine depositional environments. Most of the calcium carbonate that makes up limestone comes from biological sources, ranging from the hard, shelly parts of invertebrates such as molluscs to very fine particles of calcite and aragonite formed by algae. The accumulation of sediment

in

carbonate-forming

environments is largely controlled by

 

factors that influence the types and abundances of organisms that live in them. Water depth, temperature, salinity, nutrient availability and the supply of terrigenous clastic material all influence carbonate deposition and the buildup of successions of

limestones.

Some

depositional

environments are created by organisms, for example, reefs built up by sedentary colonial

organisms

such

as

corals.

Evaporite deposits in modern marine

Fig. (35). Settings where barred basins can result in thick successions of evaporates (Nichols, 2009). 

environments are largely restricted to coastal regions, such as evaporate lagoons and sabkha mudflats. However, evaporite successions in the stratigraphic record indicate that precipitation of evaporate minerals has at times occurred in more extensive marine settings (Fig., 35). 2.7- Adjacent sea basins and epicontinental seas are connected with the open sea and therefore exchange basin water with normal ocean water (Einsele, 1992). The extent of this water exchange and thus the salinity of the basin water strongly depend on the width and depth of the opening to the ocean. In humid regions, adjacent basins with a limited opening tend to develop brackish conditions, while arid basins frequently become more saline than normal sea water. Adjacent basins and epicontinental basins on continental crust are commonly shallow, but basins on oceanic or mixed crust may also be deep. All these basins may show either symmetric or asymmetric cross sections, and they may represent either simple morphological features or basins subdivided by shallow swells into several sub-basins (segmented basins). In the latter case, markedly differing depositional sub-environments have 49   

to be taken into account. Most of these adjacent basins are still strongly influenced by the climate and relief of peri-basin land regions, which control the influx of terrigenous material from local sources. In addition, more distant provenances may contribute to the sediment fill. In summary, adjacent basins may exhibit a particularly great variety of facies (Einsele, 1992). 2.8- The shallow seas and continental shelf sediments are still considerably affected by processes operating in neighboring land regions, which generally provide sufficient material to keep these basins shallow. Strong waves, and surface and bottom currents usually tend to distribute the local influx of terrigenous sediment over large areas. Especially in shallow water, the high-energy, sediment-transporting systems prevent the deposition of fine-grained materials, partially including sands. Therefore, such areas often persist over long time periods without being filled up to sea level. This is also true for widely extended shallow-marine basins, as long as excess sediment volume (in relation to space provided by subsidence) can be stored in special depressions (Einsele, 1992) or be swept into a neighboring deeper ocean basin. The margin of such basins is commonly characterized by a kind of ramp morphology. 2.9- Deep-sea basins or basin plains are the deepest parts of marine environments except for the special features of deep-sea trenches. Large volumes of terrigenous material can also be collected by the troughs in a submarine horst and graben topography bordering the continent (Fig., 36). Similarly, deep sea trenches at the foot of relatively steep slopes and slope basins are sites of preferential sediment accumulation (Einsele, 1992). Thick, ancient flyisch sequences are

mostly

interpreted

as

depositions in such basins. Less

Fig. (36). Deep water environments are floored by ocean crust and are the most widespread areas of deposition worldwide (Nichols, 2009). 

important sediment accumulation features are small basins, called "ponds", which occur along oceanic ridges, and infillings of narrow troughs due to fracturing of the oceanic crust. The thin, frequently incomplete sedimentary records on the tops of submarine ridges, platforms, and seamounts strongly contrast with all other marine sediments. These deposits are mostly biogenic or chemically precipitated and usually contain only very small proportions of terrigenous or volcaniclastic materials. Although such limited sediment accumulations can hardly be referred to as basin fills, they do constitute an important and diagnostically significant part of larger marine depositional environments. The direct influence of tectonic basin evolution on sedimentary facies is only evident in areas, where tectonic movements are 50   

rapid and non-uniform, such as at the basin margins, or where sediment accumulation lags far behind subsidence. This situation is common in continental rift and pull-apart basins during their early stages of evolution, in subduction-related settings, in remnant and foreland basins, and in deep marine environments along oceanic ridges or transform faults far away from large land masses.

51   

CHAPTER II GEOTHERMAL ENERGY IN THE SEDIMENTARY BASINS In geothermal reservoirs, heat is created within the mantle or crust through the decay of radioactive isotopes (Fig.; 37). Within a sedimentary basin, this heat is transferred to the surface through conduction and convection of

fluids.

Current

geothermal

gradients are controlled by the combination of conduction and convection, and can vary due to the relative importance of each (Graf, 2009). Studies of the present day heat flows and ancient geothermal gradients

suggest

that

Fig., (37). Block model of geological formations that represent a geothermal reservoir (source:  http://www.eia.gov/cneaf/ solar.renewables/renewable.energy.annual/backgrnd/fig19.htm). 

thermal

regime closely reflects tectonic history. In particular, hypothermal (cooler than average) basins include ocean trenches and outer forearcs and foreland basins. Hyperthermal (hotter than average) basins include oceanic and continental rifts, some strike-slip basins with mantle involvement, and magmatic arcs in collisional settings. Mature passive margins that are old compared with the thermal time consist of the lithosphere tend to have nearaverage heat flows and geothermal gradients (Allen and Allen, 2006). Changes in physical and chemical conditions during basin evolution control the interaction

Fig., (38) The major geothermal energy location around the world (after Bhattacharya, 2011) 

between pore fluids and rocks. Sedimentary rocks consolidate and may be cemented or dissolved, thereby changing their chemistry, texture and ability to transmit fluids, solutes and heat. At the same time, fluids change their hydrochemical composition and may become more

52   

diluted or more concentrated (Bitzer et al., 2001). As these fluids change in temperature, they may dissolve when mixed with other fluids or if a further change of temperature occurs. Flowing groundwater takes up the geothermal energy from the Earth’s crust and transports part of it in the direction of water flow. Productivity of a geothermal reservoir is controlled predominantly by the geothermal gradient (i.e., temperature variation with depth) encountered in a basin. Extremely high gradients (200°C/km) are observed along oceanic spreading centers (e.g., the Mid- Atlantic Rift) and along island arcs (e.g., the Aleutian chain) (Fig., 38). In Iceland, geothermal energy, the main source of energy, is extracted from areas with geothermal gradients ≥40°C/km. Low gradients are observed in tectonic subduction zones because of thrusting of cold, water-filled sediments beneath an existing crust. Tectonically stable shield areas and sedimentary basins have average gradients that typically vary from 15 C/km to 30°C/km (Graf, 2009). The geothermal provinces at India are associated with major rifts or subduction tectonics and registered high heat flow and high geothermal gradient. The reservoir temperatures estimated are 120° C (west coast), 150° C (Tattapani) and 200° C (Cambay). The depth of the reservoir in these provinces is at a depth of about 1 to 2 km. These geothermal systems are liquid dominated and steam dominated systems prevail only in Himalayan and Tattapani geothermal provinces (Bhattacharya, 2011) Geothermal energy indicates that part of the heat within the Earth that can or might be recovered and exploited by mankind. Due to the long-term availability and the large extent of geothermal heat, geothermal energy represents an efficient renewable energy worldwide. Making geothermal heat an effective source for a sustainable supply of energy requires a quantitative reserve and resource assessment. Though immense in its nature, only a fraction of the Earth’s heat can be utilized in practice, its exploitation being limited to areas characterized by favorable hydrogeological conditions for geothermal resources to develop. A proper geothermal exploration involves different stages comprising: (1) a correct localization of potential areas to ascertain the existence of a particular geothermal field; (2)an accurate estimate of the size of the resource to determine the type of geothermal field; and (3) an appropriate identification of the main physical transport processes involved to properly identify geothermal phenomena. This requires an integrated approach involving different disciplines and methodologies including geological field measurements, laboratory-based investigations as well as mathematical modeling. It is well known (Bethke et al., 1988; Raffensperger and Vlassopoulos, 1999) that the most significant portion of the world’s mineral, energy and water resources is hosted in sedimentary basins. Formation of these resources results from interactions between different coupled processes comprising 53   

groundwater flow, mechanical deformation, mass transport and heat transfer and different water–rock interaction mechanisms. Understanding the relative impact of fluid and other heat driving processes on the resulting geothermal field as well as the resulting subsurface flow dynamics is of crucial importance for geothermal energy production. For geothermal exploration it is essential to quantify the above-mentioned processes by interpretation of their characteristic thermal signatures in the subsurface. This requires a correct interpretation of the impact of all processes contributing to the temperature field to not misinterpret similar, but distinct in nature, thermal signatures. Heat in the crust is mainly transferred by diffusion. In sedimentary basins, an additional mean of heat transport is provided by advective forces by ground water circulating through permeable aquifers (Andersaon, 2005). Temperatures in the model are governed by the effects of vertical and horizontal thermal conduction such that the lithosphere-asthenosphere boundary is defined as a partial melt isotherm or phase change boundary which migrates vertically depending on the transient thermal state. Vertical deformations of the lithosphere result from the purely mechanical effects of sediment loading as well as from changes in the ambient temperature field. The temperature anomalies contribute to these deformations not only by setting up body forces but also by creating thermal in plane forces and associated bending units (Stephenson et al., 1989). 1- Geothermal Gradient The geothermal or temperature gradient is the rate of increase in temperature per unit depth in the Earth due to the outflow of heat

from

the

centre.

The

temperature gradient between the centre of the Earth and the outer limits of the atmosphere averages about 1°C per kilometer (Fig., 39).

To

classify

geothermal

systems, Tester et al., (2006) divided geothermal resources into high- (>150⁰C), medium- (50– 150⁰C)

and

low-

(<50⁰C)

Fig., (39) Structure of the Earth and the geothermal gradient. (Source: http://www.mpoweruk.com/geothermal_energy.htm). 

temperature resources. The low 54   

temperature resources were used for direct heating applications. The temperature gradient in the Earth's fluid layers and the magma tend to be lower because the mobility of the molten rock tends to even out the temperature. This mobility however does not exist in the solid crust where temperature gradient is consequently much higher, typically between 25°C and 30°C per kilometer depending on the location and higher still in volcanic regions and along tectonic plate boundaries where seismic activity transports hot material to near the surface. (Source: http://www.mpoweruk.com/geothermal_energy.htm)

2- Effect of the geothermal energy on hydrocarbon maturation Subsidence in sedimentary basins causes thermal maturation in the progressively buried sedimentary layers. Indicators of the thermal history include; organic, geochemical, mineralogical and thermochronometric parameters. The most important factors in the maturation of organic matter are temperature and time, pressure being relatively un-important. This temperature and time dependency describes the reaction rate increases exponentially with temperature, the rate of increase (Allen and Allen, 2006). The combined effects of sedimentary processes and heat flow are the prime control on the rate and extent of hydrocarbon maturation in potential source rocks, which is of prime interest in oilfield appraisal. Hydrocarbons generated by organic matter rich sediments may be transported towards reservoir rocks, if physico-chemical conditions and timing are appropriate. Flow, transport and reaction in the scale of sedimentary basins are in most cases slow and steady processes. However, over the scale of geologic time, its effects are of great importance as they can generate important resources (Bitzer et al., 2001). The maturation of the hydrocarbons involves the slow thermodynamic conversion of the organic matter (Kerogens) in potential source rock into oil and gas, which may then migrate to more porous reservoir rocks. The maturation process is heavily influenced by two factors; the local temperature and the duration of the thermal event. In turn, these are strongly controlled by the rates of subsidence and sedimentation. During basin forming events, large amounts of heat are transferred from the basement through the evolving sedimentary cover, providing an energy source for the hydrocarbon maturation processes (Palumbo et al., 1999; Gray et. al. 2012). As in any ‘slow cooking’ process, however, maturation can occur at a given temperature only if the effective heating time is long enough. The maturation index, which depends on both the effective heating time and the thermal history, is a quantitative measure of the degree of maturation. (Pieri 1988; Cranganu and Deming 1996).

55   

3- Geothermal energy utilizations Direct-use of geothermal energy is one of the oldest, most versatile and also the most common form of utilization of geothermal energy (Dickson and Fanelli, 2003). The early history of geothermal direct-use has been well documented for over 25 countries. Cataldi et al., (1999) documents that the geothermal uses are for over 2,000 years. Now, there are 78 countries having direct utilization of geothermal energy, is a significant increase from the 72 reported in 2005, the 58 reported in 2000, and the 28 reported in 1995 (Lund et al., 2010). The thermal energy used is 438,071 TJ/year (121,696 GWh/yr), about a 60% increase over 2005, growing at a compound rate of 9.9% annually. The distribution of thermal energy used by category is approximately 49.0% for ground-source heat pumps, 24.9% for bathing and swimming (including balneology), 14.4% for space heating (of which 85% is for district heating), 5.3% for greenhouses and open ground heating, 2.7% for industrial process heating, 2.6% for aquaculture pond and raceway heating, 0.4% for agricultural drying, 0.5% for snow melting and cooling, and 0.2% for other uses (Table 2). About Egypt, no data were submitted for WGC2005 or WGC2010. A spa at Hammam Faraun is also reference in Lashin and Al Arifi (2010). The estimates in Lund et al. (2005) of 1.0 MWt and 15 TJ/yr are assumed to still be valid. There are two main exploitable sources of geothermal energy. Hydrothermal systems, first demonstrated in 1904, used the naturally occurring hot water or steam trapped in

or

through

circulating permeable

rock, to drive steam powered generators.

Table (2): Summary of geothermal energy use by continent in 2000, showing contribution of Europe (Fridleifsson, 2002; based on Huttrer, 2001; Lund and Freeston, 2001)

electricity More

recently, since 1970, technology has been developed to extract the heat from hot rock by artificially circulating water through the rock to produce super-heated water or steam to drive the generators. For cost efficient electricity generation, suitable temperatures for hot water and steam range upwards from 120°C to 370°C. Such naturally occurring hydrothermal resources are not widely available and are found in only a few regions of the world where the Earth's crust is very thin, usually around the edges of the crustal tectonic plates. Geothermal electricity 56   

generating plants have been installed in over twenty countries with new installations planned in several more. In shallow reservoirs or regions where the water or steam temperature may range between 21°C to 149°C and not be hot enough for efficient electricity generation, the hot water can be used directly for local heating applications. Iceland is widely considered the success story of the geothermal community. The country of just over 300,000 people is now fully powered by renewable forms of energy, with 17% of electricity and 87% of heating needs provided by geothermal energy (fossil fuels are still imported for fishing and transportation needs (Blodgett and Slack, 2009). 3.1- Hydrothermal Systems - Geothermal Aquifers Conventional hydrothermal systems make use of geothermal aquifers which are naturally occurring geological formations of permeable rock or unconsolidated sediment (gravel, sand, silt, or clay) in which water may accumulate, between layers of impermeable rock. Where these aquifers occur in fractured volcanic rocks where temperatures are relatively high near the surface or in non volcanic areas where the crustal heat flow is very high, the water temperature may be high enough to provide steam for powering a conventional prime mover driving an electricity generator. The hot water can be extracted from these hydrothermal reservoirs using boreholes and, after the heat has been extracted, the cooled water is pumped back into the ground to maintain the water table and pressure. Energy from geothermal aquifers is not completely renewable since heat is usually extracted at a rate quicker than it is replenished by the surrounding rocks. 3.2- Hot Dry Rocks (HDR) Enhanced Geothermal Systems (EGS) Hot rock systems extract energy from dry rocks with temperatures up to 1000°C deep in the Earth's crust, rather than from hydrothermal aquifers, but first the solid rock must be made permeable to allow the circulation of water into which the rocks give up some of their heat. Such Hot Dry Rock (HDR) systems (Fig., 40) need Enhanced Geothermal Systems (EGS) to extract the available energy and these involve much higher investments and exploration risks than extracting energy from naturally occurring hydrothermal reservoirs (Pruess 2007). Like hydrothermal systems, practical HDR systems depend on particular natural geological formations. They need access to hot granite or similar rocks with temperatures of 250°C or more, maintained by the heat flow from the Earth's hot core and such high temperatures are normally found at depths of over 3 kms. The deeper the rock, the higher the temperature but current drilling technology limits the practical working depths to about 5 57   

kms. The ideal geological formation also includes

an

insulating

blanket

of

sedimentary rocks, particularly shales, siltstones and coal seams, on top of the hot granite which effectively entrap the heat from the granite preventing it from being dissipated. Water is used as the thermal fluid to get the heat out of the rock and to enable this, the solid granite must be broken up (fractured) to allow horizontal water flow through the hot rock layer, and equally important, to provide the largest possible surface area of the hot rock through which the heat can be transferred into the water (Kitsou et al., 2000).

Fig., (40) Geothermal Energy Capture from Hot Rocks, Australian National University (Modified by Geothermal Resources Ltd) Source: http://www.mpoweruk.com /geothermal_energy.htm.

The water circulation system needs at least two bore holes, an injection bore hole through which cold water is pumped at high pressure down into the hot rock layer and an extraction borehole through which the hot water is returned to the surface. The fracturing of the hot rock is achieved by the injection of

water

from

the

surface

under

extremely high pressures. The water pressure forces open existing fractures in the hot rock, which do not completely

Fig., (41) The diagram shows the temperature gradient in the Earth's crust at different locations. Source: http://www.mpoweruk.com/geothermal_energy.htm

close again when the water pressure is removed, creating a passage through the rock between the injection and extraction boreholes. This is not an easy process because the immense pressures due to the weight of the overlying rocks tends close up any gaps in the rock. Nevertheless this EGS hydro-fracturing stimulation technology is commonly used in the oil industry to improve flow rates by enhancing the permeabilities of the host rock. The diagram below shows the main components of a 58   

geothermal power plant used to capture energy from hot dry rocks. The temperature profile varies, depending on factors such as the porosity of the rock, the degree of liquid saturation of the rock and sediments, their thermal conductivity, their heat storage capacity and the vicinity of magma chambers or heated underground reservoirs of liquid. 3.3- Geothermal energy in contemporary balneotherapeutics and Tourism In many countries, bathing and swimming are important and attractive aspects of geothermal direct uses. Geothermal is utilized in this way in at least 51 countries, i.e. over 11% of total installed power and 22% of thermal energy for direct uses worldwide (Fig., 41). Nowadays, recreation and healing based on geothermal water, steam, and energy are a very attractive and perspective branch of tourism where the demand exceeds the supplies. Geothermal plays a number of functions in tourism, e.g. swimming and therapeutic pools, curative geothermal by-products (e.g. salts), ecological heating of hotels and spas. Hydrothermal phenomena themselves (warm springs, geysers, hydrothermal minerals, etc.) are tourist attractions, similar to the historical objects or ruins related with geothermal use (Antics and Sanner 2007). Incorporation of these phenomena and objects in the common domain of tourism favours the idea of “sustainable development” and pro-ecological development of many regions and countries (Kępińska, 2004). 4- Healing and therapeutic value of geothermal waters Generally, cold mineral and geothermal waters can be treated as “therapeutic” or “having healing properties” if they meet at least one of the following criteria: 1) chemical (chemical composition); and 2) physical (temperature, radioactivity). Both these criteria are met by geothermal waters which can, owing to their physical (over 20⁰C) and chemical properties, naturally play healing or therapeutic functions (Antics and Sanner 2007; Kępińska, 2004). Temperature is one of the main factors thanks to which geothermal waters (just like regular mineral waters heated to a proper temperature) are applicable to healing, rehabilitation, and prophylaxy of diseases and dysfunctions of muscles, rheumatism, neurological diseases and many other ailments. Chemical composition greatly determines the application of geothermal waters for a spectrum of skin and internal diseases (Antics and Sanner 2007; Kępińska, 2004). Geothermal waters are also used for the production of therapeutic salt, leaches and evaporated salt. The total dissolved solids of such waters cannot exceed 60 g/dm3 and pharmacological-dynamic factors are taken into account. These minimum concentrations of chemical components dissolved in water or physical properties of water make up a threshold 59   

for biologically active waters. Therapeutic waters cannot be contaminated with bacteria or chemical compounds. Their curative properties must be proven by tests, and the oscillations in chemical composition and physical properties of waters may change only in a very small range (Antics and Sanner 2007; Kępińska, 2004). 4.1- Therapeutic tourism Geothermal balneotherapy and spas are basic elements of therapeutic tourism, one of the most important forms of recreation nowadays. Healing purposes can be acquired through various forms of tourism (spas, weekend tours, general healing tours, healing tours dedicated to specific diseases, etc.) Today, therapy is one of the fundamental functions of tourism, thanks to which the negative effects of civilization, e.g. stress can be reduced, and the inner force and feeling of integration reinforced (Kępińska, 2004). Over the centuries, these purposes have been most successfully realized in health resorts, i.e. spas, especially those with geothermal water. Spas are also attributed to a specific lifestyle, leisure, healing and biological rejuvenation, and an aspect of cultural and social life (Kępińska, 2004). 4.2- Geothermal Electricity Production around the world  Many regions of the world are already tapping geothermal energy as an affordable and sustainable solution to reducing dependence on fossil fuels, and the global warming and public health risks that result from their use. For example, more than 8,900 megawatts (MW) of large, utility-scale geothermal capacity in 24 countries now produce enough electricity to meet the annual needs of nearly 12 million typical

U.S.

households

(GEA

Fig., (42) Three different systems applied in Geothermal Electricity production. Source: http://www.ucsusa.org/clean_energy/our-energychoices/renewable-energy/how-geothermalenergy-works.html

2008a).

60   

Geothermal plants produce 25 percent or more of electricity in the Philippines, Iceland, and El Salvador (Fig., 42). The United States has more geothermal capacity than any other country, with more than 3,000 megawatts in eight states. Eighty percent of this capacity is in California, where more than 40 geothermal plants provide nearly 5 percent of the state’s electricity.1 In thousands of homes and buildings across the United States, geothermal heat pumps also use the steady temperatures just underground to heat and cool buildings, cleanly and inexpensively. The largest geothermal system now in operation is a steam-driven plant in an area called the Geysers, north of San Francisco, California. Despite the name, there are actually no geysers there, and the heat that is used for energy is all steam, not hot water. Although the area was known for its hot springs as far back as the mid-1800s, the first well for power production was drilled in 1924. Deeper wells were drilled in the 1950s, but real development didn't occur until the 1970s and 1980s. By 1990, 26 power plants had been built, for a capacity of more than 2,000 MW. (Source: http://www.ucsusa.org/clean_energy/our-energychoices/renewable-energy/how-geothermal-energy-works.html). Geothermal energy supplies more than 10,000 MW to 24 countries worldwide and now produces enough electricity to meet the needs of 60 million people. The Philippines, which generates 23% of its electricity from geothermal energy, is the world's second biggest producer behind the U.S. Geothermal energy has helped developing countries such as Indonesia, the Philippines, Guatemala, Costa Rica, and Mexico. The benefits of geothermal projects can preserve the cleanliness of developing countries seeking energy and economic independence, and it can provide a local source of electricity in remote locations, thus raising the quality of life. Iceland has been expanding its geothermal power production largely to meet growing industrial and commercial energy demand. In 2004, Iceland was reported to have generated 1465 gigawatt-hours (GWh) from geothermal resources; geothermal production is expected to reach 3000 GWh at end of 2009 (Blodgett and Slack, 2009).

61   

CHAPTER III MINERAL RESOURCES OF THE SEDIMENTARY BASINS The diversified geology of various regions and stratigraphic levels within the basins have given rise to a wide variety of minerals, more than 50 different kinds other than oil, gas and coal, that have an existing or potential resource value. The minerals are divided into industrial (or nonmetallic) minerals and metallic minerals. Under these broad categories the minerals are grouped into the various mineral types shown, with each type having common geological characteristics or elemental associations or both. With respect to the origin of basin fluids, Lawrence and Cornford (1995) distinguish between internally derived fluids such as formation waters (connate waters) and hydrocarbons, and externally derived fluids such as meteoric and metamorphic fluids. This hydrothermal circulation also extracts minerals and salts from rock. Minerals precipitate out of the hot waters and build spectacular vents, tens of meters high, on the mid-ocean ridges. Another internal source of fluid is related to clay diagenesis, which may contribute to overpressure build-up in subsiding basins (Bethke, 1986). Subsurface fluid flow plays a significant role in many geologic processes and is increasingly being studied in the scale of sedimentary basins and geologic time perspective. Many economic resources such as petroleum and mineral deposits are products of basin scale fluid flow operating over large periods of time. Such ancient flow systems can be studied through analysis of diagenetic alterations and fluid inclusions to constrain physical and chemical conditions of fluids and rocks during their paleohydrogeologic evolution. Basin simulation models are useful to complement the paleohydrogeologic record preserved in the rocks and to derive conceptual models on hydraulic basin evolution and generation of economic resources. Different types of fluid flow regimes may evolve during basin evolution (Bitzer et al., 2001). The most important with respect to flow rates and capacity for transport of solutes and thermal energy is gravitational fluid flow driven by the topographic configuration of a basin. Such flow systems require the basin to be elevated above sea level. Consolidational fluid flow is the principal fluid migration process in basins below sea level, caused by loading of compressible rocks. Flow rates of such systems are several orders of magnitude below topography driven flow. However, consolidation may create significant fluid overpressure. Episodic dewatering of over-pressured compartments may cause sudden fluid release with elevated flow velocities and may cause a transient local thermal and chemical disequilibrium between fluid and rock. This paper gives an overview on subsurface 62   

fluid flow processes at basin scale and presents examples related to the Penedès basin in the central Catalan continental margin including the offshore Barcelona half-graben and the compressive South-Pyrenean basin (Bitzer et al., 2001). I- Organic Mineral Resources I.1- Oil and Natural Gas Resources The world was divided into 8 regions and 937 geologic provinces. These provinces have been ranked according to the discovered known oil and gas volumes (Klett et al., 1997). Then, 76 “priority” provinces (exclusive of the United States and chosen for their high ranking) and 26 “boutique” provinces (exclusive of the United States) were selected for appraisal of oil and gas resources. Boutique provinces were chosen for their anticipated petroleum richness or special regional economic or strategic importance (Klett, 2000). A geologic province is an area having characteristic dimensions of hundreds of kilometers that encompasses a natural geologic entity (for example, a sedimentary basin, thrust belt, or accreted terrane) or some combination of contiguous geologic entities. Each geologic province is a spatial entity with common geologic attributes. Province boundaries were drawn as logically as possible along natural geologic boundaries, although in some places they were located arbitrarily (for example, along specific water-depth contours in the open oceans) (Klett, 2000). Total petroleum systems and assessment units were delineated for each geologic province considered for assessment. It is not necessary for the boundaries of total petroleum systems and assessment units to be entirely contained within a geologic province. Particular emphasis is placed on the similarities of petroleum fluids within total petroleum systems, unlike geologic provinces and plays in which similarities of rocks are emphasized (Klett, 2000). The total petroleum system includes all genetically related petroleum that occurs in shows and accumulations (discovered and undiscovered) generated by a pod or by closely related pods of mature source rock. Total petroleum systems exist within a limited mappable geologic space, together with the essential mappable geologic elements (source, reservoir, seal, and overburden rocks). These essential geologic elements control the fundamental processes of generation, expulsion, migration, entrapment, and preservation of petroleum within the total petroleum system (Klett, 2000). 1.1- Sedimentary basins and petroleum formation in the Middle East Middle East is divided into three major sedimentary basins: the Greater Arabian Basin, the Zagros Basin and the Oman Basin. Each basin is further divided into sub-basins, and each of these has its own style and time of origin reflected by differences in thickness and 63   

lithology.

The

megatectonic

framework of the Middle East (Alsharhan and Nairn 2003) shows that the area is dominated by the many sub-basins, broad regional highs,

anticlines

reflecting

and

deep-seated

flexures basement

faults and salt diapirisms. From the early

Mesozoic

onwards,

the

pattern of sedimentation in the Middle East was influenced by periods

of

increased

activity

alternating with quiet intervals. During the late Turonian to the early Campanian (Fig., 43), a major

Fig., (43) Paleogeographic map of the Albian–upper Cenomanian strata of the Arabian Gulf basin (modified from Murris, 1980).

change in basin configuration took place, heralding the first phase of Alpine compressive tectonics (Murris, 1980). During the Late Cretaceous orogenic period in Syria, northwestern Iraq and Southeast Turkey, dextral and sinistral strikeslip faults, fault zones and grabens were formed. The grabens were filled by a thick sequence of Sediments, which were inverted during the late Tertiary compressive phase, giving rise to en-echelon fold belts. The only areas in the Middle East with production and potential approaching that of the Middle East are in the Pricaspian Basin and the West Siberian Basin of the former USSR. Saudi Arabia ranks second in proven reserves and first in exporting oil, replacing the former Soviet Union with its rapidly declining production. Exploration in the producing areas of the Arabian Gulf and in the Zagros generally is in the mature phase; after many years of increasing reserve estimates, the figures are beginning to decline, despite the dramatic increase in reserve estimates of gas. However, there still are major untested areas, particularly in Iraq, Jordan and Yemen, and new play concepts and the introduction of new technologies may reverse the decline, at least temporarily (Powers, et al., 1985). The Albian–Cenomanian consists mainly of Orbitolina-bearing limestone with local basin margin rudist buildups in the offshore North field of Qatar and northeast Iraq. There are two main oil provinces where the Mauddud Formation is a major oil-producing reservoir. The Northern Province includes Iraq’s oil fields such as Ain Zalah, Bai Hassan, and Jambur. The 64   

southern province includes the Ratawi field in southern Iraq, Raudhatain, Sabriya, and Bahra fields in Kuwait, Bahrain (Awali) field in Bahrain, and Fahud and Natih fields in Oman. The formation has high oil potential in the southern and southeastern fields of Iraq and the offshore areas of Qatar and Saudi Arabia (Sadooni and Alsharhan 2003). 1.2- Petroleum prospectivity of the principal sedimentary basins on the United Kingdom Continental Shelf The main sedimentary basins within the UKCS can be broadly divided into a number of separate provinces, on the basis of petroleum geology and location. These provinces comprise the North Sea Oil Province, the North Sea Gas Province, the Irish Sea and the Atlantic Margin (Fig., 44). The remaining and future petroleum potential of these provinces is summarised below. The following six exploration plays, in particular are anticipated to offer significant hydrocarbon potential (Munns et al., 2005): Upper Jurassic syn-rift deep-water play, Upper Jurassic shallow-marine ‘inter-pod’ play, Lower Cretaceous deep-water play, Paleogene deep-water play, Upper Cretaceous Chalk play, and Lower Permian basin-margin play (Gray, 2010). The North Sea Oil Province is one of the world’s major oil-producing regions. The geological history of the oil province was dominated by an episode of late Jurassic to earliest Cretaceous

crustal

extension,

which

developed the Viking Graben, Moray Firth and Central Graben rift systems. Syn-rift, organic-rich

marine

mudstones

(Kimmeridge Clay Formation) are the source rocks for virtually all of the region’s hydrocarbons. Post-rift thermal subsidence enabled these source rocks to become

mature

for

hydrocarbon

generation along the rift axes from Paleogene times onwards (Johnson and Fisher, 1998). Hydrocarbon migration has been mainly vertical. Consequently, most of the producing oil and gas fields lie

Fig. (44). Distribution of oil and gas provinces and petroleum Carboniferous source rocks on the UK Continental Shelf (After Gary, 2010).

within the geographical boundary of the mature source rocks. Hydrocarbons occur in a wide range of pre-rift, syn-rift and post-rift reservoirs (Gray, 2010). Although extensional rifting generally ceased during the earliest Cretaceous (Ryazanian), fault-controlled subsidence 65   

persisted in parts of the Moray Firth Basin, throughout much of Early Cretaceous times. This localised tectonism is considered by Oakman and Partington (1998) to have been controlled by strike-slip faulting. 1.3- prospectivity of the sedimentary basins of Irish Sea The East Irish Sea Basin is at a mature

exploration

phase.

Early

Namurian basinal mudstones are the source rocks for these hydrocarbons. Production from all fields is from faultbounded

traps

of

Lower

Triassic,

principally aeolian Sherwood Sandstone reservoir, top-sealed by younger Triassic continental mudstones and evaporates (Fig.,

45).

Future

exploration

Fig. (45). Hydrocarbon fields and discoveries.

will

initially concentrate on extending this play, but there remains largely untested potential also for gas and oil within widespread Carboniferous fluvial sandstone reservoirs. This play requires intraformational mudstone seal units to be present, as there is no top-seal for reservoirs subcropping the regional base Permian unconformity in the east of the basin, and Carboniferous strata crop out at the sea bed in the west (Gray, 2010). I. 2- Coal bearing formations: 2.1- Australia: The Australian coals can be separated into Permian, Mesozoic and tertiary coals. The Permian coals are in general hard coal between high volatile bituminous and anthracite rank, the Mesozoic coals are high volatile bituminous, perhydrous coals and the tertiary coals are of the lignitic rank. The coals of the Gippsland Basin are the youngest coals that can be extracted by mining. The coal has five major coal seams, which are defined in Yallourn, Morwell and Traralgon Formations. These three formations are all not older than 5 Ma; the individual seam thickness often exceeds the 100 m. The vertical stratigraphic position is over 400 m of continuous low ash coal (Crosdale, 2004; Holdgate, 2005; Golab et al., 2007). 2.2- India: Indian coal resources are confined to two distinct geological periods and basinal set-ups – (i) Permian sediments deposited mostly in the intra-cratonic Gondwana basins of Peninsular India and a few minor occurrences as thrust sheets (overriding Siwalik sediments) 66   

in the foothills of Darjeeling and Arunachal Pradesh Himalayas and (ii) Early Tertiary sediments deposited in the near shore peri-cratonic basins and shelves in the North Eastern Region. More than 99% of total coal resources are Gondwana coal. The Gondwana basins of Peninsular India, restricted to the eastern and central parts of the country are disposed as linear belts along the course of major river valleys of Damodor-Koel (with a subsidiary belt to the north). Extensive spread of Gondwana sediments beneath the Ganga and Brahmaputra alluvium in the Bengal Basin and beneath the Deccan traps in Central Indian craton, in addition to the occurrences of Gondwana outliers beyond the confines of known coalfields, is suggestive of a much wider span of the parent Gondwana basins. Gondwana sediments are represented by thick sequence of glacial, marine, fluvial and lacustrine facies (Bhattacharya, 2011). II- Inorganic Mineral Resources II.1- Volcanogenic massive sulphides (VMS) deposits are also known as volcanicassociated, volcanic-hosted, and volcano-sedimentary-hosted massive sulphide deposits. They typically occur as lenses of polymetallic massive sulphide that form at or near the seafloor in submarine volcanic environments. They form from metal-enriched fluids associated with seafloor hydrothermal convection. Their immediate host rocks can be either volcanic or sedimentary. VMS deposits are major sources of Zn, Cu, Pb, Ag and Au, and significant sources for Co, Sn, Se, Mn, Cd, In, Bi, Te, Ga and Ge. Some also contain significant amounts of As, Sb and Hg. Historically, they account for 27% of Canada's Cu production, 49% of its Zn, 20% of its Pb, 40% of its Ag and 3% of its Au. Because of their polymetallic content, VMS deposits continue to be one of the best deposit types for security against fluctuating prices of different metals. There are close to 800 known VMS deposits worldwide with geological reserves over 200,000 t (Galley et al., 2007). II.2- Metaliferous Oxides: Manganese nodules occur in all the oceans. Their accretion rate is very slow, only a few mm in 1 million years. The average nodule has 24% manganese, compared to 35 to 55% manganese in land ore bodies, so they do not offer solid economics as a manganese source, but they also contain iron (14%), copper (1%), nickel (1%), and cobalt (0.25%). II.3- Metallic and Gem Minerals in Placer Deposits: A placer deposit is an accumulation of mineral grains concentrated by sedimentary processes. When pebbles, sands, and silts are sorted by wave action or stream flow, minerals with higher specific gravity and resistance to weathering become concentrated, especially in beaches and drowned river mouths. Marine 67   

placer mineral deposits are found on the continental shelf from the beaches to the outer shelf. The strategic element titanium derived mainly from ilmenite and rutile ores, while the noble elements are gold and platinum. II.4- Evolution of a Mineralized Geothermal System, Valles Caldera, New Mexico, USA: Hot springs and fumaroles are surface manifestations of a hydrothermal reservoir (210°–300°C; 2–10 x 103 mg/kg Cl) that is most extensive in fractured, intracaldera Bandelier Tuff and associated sedimentary rocks, located in specific structural zones. Fluids are composed of deeply circulating water of (primarily) meteoric origin, which have a mean residence time in the reservoir of 3–10 kyr. Host rocks show intense isotopic exchange with hydrothermal fluids. Alteration assemblages are controlled by temperature, permeability, fluid composition, host‐rock type and depth. A generalized distribution from top to bottom of the system consists of argillic, phyllic, propylitic, and calc‐silicate mineral assemblages. Typical alteration minerals in phyllic and propylitic zones are quartz, calcite, illite, chlorite, epidote and pyrite, whereas common vein constituents consist of the above minerals plus fluorite, adularia, and wairakite. Argentiferous pyrite, pyrargyrite, molybdenite, sphalerite, galena, chalcopyrite, arsenopyrite, stibnite and barite have been found at various depths and locations in the Valles system (Goff, 2012). II.5- Mineral Resources of the Western Canada Sedimentary Basin Minerals other than oil, gas and coal occur in abundance and variety in the Western Canada Sedimentary Basin. They include the industrial (or nonmetallic) and metallic minerals and together account for a significant proportion of Western Canada's wealth. Metallic minerals are much less developed; known deposits are few and generally small, although they include the world-class Pine Point (Pb-Zn) ore-body. For the industrial minerals, most production comes from the Interior Plains region, where Phanerozoic rocks form a northeast-tapering wedge of undeformed strata. These strata include Paleozoic carbonates and evaporites that give rise to rich resources of sulphur, potash, salt, gypsum, limestone and dolomite. The Paleozoic strata are succeeded by Mesozoic and Tertiary clastic rocks that are sources for economic deposits of kaolin and structural clays, bentonite, silica sand, and constructional sands and gravels. Important production also comes from the Cordilleran region, where deformed and upthrusted basin strata in the Rocky Mountain belt expose economic deposits of limestone, magnesite, gypsum and quartzite. For the metallic minerals, except for Pine Point, most deposits have been found in the Cordilleran region. These are mainly lead-zinc deposits of the Mississippi Valley type, few in 68   

number and widely separated; limited past production came from small localized orebodies in southeastern British Columbia. In the Interior Plains, the Pine Point lead-zinc deposit is the largest and only significant economic deposit. Some placer gold is still produced from Tertiary and recent gravels. Sedimentary iron deposits (in the Clear Hills region of Alberta) are large, but remain undeveloped (Hamilton and Olson, 1990). II.6- Mineral Resources of the Australian Sedimentary Basins 6.1- Heavy minerals: the Murray Basin, much of which is in New South Wales, has the potential to become one of the world’s major new mineral sands provinces. Total resources of coarse-grained heavy minerals (rutile, zircon, ilmenite and weathered ilmenite) identified in the Murray Basin exceed 100 Mt, of which over 80 Mt occurs in the New South Wales part of the basin. Beach placers along much of the coast north of Sydney were formerly major sources of rutile, zircon and ilmenite. These heavy minerals form as accessory minerals in many igneous and metamorphic rocks, nearly all major economic deposits of these minerals, principally rutile, zircon and ilmenite, occur as detrital accumulations in young (Pliocene or younger) shoreline or beach placer deposits (Force 1991; Roy and Whitehouse 2003). 6.2- Bauxite: Australia is the largest producer of bauxite in the world. New South Wales has numerous comparatively small, scattered deposits of bauxite/laterite that typically occur as discontinuous deposits along the crests of flat hills, ridges and in areas of generally subdued topography (Holmes et al. 1982). 6.3- Sedimentary phosphate deposits occur on every continent and range in age from Precambrian to Recent. Large resources of phosphates occur on the continental shelves. Phosphorite beds consist of grains, pellets or fragments of cryptocrystalline apatite (collophane) and are typically a few centimeters to tens of meters thick (McHaffie and Buckley 1995). These deposits typically show extensive reworking, secondary enrichment and replacement. Shallow oceanic areas and continental shelves commonly have thick accumulations of phosphorus-rich organic debris, mainly derived from deep oceanic sources associated with upwelling currents of cold, nutrient-rich water. The Phosphate Hill deposit is the only commercial phosphate rock mine in Australia. The southern Eromanga Basin (Great Australian Basin) may have potential for economic phosphate rock deposits (Wallis, 2004). 6.4- Other Metals: Gold, Copper, Silver, Uranium, Gypsum, Iron oxide, Kaolin, Limestone, Magnesite,, Magnetite , Manganese, Mica, Olivine, Opal, Asbestos and Dolomite (McHaffie and Buckley 1995).   69   

CHAPTER IV MEDITERRANEAN SEA The Mediterranean Sea is a mid-latitude semi-enclosed sea, or almost isolated oceanic system. Many processes which are fundamental to the general circulation of the world ocean also occur within the Mediterranean, either identically or analogously. The Mediterranean Sea exchanges water, salt, heat, and other properties with the North Atlantic Ocean. The North Atlantic is known to play an important role in the global thermohaline circulation, as the major site of deep- and bottom-water formation for the global thermohaline cell (conveyor belt) which encompasses the Atlantic, Southern, Indian, and Pacific Oceans. The salty water of Mediterranean origin may affect water formation processes and variabilities and even the stability of the global thermohaline equilibrium state (Robinson et al., 2001). I- Mediterranean Geosynclinal Belt Mediterranean Sea is one of the largest mobile regions of the earth’s crust, separating the Eastern European, Siberian, Sino-Korean, and South China platforms from the AfricanArabian and Indian platforms. The Mediterranean geosynclinal belt stretches across Eurasia (Europe-Asia), from the Strait of Gibraltar in the west to the Indonesian archipelago, where it joins the Pacific geosynclinal belt. It encompasses a large part of Western and Southern Europe, the Mediterranean Sea, North Africa (Morocco, Algeria, and Tunisia), and Southwest Asia. Geosynclines are all down-warped and down-faulted basins within the craton exclude thick continental terrace-type deposits some of which have been designated geosynclines as the Gulf Coast (Pettijohan, 1984). They are characterized by; 1) location marginal to or between cratons, 2) mobility expressed by intense folding and thrusting, 3) initial somatic igneous phase marked by ophiolitec emission I internal (away of the craton) zones and, 4) synorogenic and post-orogenic igneous activity in internal zones (Pettijohan, 1984). The Mediterranean geosynclinal belt includes the Hercynian (Variscan) folded regions of Western and Central Europe, the Alpide geosynclinal (folded) region, and the Indonesian folded region. The vast Tethys Sea was located on the site of the Mediterranean geosynclinal belt during the Paleozoic and Mesozoic. The Mediterranean was once a deep, dry valley, some five million years ago, dividing the three continents, Europe, Africa and Asia, until a cataclysmic broach was made in the retaining wall, which kept out the Atlantic Ocean in the West, towards present-day Gibraltar. 70   

A huge cascade of water began, flooding the whole Mediterranean basin, in a process that lasted many, many years, and a new sea was born. Analysing the geographical configuration of this new sea more closely, we find that it is rather formed from a number of seas: the Alboran, the gulf of Lione, the Tirrhennian Sea, the Ionian Sea, the Aegean Sea, the Adriatic Sea etc, each with its own characteristics (Pettijohan, 1984; Robertson and Mountrakis, 2006). II- Origin and evolution of Mediterranean geosyncline The Eastern Mediterranean is one of the key regions for the understanding of fundamental tectonic processes, including continental rifting, passive margins, ophiolites, subduction, accretion, collision and post-collisional exhumation. It is also ideal for understanding the interaction of tectonic, sedimentary, igneous and metamorphic processes through time that eventually lead to the development of an orogenic belt. Tethyan nomenclature remains controversial and we will suggest an appropriate informal terminology for the various oceanic basins that existed. This envisaged southward subduction of a Late Palaeozoic-Early Mesozoic ocean (Palaeo-Tethys) and the related opening of several marginal basins along the northern margin of Gondwana. Closure of this ocean culminated in continental collision by the latest Triassic-Early Jurassic time, and was followed by opening of a new, Jurassic ocean basin (Northern Neotethys) (Robertson and Mountrakis, 2006). Geologic features in the present-day Mediterranean essentially result from two major processes: the tectonic displacement caused by the subduction of the African plate underneath the Eurasian plate; and the progressive closure of the Mediterranean Sea involving a series of submarine-insular sills. The development of the Mediterranean basin begins with the breakup of the supercontinent Pangea in the Mesozoic Era. During this time, sea-floor spreading triggered the development of the Atlantic Ocean in the Triassic period, which separated the African and Eurasian plates from the North American plate. Sea-floor spreading in another geographical location caused the development of the Tethys Ocean, separating the African plate from the Eurasian. In the late Cretaceous period, these African and Eurasion plates began to converge, closing the Tethys ocean basin, and the remnants of this ancient ocean (Smith 1993; Dercourt et al., 2000; Robertson and Mountrakis, 2006). In the Cainozoic age, the area of the Mediterranean Sea was a huge ocean that slowly shrank into a few secondary basins. The main one then turned into the Mediterranean Sea. This was caused by the African and Eurasian continental plate moving closer to each other. 71   

The powerful thrusts coming from the south caused the sediments built up at the bottom of the ocean to raise, thus originating the mountain ridges of the Atlantis, the Pyrenees, the Alps, the Balkans and Asia Minor. During the late Miocene, the ancient ocean became an internal sea, even if different from today’s Mediterranean Sea (Robertson and Mountrakis, 2006). During the Pliocene, the Mediterranean Sea dried up. The geological phenomena associated with this period, such as the opening of huge fractures, volcanic activity, the raising of coastal areas, etc., prompted the formation of the ecological and geographical complexity of the Mediterranean region. This phase boosted the expansion of salt-resistant plants (Halophytes of the genera: Limonium, Salicornia, Arthrocnemum, Salsola, Artemisia) and the appearance of small and sparse species whose adaptability to particular conditions made them develop quickly. In the end, today’s Straits of Gibraltar broke up because of the earth's crust moving, and the water of Atlantic sea flew into the Mediterranean basin. The current configuration of this basin came into being approximately five million years ago (Dercourt et al., 2000; Robertson and Mountrakis, 2006). There are three major geomorphical settings within the Mediterranean basin; areas with stable margin characteristics, areas with unstable convergent margin characteristics, and areas with extensional margin (rifting) characteristics. Thus the Mediterranean basin is a location of an intercontinental interplate system; with compressional and extensional events occurring within close proximity (Robertson and Mountrakis, 2006). Geologists have yet to come to a consensus about which plates in addition to the African and Eurasian ones, if any, are involved in Mediterranean tectonics. Subsidence-related and other vertical displacements are also found in compressional and extensional areas. A few notable events occurred during the Cenozoic which affected the entire Mediterranean; the Messinian "salinity crisis", when the closing off of the Mediterranean-Atlantic seaway caused complete isolation of the Mediterranean and thus widespread evaporation; and then the Pliocene "revolution", when the channel opened back up, causing reestablishment of marine conditions; and the Quaternary "transgressive raised terraces," of controversial geological origin; among others (Dercourt et al., 1986; 1993; 2000). The Central portion of the Mediterranean basin exemplifies the juxtaposition of compressional and extensional tectonic activity in the area. The region bordered to the west by Sicily and to the east by Turkey's west coast (encompassing the Aegean, Ionian, and Adriatic seas) exhibit a particular set of features. There were four major periods of extension 72   

in this area. The first one occurred in the Mid-Upper Jurassic; evidence of this phase is seen in the Strepanosa Trough and Ionian plain. A second one occurred in the Mid-Late Triassic, opening up the Ionian Sea and the Eastern Mediterranean. A third extensional phase occurred in the Mid-Upper Cretaceous, as evidenced by the stretched features of the Sirte Rise, a monocline with normal faults and tilted blocks. The fourth one, occurring in the Mid-Upper Miocene through to the Quaternary period, affected many areas of the Central Mediterranean (Dercourt et al., 2000). This extensional phase is closely associated with compressive motions; it is part of the reason for a counter-clockwise rotation of the Southern Appennine area which begins in the upper Cretaceous. All four of these extensional phases are the cause of geologic features found in the area, such as volcanic activity and rift-related sedimentary processes. Due to such extension, the oceanic crusts of the Central Mediterranean are considerably thinned in some places. The opening of small oceanic basins of the central Mediterranean follows a trench migration and back-arc opening process that occurred during the last 30 Myr. This phase was characterized by the counterclockwise rotation of the Corsica-Sardinia block, which lasted until the Langhian (ca.16 Ma), and was in turn followed by a slab detachment along the northern African margin. Subsequently, a shift of this active extentional deformation led to the opening of the Tyrrenian basin (Dercourt et al., 1986; 1993; 2000). The Mediterranean Ridge or Outer Median Ridge is a sea-floor feature that marks the unstable (convergent) margin between two or more oceanic plates. The first stages of the major collision between the North of the African plate and the South of the Eurasion plate are believed to have occurred in the lower-middle Miocene (Dercourt et al., 1993; 2000). This collision is also associated with the counter-clockwise rotation of the Appennine area, and both of these associations are exhibited in the Calabrian (Italy and Sicily) and Hellenic (Greece) orogenic arcs which are situated among both compressive and extensional dynamics. The ridge extends geographically from Sicily to Cyprus along a generally E/W strike. It is an extensive fold-fault system corresponding to recent uplift and folding of past abyssal plains (Smith, 1993; Dercourt et al., 2000). III- Paleoenvironmental analysis Its semi-enclosed configuration makes the oceanic gateways critical in controlling circulation and environmental evolution in the Mediterranean Sea. Water circulation patterns are driven by a number of interactive factors, such as climate and bathymetry, which can lead 73   

to precipitation of evaporites. During late Miocene times, a so-called "Messinian Salinity Crisis" (MSC hereafter) occurred, which was triggered by the closure of the Atlantic gateway. Evaporites accumulated in the Red Sea Basin (late Miocene), in the Carpatian foredeep (middle Miocene) and in the whole Mediterranean area (Messinian) (Cendon et al., 2004). An accurate age estimate of the MSC—5.96 Ma—has recently been astronomically achieved; furthermore, this event seems to have occurred synchronously. The beginning of the MSC is supposed to have been of tectonic origin; however, an astronomical control (eccentricity) might also have been involved. In the Mediterranean basin, diatomites are regularly found underneath the evaporitic deposits, thus suggesting (albeit not clearly so far) a connection between their geneses. The present-day Atlantic gateway, i.e. the Strait of Gibraltar, finds its origin in the early Pliocene. However, two other connections between the Atlantic Ocean and the Mediterranean Sea existed in the past: the Betic Corridor (southern Spain) and the Rifian Corridor (northern Morocco). The former closed during Tortonian times, thus providing a "Tortonian Salinity Crisis" well before the MSC; the latter closed about 6 Ma, allowing exchanges in the mammal fauna between Africa and Europe. Nowadays, evaporation is more relevant than the water yield supplied by riverine water and precipitation, so that salinity in the Mediterranean is higher than in the Atlantic. These conditions result in the outflow of warm saline Mediterranean deep water across Gibraltar, which is in turn counterbalanced by an

inflow

of

a

less

saline

surface

current

of

cold

oceanic

water

(Source:

http://www.princeton.edu/~achaney/tmve/wiki100k/docs/Mediterranean_Sea.html).

IV- Mediterranean basins The Mediterranean, seen from the surface, looks like a single sea divided into a number of basins with different characteristics, a different geological history, and also different morphologies of the sea floor. The main division is that of the Western Mediterranean and Eastern Mediterranean: two basins separated by an underwater ridge that crosses the sea from Sicily to the coasts of Tunisia. The morphological differences between the two basins also provoke differences in the temperature and in the chemical characteristics of the water. The western basin has a temperature of 12 °C in winter and 23 °C in summer, its salinity is 36‰ , while the eastern basin is warmer and more salty, it temperature is 16 °C in winter and 26-29°C in summer, its salinity is 39‰. The geological and morphological differences of the two basins also have consequences on the distribution of the living forms.

74   

The two main basins are in turn divided into smaller sub-basins, whose characteristics depend greatly on the geological history that led to their formation (Mazzoleni et al., 1992). IV. 1- Tectonic settings of Eastern Mediterranean basin

The Eastern Mediterranean is one of the key regions for understanding of fundamental tectonic processes, including continental rifting, passive margins, ophiolites, subduction, accretion, collision, and post-collisional exhumation. It is ideal for understanding the interaction of tectonic, sedimentary, igneous and metamorphic processes through time that eventually

lead

to

the

development of an orogenic belt

(Robertson

and

Mountrakis,

2006).

origin

the

of

The

Eastern

Mediterranean

basin

(EMB) by rifting along its passive

margins

is

reevaluated. Evidence from these margins shows that

Fig. (46). The East Mediterranean Basin (EMB) and the extent of its subducted parts (Garfunkel, 2004).

this basin formed before the Middle Jurassic; where the older history is known, formation by Triassic or even Permian rifting is indicated. Off Sicily, a deep Permian basin is recorded. In Mesozoic times, Adria was located next to the EMB and moved laterally along their common boundary, but there is no clear record of rifting or significant convergence (Fig., 46). Farther east, the Tauride block, a fragment of Africa–Arabia, separated from this continent in the Triassic. After that the Tauride block and Adria were separate units that drifted independently. The EMB originated before Pangaea disintegrated. Two scenarios are thus possible. If the configuration of Pangaea remained the same throughout its life span until the opening of the central Atlantic Ocean (configuration A), then much of the EMB is best explained as a result of separation of Adria from Africa in the Permian, but this basin was modified by later rifting. The Levant margin formed when the Tauride block was detached, but space limitations require this block to have also extended farther east. Better constraints on the history of Pangaea are thus required to decipher the formation of the Eastern Mediterranean basin (Garfunkel, 2004). In middle Miocene times, the collision between the Arabian microplate and Eurasia led to the separation 75   

between the Tethys and the Indian Oceans. This process

determined

profound changes in the oceanic patterns,

circulation which

shifted

global climates towards colder

conditions.

The

Hellenic Arc, which has a

Fig. (47). The East Mediterranean Basin (EMB) and the extent of its subducted parts (Garfunkel, 2004).

land-locked configuration, underwent a widespread extension for the last 20 Myr due to a slab roll-back process (Fig., 46). In addition, the Hellenic Arc experienced a rapid rotation phase during the Pleistocene, with a counterclockwise component in its eastern portion and a clockwise trend in the western segment. To the east and south, its original passive margins are preserved, whereas its present northern and western margins were shaped by later subduction and plate convergence. Seismic refraction studies show that the EMB has an up to 10 km thick probably oceanic crust (and/or strongly attenuated continental crust) overlain by 6 to >12 km of sediment (DeVoogd et al., 1992; Ben-Avraham et al., 2002). In contrast, the Africa–Arabia continent next to the passive margins of this basin has 30- to 35-km thick continental crust (Makris et al., 1988). Such a change in crustal structure allows the interpretation that the EMB formed as a result of rifting, which led to detachment and northward drifting of blocks away from these passive margins. This view is widely accepted, but the history of rifting, the identity and original location of the detached blocks, and the growth history of the basin remain incompletely understood (Robertson et al., 1996). IV.2- Tectonic Settings of the Western Mediterranean The western Mediterranean is the younger part of the Mediterranean, being a basin formed from late Oligocene to present. The western Mediterranean consists of a series of subbasins such as the Alboran, Valencia, Provençal, Algerian and Tyrrhenian seas. These basins have in general a triangular shape and they generally rejuvenate moving from west to east. They are partly floored by oceanic crust (Provençal and Algerian basins, and two smaller areas in the Tyrrhenian Sea). The remaining submarine part of the western Mediterranean basin is made of extensional and transtensional passive continental margins. The continental 76   

crust is composed of Paleozoic and prePaleozoic rocks deformed by the Caledonian and Variscan orogenic cycles (Carminati et al., 1998a, b). The geological evolution of the western

Mediterranean

exhibits

interactions

between

complicated

orogenic processes and widespread extensional tectonics. The region is located in a convergent plate margin separating Africa and Europe, and consists of marine basins – the Alboran Sea, the Algerian-Provençal Basin, the Valencia Trough, the Ligurian Sea and the Tyrrhenian Sea which formed as back-arc basins since the Oligocene. The

evolution

simultaneously

of

these

with

basins, ongoing

convergence of Africa with respect to Europe, has been the subject of numerous studies (e.g., Stanley and Wezel 1985, Durand et al. 1999). Widespread extension associated with the formation of these basins led to considerable thinning of the continental crust (i.e., in the Alboran Sea and the northern Tyrrhenian) or to the local initiation of sea floor spreading (i.e., in the southern Tyrrhenian and Provençal

Fig., (48) Western Meditteranean reconstruction through Oligocene (Rosenbaum, et al. 2002).

Basin). Furthermore, extensional tectonism in the western Mediterranean was coeval with orogenesis in the adjacent mountain chains of the Rif-Betic cordillera, the Maghrebides of northern Africa and Sicily, the Apennines, the Alps and the Dinarides (Malinverno and Ryan, 1986; Crespo-Blanc et al., 1994; Tricart et al., 1994; Cello et al., 1996; Azañón et al., 1997; Frizon de Lamotte et al., 2000; Faccenna et al., 2001). Since Mesozoic to Tertiary times, 77   

during convergence between Africa and Iberia, it developed the Betic-Rif mountain belts. Tectonic models for its evolution include: rapid motion of Alboran microplate, subduction zone and radial extentional collapse caused by convective removal of lithosferic mantle. The development of these intramontane Betic and Rif basins led to the onset of two marine gateways which were progressively closed during the late Miocene by an interplay of tectonic and glacio-eustatic processes (Fig., 48). The simultaneous formation of extensional basins together with thrusting and folding in adjacent mountain belts has led to several tectonic models that acknowledge the role of large-scale horizontal motions associated with the retreat of the subduction trench (hereafter termed subduction rollback) (Malinverno and Ryan, 1986; Royden, 1993a; Lonergan and White, 1997). These provide an explanation for the origin of allochthonous terranes, which drifted great distances to their present locations (e.g., Calabria). However, some issues are yet to be resolved and have been the subject of considerable debate. Different models have been proposed to explain the evolution of the Alboran Sea, namely, as a back-arc basin associated with a retreating slab (Lonergan and White, 1997), or as the result of an extensional collapse of thickened lithosphere (Platt and Vissers, 1989; Houseman, 1996). The evolution of the Tyrrhenian Sea is also controversial, with some fundamental problems in the current explanations of the evolution of this basin (Rosenbaum et al., 2002). According to Rosenbaum et al., (2002), the reconstruction shows that during Alpine orogenesis, a very wide zone in the interface between Africa and Europe underwent extension.

Extensional

tectonics was governed by rollback

of

subduction

zones

triggered

by

gravitational instability of old

and

dense

lithosphere.

oceanic Back-arc

extension occurred in the overriding plates as a result of slow convergence rates combined

with

rapid

Fig., (49). Tectonic sketch of the Western Mediterranean Region (modified from Barrier et al., 2004).

subduction rollback (Fig., 49). This mechanism can account for the evolution of the majority of the post-Oligocene extensional systems in the western Mediterranean. Moreover, extension led to drifting and rotations of continental terranes towards the retreating slabs in excess of 78   

100-800 km. These terranes - Corsica, Sardinia, the Balearic Islands, the Kabylies blocks, Calabria and the Rif-Betic - drifted as long as subduction rollback took place, and were eventually accreted to the adjacent continents. We conclude that large-scale horizontal motions associated with subduction rollback, back-arc extension and accretion of allochthonous terranes played a fundamental role during Alpine orogenesis. The formation of the Western Mediterranean back-arc Basin is related to the northwards convergence of the Nubia (Africa) Plate relative to the Eurasia Plate since the late Cretaceous (Olivet, 1996). The plate boundary between Nubia (Africa) and Eurasia plates is clearly delineated in the Eastern Atlantic Ocean then becomes diffuse in the Alboran Sea and the adjacent areas in Spain and Morocco. Along the North Algeria the earthquakes, with reverse or strike-slip focal mechanisms compatible with NW–SE compression, are localized along a 200 km large stripe from the coast to in-land (Mauffret, 2007) V- Origin and Tectonic History of Mediterranean Sub-basins V.1- The Levantine Basin The Eastern Mediterranean, and with, it the Levantine Basin, is a relic of the Mesozoic NeoTethys Ocean (Stampfli and Borel,

2002;

Garfunkel,

2004). The Levantine Basin is confined by the Israeli and the Egyptian coasts, Cyprus and the Eratosthenes Seamount. In the Miocene, the

so-called

‘Messinian

Salinity Crisis’ was initiated by the disconnection of the Mediterranean

to

the

Atlantic (Fig., 50). This was caused by a combination of tectonic uplift and sea level changes and led to a drop of sea level, a rise in salt

Fig., (50) Levantine Basin, surrounding and related basins in the Eastern Mediterranean. Source: http://my.opera.com/talatkm/blog /index.dml/tag/ renewable%20energy

concentration and finally to precipitation (e.g., Gradmann et al., 2005). 79   

The early evolution of the Levantine Basin in the Southeastern Mediterranean Sea is closely related to the history of the Neo-Tethys. Determining whether the crust in the basin is continental or oceanic is crucial for reconstruction of the Neo-Tethys opening and the position of its spreading axes (Netzeband et al., 2006). Whereas the continental character of the crust under the Eratosthenes Seamount and Cyprus is undisputed (Garfunkel, 1998; Robertson, 1998a), the nature of the crust underlying the Levantine Basin is still a matter of debate. According to these theories, the Eratosthenes Seamount was separated from the African margin in the Permian (Garfunkel, 1998) along with other continental fragments (BenAvraham and Ginzburg, 1990) and migrated northwards, where it presently collides with Cyprus (Robertson, 1998b) with an annual collision rate of approximately 1cm/year (Kempler and Garfunkel, 1994; Albarello et al., 1995). The basin has undergone significant subsidence for more than 100Ma (Almagor, 1993; Vidal et al., 2000), over 2km since Pliocene and is still subsiding (Tibor et al., 1992). A fold belt, which extends from the Western Desert of Egypt through Sinai into the Palmyra folds of Syria has been termed the Syrian Arc (Walley, 1998). The evolution of this regional compressional tectonic feature began in the Late Cretaceous and continued until the Early– Middle Miocene (Walley, 1998). Its evolution was related to the closure of the Neo-Tethys (Garfunkel, 1998, 2004). V.2- Aegean Sea basin The Aegean Sea and its surroundings regions comprise one of the most rapidly deforming parts of the Alpine-Himalayan mountain belt. Through the deformation of the belt as a whole is related to the northward movement of Africa, Arabia and India relative to Eurasia, the tectonics of the Aegean region itself is dominated by strike-slip and extensional motions (Jackson, 1994). The sedimentary basin of Aegean Sea was formed by lithosphere stretching (McKenzie, 1979). McKenzie and Jackson (1983), noticed that the association of thinned crust with high heat flow, normal faulting and subsidence. The oceanic lithosphere to the SW is subducted in the Hellenic trench beneath the continental lithosphere of the Aegean to the NE, leading to the formation of an inclined seismic zone with active volcanos above it (Jackson, 1994).

80   

The Aegean Sea region was an active tectonic region, it was shortening by a series of collisional events in the late Mesozoic and early Tertiary, which imparted a strong structural fabric in the form of folds, thrust faults, and sutures that trend NW-SE in mainland of Greece, and then change to more E-W or ENE-WNW orientation across the central Aegean and into western Turkey (Sengor et al., 1984). The Aegean region, located in the overriding plate of the Hellenic subduction zone, has been subjected to extensional tectonics since the late Eocene-Early Oligocene (~35 Ma) (Jolivet and Faccenna, 2000; Jolivet and Brun, 2010). Earlier extension may have occurred to the North in the Rhodope massif since some 45 Ma at a slower pace (Brun and Sokoutis, 2007). The Hellenides formed from the Late Jurassic to the Present above the Hellenic subduction

(Royden

Papanikolaou,

and 2011;

Philippon et al., 2012). They result from the off scrapping of crustal units from the Pelagonian in the north and then the Pindos Ocean and Apulian block further south that were subducted below Eurasia after the closure of the Vardar ocean in the Late Cretaceous (Fig., 51a). The Aegean

Sea

considerable

Fig., (51a). Sedimentary basin development of Aegean Sea: Three reconstructions of the section showing the progressive slab retreat, fter Jolivet and Brun (2010) and a velocity field of particles after the analogue model of Funiciello et al. (2003). Partially molten lower crust is shown in red After (Jolivet et al., 2012).

experiences amounts

of

extensional features as well, related to the subduction of the African plate underneath the

Hellenic

Arc.

The

shortening of the Pindos and Apulian blocks led to the formation of a series of large scale nappes, all emplaced

Fig., (51b) Tectonic map of Aegean-Anatolian Region (After (Jolivet et al., 2012).

with a south or southwest

81   

vergence of the thrust front, from the Eocene to the present (Sotiropoulos et al., 2003; Van Hinsbergen et al., 2005).The Aegean domain, since the Oligo-Miocene, in a geodynamic sense, also encompasses a part of western Anatolia (Fig., 51b). The Menderes massif has indeed recorded tectonic events that are typically Aegean and it is thus useful to review the evolution of ideas on this region as well. Moreover, the crust is thicker in the Menderes massif and the pre-extension structures are thus better preserved than in the Cyclades (Jolivet et al., 2012). Subsidence in the late Miocene also had a grand affect on the region, resulting in the fragmentation of an Aegean landmass from vertical displacement. Extension in the Hellenic arc area runs generally N/S, and crustal shortening forms an E/W insular platform. Here the oceanic crust is thinned to almost 1/2 its original thickness. The counter-clockwise motion is further expressed in the area by transcurrent faulting in the Northern Aegean,

beginning

in

the

fourth

extensional phase of the Mid-Upper Miocene. The outer regions of the Hellenic zones, by contrast, exhibit compressive geology. It is characterized by the presence of over 200 islands and is subdivided into various minor basins, such as the Crete sub-basin (Fig., 52), surrounded by a trench that is 2,500 m deep (Jolivet and Brun, 2010; Ring et al., 2010).

Fig., (52) Aegean Basin constructions, sub-basis and boundaries (after: Jolivet et al., 2012).

V.3- Adriatic Sea basin The Adriatic basin has geological and morphological characteristics that are quite particular. Over one third of the area of the sea bottom is no more than 50-60 m deep. The Adriatic basin lies between the Apennine mountain range and the area of the Balkans. It is a zone of great compression with the margin of the European plate dipping below the Adria plate. It is not a very deep sea, it filled rapidly with the sediments from the erosion of the two 82   

mountain ranges facing each other, and in a near geological future it is destined to disappear. It is subdivided into three different basins (Casero and Bigi 2012). The northern part, or Upper Adriatic Sea, is entirely covered with the alluvial deposits of the large rivers of the North East, specially of the Po river, and is characterized by a sea-bottom that degrades gently to a maximum depth of 75 m. The central part is a closed and more variable depression, the so-called Middle Adriatic trench (Fig., 53), which is 266 m deep. The southernmost part is known as the Lower Adriatic and is characterized by a plain that is 1,000 m deep on average. The basin reaches a maximum depth of 1,230 m near the coast of the Puglia region. Toward the south, the sea bottom rises to a depth of 800 m near the Strait of Otranto, which separates the Adriatic Sea from the

Fig., (53) tectonic setting of Adriatic Sea (After Mantovani et al., 2009) 

Ionian Sea (Cattaneo and Trincardi, 1999; Asioli et al., 2001). The south Adriatic is a deeper basin showing a complex morphology and a maximum depth of about 1200 m (Maselli et al., 2010). Overall, the Adriatic Sea is a mud‐dominated system where the Po River is the most important source of sediment. The flexure of the lithosphere belonging to the Adria margin started from the most internal areas and migrated eastward through time, forming foredeep basins oriented subparallel to the belts and filled by large quantities of terrigenous (siliciclastic) sediments, derived from the erosion of the incipient inverted margin (orogen and former foredeep). Each flexural phase was accommodated either by the sedimentation of a flysch wedge, or by the sub-marine gravitational emplacement of large rock masses detached from the inverted margin sequence (Casero and Bigi 2012).

83   

During the last 25 Ma the westward subduction of the Adria plate led to the formation of the Apennine chain, while the Adriatic basin became a foreland domain. During the Pliocene and Pleistocene, the central Adriatic basin was characterized by a high subsidence rate because of the eastward rollback of the hinge of the Apennine subduction (Royden et al., 1987). The southern Adriatic basin was, instead, characterized by a different tectonic style, showing uplift since the middle Pleistocene (Scrocca, 2006; Ridente and Trincardi, 2006). This different tectonic behavior has been ascribed to differences in the thickness of the Adriatic lithosphere subducted toward the west (Doglioni et al., 1994). The Sicily Channel Rift area is an example of the Miocene-Quaternary extensional phase. The Adriatic Sea itself is relatively shallow, and almost the ocean floor (a thick carbonitic platform underlain by continental crust) exhibits compressional deformation structures, except for the Ionian Abyssal Plain, which is thought to be underlain by Paleoceanic crust. The western Adriatic margin (eastern Mediterranean), part of the Apennine foreland, is characterized by a differentiated tectonic setting, showing high subsidence rates (up to 1 mm/yr) in the northern area and tectonic uplift (on the order of 0.3–0.5 mm/yr) in the southern part corresponding with the so‐called Apulia swell. The average subsidence rate of about 0.3 mm/yr appears greater than the average sediment supply rate (0.15 mm/yr), and this fact explains the overall back stepping of the 100 kyr regressive depositional sequences on the margin. The results obtained help to improve the understanding of the regional tectonics and can be used for quantitative reconstruction of Quaternary sea level changes in the Adriatic region (Maselli et al., 2010). The history of the Alpine orogeny, constituting the northwestern portion of the Adriatic, really begins in the Mesozoic as well, for the sedimentary strata which constitutes most of its orogenic elements was laid down in the continental margins of the ancient Tethys Ocean. The Alpine orogeny and the Calabrian arc orogeny are both results of convergent plate margin movement between Africa and Europe, and display some vertical uplift associated with the subsidence of Mediterranean sea-floor deposits during the Cenozoic (Mantovani et al., 2009). V.4- Ionian Sea basin The Ionian Sea represents a key area for the understanding of the evolution of the Mediterranean geodynamics, both for the Apennines and Hellenic subduction zones (Scandone, 1980; Angelier et al., 1982; Royden et al., 1987), and for the Mesozoic Tythyan paleogeography (Dercourt, 1986; Lemoine et al., 1986). This basin has been considered by Le 84   

Pichon (1982) as a landlocked basin or a trapped crust (Letouzey, 1986). The Ionian Sea occupies

the

central

southern Mediterranean.

part

of

Here

the the

maximum depth of the Mediterranean is reached (5,093 m in the Hellenic trench). It is characterized by deep trenches (Hellenic trench, Herodotus trench near the Libyan coast, Malta trench and Pantelleria trench), vast deep abyssal plains to the East, and less deep plains towards the West, as in the area near Sicily and the Sirte plain, near the Libyan coast. The experiences

Ionian the

major

Sea

perhaps

amounts

of

subsidence in the Central Mediterranean. The Ionian lithosphere is subducting underneath Calabria to the northwest (Selvaggi and Chiarabba, 1995). The associated accretionary wedge widely advanced in the Ionian Sea, particularly involving the sedimentary cover on the top of it. The southern and southwestern margins of Ionian Sea are the areas which have not yet been involved in Tertiary and Quaternary shortening of the Apennines and Hellenic subduction zones (Catalano et al., 2000). The Ionian Sea is characterized by the subduction of the African plate under the Calabrian Arc, making it one of the most

Fig., (54). Paleogeographic and palaeoenvironmental map of the western-central Tethys during Early Aptian, and position of the studied Apulia Carbonate Platform (Ap) to Ionian Basin. 2- Tectonic framework of the southern Apennine foldand-thrust belt and Gargano-Murge foreland (Puglia, southeastern Italy) (After Graziano, 2013). 

geologically active areas in our country. Even though, geologically speaking the Calabrian Arc, belongs geographically to the Apennine range of mountains, it is a small portion of the 85   

Alpine range of mountains, like Corsica and Sardinia (Fig., 54). The superficial expression of the subduction is the volcanic arc of the Aeolian islands (Doglioni et al., 1999). The Ionian Abyssal Plain in this region is characterized by differentially subsiding areas but generally experiences more than adjacent regions, contributing greatly to the uplift associated with the Alpine orogeny and the Quaternary coastal blocks. The Hellenic trench (a thrust fault linked to the convergent activity in the Mediterranean ridge) began propagation in Miocene and continues today; it constitutes a major element of Ionian seafloor topography. The extensional features in the Ionian region are somewhat subdued, the dominant tectonic activity is convergent and/or related to vertical movement. Towards the West the Ionian Sea is bordered by the deep Malta Slope, this 3,000 m drop separates the Ionian Sea from the Western Mediterranean. Between the two basins is the so-called Pelagian Block, an offshoot of the African coast that extends between Tunisia and Sicily, forming a submerged ridge, of which Malta and the Pelagian Islands (Lampedusa and Lampione) are the highest tops, so high that they emerge from the sea. The Apulia Carbonate Platform (ACP) and the bounding Ionian Basin (IB) (southeastern Italy) were two major paleogeographic domains of the Mesozoic-Cenozoic central Tethys. During Aptian times they were located apart from the European-African landmasses and their related influence (Graziano, 2013) V.5- The Tyrrhenian Sea The Tyrrhenian Sea is an almost triangular shaped depression, between Sardinia and peninsular Italy, and is the youngest of the deep Mediterranean basins. It has a depth of 3,800m and is the deepest of the western basins. The Tyrrhenian Sea is located in the center of the Mediterranean and it is a small back-arc basin developed behind the east-migrating Apennine chain. The Tyrrhenian rifting started about 10-12 Ma. A slab rollback mechanisms, which occurred at different rates between the northern and southern part of the basin, is commonly invoked to explain, respectively, the counterclockwise eastward migration of the Apennines and the SE migration of the Calabro-Peloritano arc (Pastore et al., 2011). Morphologically, it is a deep basin surrounded by sharp and deep slopes, cut by deep submerged valleys. The Tyrrhenian Sea is the youngest basin in the western Mediterranean, forming since the Tortonian (~9 Ma). It was opened, according to this reconstruction, as a result of a southeastward rollback of subduction systems near the margins of the Adriatic plate (Malinverno and Ryan, 1986) due to the collision of Corsica and Sardinia with the Apennines at ~18 Ma that led to a relative quiescence in back-arc extension between 18-10 86   

Ma. During this period, continental crust of Apennine units incorporated in the subduction zone, and impeded further eastward subduction rollback (fig., 55a). Thus, considerable crustal shortening occurred in the Apennines accompanied by thrust systems that propagated eastward (Rosenbaum et al., 2002). During the latest Miocene or the Early Pliocene (5 Ma) extension ceased in the northern Sea

Tyrrhenian

and

migrated

southward

to

southern

the

Tyrrhenian

Sea. This stage was characterized

by

considerable extension that culminated during the

Pliocene-

Pleistocene, when new oceanic crust formed. Contemporaneously, crustal

shortening

occurred

in

Southern

Apennines

and

the

in

Fig., (55a). Main physiographic and geophysical features of the Tyrrhenian Sea. (A) Bathymetric map of the Tyrrhenian Sea. (B) Moho Isobaths Map; three different Moho can be recognised: a new Neogene-Quaternary below the back-arc basins, an old Mesozoic Moho in the Adriatic-Ionian foreland areas (Adriatic Moho), and another old Moho below the Sardinia-Corsica block (C) Bouguer Gravity Anomaly Map. (D) Heat Flow Map (after Roberts and Bally, 2012).

Sicily

accompanied

by

counterclockwise block rotations in the former and clockwise rotations in the latter. These processes

have

been

controlled

by

rapid

rollback Ionian

of

oceanic

lithosphere

beneath the Calabrian

Figure (55b) The crustal structure of the Tyrrhenian Sea is shown along a regional cross section, The original uninterpreted seismic profiles are available in Scrocca et al. (2003; Roberts and Bally, 2012). 

arc (Rosenbaum et al., 2002). A deep, narrow, and distorted Benioff zone, plunging from the Ionian Sea towards the southern Tyrrhenian basin, is the remnant of a long and eastward migrating subduction of 87   

eastern Mediterranean lithosphere (Fig., 55b). From Oligocene to Recent, subduction generated the Western Mediterranean and the Tyrrhenian back-arc basins, as well as an accretionary wedge constituting the Southern Apenninic Arc (Sartori, 2003). It communicates with the other basins through 4 passages: a 300-400 m deep channel puts it in communication with the Ligurian Sea; a wide, 2,000 m deep channel between Sicily and Sardinia connects it to the Algerian basin; the Boniface Strait (that is max. 50 m deep) connects it to the Provence basin; and finally, the Strait of Messina is the connection (100 m deep) with the Ionian Sea. Large volcanic structures, which, for the time being are quiescent, rise from the sea bottom here. In the Tyrrhenian Sea, stretching started in late Miocene and eventually produced two small oceanic areas: the Vavilov Plain during Pliocene (in the central sector) and the Marsili Plain during Quaternary (in the southeastern sector). They are separated by a thicker crustal sector, called the Issel Bridge. Back-arc extension was rapid and discontinuous, and affected a land locked area where continental elements of various sizes occurred. Discontinuities in extension were mirrored by changes in nature of the lithosphere scraped off to form the Southern Apenninic Arc. Part of the tectonic units of the southern Apennines, accreted into the wedge from late Miocene to Pliocene, had originally been laid down on thinned continental lithosphere, which should constitute the deep portion of the present slab. After Pliocene, only Ionian oceanic lithosphere was subducted, because the large buoyancy of the wide and not thinned continental lithosphere of Apulia and Africa (Sicily) preserved these elements from roll back of subduction. After Pliocene, the passively retreating oceanic slab had to adjust and distort according to the geometry of these continental elements (Sartori, 2003). V.6- The Alboran Sea

It extends from the Strait of Gibraltar to the Balearic Basin. Its maximum depth is 1,500 m, that drops to 1,800 m in the Alboran rift that separates it from the Algerian basin. In the centre there is a small volcanic island, 10 m above sea level that rises from the sea bottom that is 1,500 m below. This part of the Mediterranean receives the direct influence of the Atlantic, because it is where the sea water mixes with the ocean water. The water here is generally colder and less salty and rich with organisms coming from the Atlantic (Ammar et al., 2007). The total volume of Neogene sediments deposited in these basins is ~209,000 km3 and is equally distributed between the internal (Alboran Basin and intramontane basins) and the external basins (foreland basins and Atlantic Margin). The largest volumes are recorded

88   

by the Alboran Basin (89,600 km3) and the Atlantic Margin (81,600 km3) (Iribarren et al., 2009). It is located in the western Mediterranean Sea, connected to the Atlantic Ocean through the Straits of Gibraltar to the W, and to the Balearic Basin through the Alboran Trough to the east (Fig. 56a). The Alboran Sea, in continuity to the east with the South Balearic Basin, is located in the inner

part

arcuate

of

belt.

this The

region as a whole is

Fig., (56a). Sedimentation basin of Alboran Sea, Western of Mediterranean (after Platt and Vissers 1989). 

bounded to the north and

south

Iberian

and

by

the

African

forelands, to the west by the Atlantic Ocean, and to the east it is connected

to

the

oceanic

Sardino-

Balearic Basin Comas et

al.,

1999).

The

Alboran Sea is a rift basin developed from the early Miocene to

Fig., (56b). A. Schematic true-scale section from the Gibraltar Arc to the South Balearic basin to illustrate the east-west crustal structure of the westernmost Mediterranean (after Comas et al., 1999) .

the present under a convergence regime between African and European plates (Comas et al., 1992; Garcia-Duenas et al., 1992). The northern margin of the Alboran Sea is a tectonically active margin located on the inner side of the Betic-Rifian alpine orogenic belt, whose formation is linked to the Neogene convergence regime (Perez-Belzuz et al., 1997). The onland geology is dominated by orogenic nappes of the Alpuja´ rride Complex, composed of Palaeozoic and Triassic rocks (Aldaya and Garcia-Duenas, 1976). Cliffed coastal segments 89   

occur in the study area due to the proximity of orogenic nappes. Depressed areas are filled with Plio-Quaternary deposits, represented by alluvial fans at the piedemont of orogenic nappes and by deltaic deposits (Lobo et al., 2006). Plio-Quaternary sediments in the Guadalfeo River deltaic plain show high granulometric variability, ranging from medium sands to gravels, or even boulders. Beaches along the Guadalfeo River prodelta are composed by sandy sediments and gravels (IGME, 1980). The formation of the Alboran Sea occurred during the westward migration of the subduction hinge. Rapid rollback was compensated by wholesale extension in the overriding continental crust, which was thinned to ~15 km between 23-10 Ma (Lonergan and White, 1997). Contemporaneously, fragments of continental crust were thrust onto the passive margin of Africa and Iberia (the External Zone), forming rotation patterns consistent with oblique thrusting derived by the westward rollback of the subduction zone. Final accretion of the Rif-Betic Cordillera occurred at ~10 Ma (Fig., 56b), when the subduction zone rolled back as far as Gibraltar. Subduction rollback then ceased, together with the cessation of backarc extension in the Alboran Sea (Lonergan and White, 1997; Rosenbaum et al., 2002). Depositional geometries and distribution patterns of shelf sediment wedges mainly derived from small rivers located in the northern margin of the Alboran Sea, Western Mediterranean Basin, are reported in this study, in order to understand: (1) their generation under particular physiographic and climatic conditions of river basins; (2) the interaction of shallow water wedges with submarine valleys. A high amount of data has been used in this study, including river discharge and wave climate data, multibeam bathymetry, highresolution seismic profiles and surficial sediment samples (Lobo et al., 2006; Ammar et al., 2007). The basins include the Alboran Sea, the intramontane basins, the Guadalquivir and Rharb foreland basins and the Atlantic Margin of the Gibraltar Arc. V.7- The Algerian Basin This is the vastest basin of the western Mediterranean area. Leaving the Alboran Sea to the west, it extends with a triangular shape from the Gulf of Valencia to the Ligurian Sea. Its maximum depth is 2,800 m, near the western coasts of Sardinia. It is characterized in its most western part, by the large deep sea cone of the Ebro River, where the continental shelf reaches a width of 60 km (Fig., 57a). Along the northern coasts, up to Genoa, the continental shelf is practically absent, it is no wider than 3-9 km. The sea bottom descends rapidly to 90   

depths over 2,000 m and is characterized by a number of submarine canyons that cut

across

canyons

it.

These

carry

large

quantities of material from the erosion of the emersed land toward the abyssal depths.

The

tectonic

evolution of the Algerian Alpine belt starts during the Eocene with the subduction of the Tethyan oceanic domain (Roca et al., 2004)

Fig., (57a) Simplified present-day geodynamic scenario of the Central–Western Mediterranean region superimposed on the topography and bathymetry. GL: Giudicarie Lineament; IL: Insubric Line (after Carminati et al., 2012).

in a context of a 15 mm a-1 N-S convergence between the European and African plates (Dewey et al., 1989). Simultaneously,

the

opening of the Algerian basin

commences

in

a

back-arc position and is associated Tethysian

with slab

the

roll-back

(Fig., 57b) and possibly with

slab

break-off

Fig. (57b) Representative 6-channel seismic profile across the Algerian margin, east of Algiers, off Dellys, crossing the Sebaou canyon on the slope (see location in figure 1). Top: line drawing of the whole section; bottom: interpreted enlargements of the seismic line. p.q.:Plio-Quaternary deposits, Mess.: Messinian deposits, Ant 1, 2 and 3: anticlines, (After . Strzerzynski et al., 2010b)

(Carminati et al., 1998a). After the splitting of the forearc and closure of the Tethyan ocean, the convergence rate between the European and African plates decreases to 5 mma-1 and deformation occurs onshore mostly on S-dipping thrusts progressively sealed by Miocene deposits and volcanic rocks (Strzerzynski et al., 2010b). The Algerian margin is affected by a Messinian sea-level fall responsible for subaerial erosion expressed by fluvial canyons. After the subsequent final sea level rise, the building of 91   

prograding Gilbert type fan deltas induces the infill of Early Pliocene rias in coastal basins such

as

the

Mitidja

basin. A coeval change of the motion of Africa relative to Europe, ~3 Ma ago (Calais et al., 2003; Mauffret, 2007): the

convergence

direction rotates about 20º

counter-clockwise

Fig., (58) The southern passive margin of the Algerian basin, behind the former subduction suture after the closure of the Tethys ocean. To the south: southverging Tellian fold and thrust belt After (Domzig et al., 2010). 

and becomes NW-SE at the longitude of the central Algerian margin (Fig., 58). Onshore, the late Pliocene to Quaternary deformation is expressed by the folding of the Early Pliocene Mitidja deposits at its southern boundary and near the Algiers Sahel anticline, where Late Pliocene to Quaternary beach deposits are located up to 350 m and are directly correlated to the anticline growth. East of Algiers, Late Pliocene to Quaternary deformation is also evidenced by eastward migration of the Isser River bed and uplifted beaches (Boudiaf et al., 1998). A first estimate of the beginning of the new deformation regime is given by the Piancenzian, i.e. 2.6 Ma, age of the last deposits below the uplifted beach sediments (Boudiaf et al, 1998). Offshore Algeria is a key area to study the reactivation in compression of a Cenozoic passive margin. This region is often affected by Mw=6-7.5 earthquakes (Mauffret, 2007; Domzig et al., 2010). The Algerian margin has originated from the opening of the Algerian basin about 25–30 Ma ago. The central margin provides evidence for large-scale normal faults of Oligo-Miocene age, whereas transcurrent tectonics characterizes the western margin. A set of NW–SE oriented dextral transform faults was active during basin opening and divided the 600 km long central margin into segments of

120–150 km (Carminati et al., 2012;

Strzerzynski et al., 2010a). The morphology of the margin and the structure of the Neogene sediments on the slope and in the basin, particularly the Plio-Quaternary sediments, are shaped by recent fault-related folds and near-surface faults distributed across the margin and also found far on land. Morphological and structural interpretation of the available data along 92   

the ~1000 km of the margin leads us to characterize several fault segments with a variable length and position. In Central Algeria (Algiers region), the main contractional structures are active blind thrusts (Plio-Quaternary) generally located near the ocean-continent transition and verging to the north (opposite to preexisting features). They form generally large asymmetrical folds sub-perpendicular to the present-day convergence direction, which are often arranged in en echelon segments at different scales. Offshore Boumerdes (east of Algiers), we show that the faults have typically a flat-and-ramp geometry creating a succession of perched basins from the mid-slope down to the deep basin, and prograding towards the basin (Carminati et al., 2012). Although the Messinian salt tectonics and the sedimentary fluxes at the outlets of canyons play a significant role, the sediment deposition as well as the morpho-structure of the margin appear to be controlled at first order by these slowrate tectonic movements, indicating a clear interaction between crustal-scale tectonics and sedimentation. We discuss the implications of these results in terms of seismic hazard and sedimentary architecture (turbidites) in deep environments (Domzig et al., 2010). The upper Miocene, Plio-Quaternary, and present-day tectonic setting is, however, compressional and supports the occurrence of a margin inversion, a process still poorly documented worldwide (Strzerzynski et al., 2010a). The central Algerian margin represents a rare example of inverted margin, where the process of subduction inception is particularly well expressed and helps understand how extensional and transtensive structures are involved in margin shortening (Strzerzynski et al., 2010a). Pre-Miocene structures such as basement highs and transform faults appear to control changes of the deformation pattern along this part of the margin, resulting in different widths, geometries, and relative positions of folds and faults. Plio-Quaternary and active blind thrust faults do not reuse Oligo-Miocene normal and transform faults during inversion, but instead grow within the continental margin, at the foot of the continental slope and at the northern sides of basement highs interpreted as stretched continental blocks of the rifted margin. The inherited structures of the margin appear, therefore, to determine this deformation pattern and ultimately the earthquake and tsunami sizes offshore. The complex geometry of the fault system along the Algerian margin suggests a process of initiation of subduction in its central and eastern parts (Strzerzynski et al., 2010a).

93   

VI- Geothermal Potentials and Uses of the Mediterranean Geothermal resources are suitable for many different types of uses but are commonly divided into two categories, high and low enthalpy and according to their energy content. High enthalpy resources (>150 °C) are suitable for electrical generation with conventional cycles, low enthalpy resources (<150 °C) are employed for direct heat uses and electricity generation using a binary fluids cycle (EC, 1999). The large geothermal potential worldwide available within a few km depth in several on land and marine areas of the Mediterranean Sea is encouraging investors and enterprises to invest in geothermal exploration for power generation and for combined heat and power co‐generation. VI. A- Geothermal potentials Geothermal energy is the natural heat o f the earth. Immense amounts of thermal energy are generated and stored in the earth’s core, mantle, and crust. The heat is transferred from the interior towards the surface mostly by conduction. This heat flow makes temperatures rise with increasing depth in the crust on average by between 25-30°C/km. An average thermal gradient o f 30°C/km means that at a depth of 2 km the temperature in the rocks is around 70°C in areas where there is no volcanic activity and where ground water is not affecting the thermal gradient (EC, 1999). The exploitable geothermal resources in the Mediterranean are generally related not to conductive systems but to convective ones. This means that the heath is brought near the surface by fluids (mainly waters) flowing vertically from depth toward the surface, so that sufficiently high temperature may be reached by drilling at economical depth. Geothermal resources are suitable for many different types of uses and according to their temperature are commonly divided into two categories, high and low enthalpy. High enthalpy is suitable for electrical generation with conventional cycles, low enthalpy resources are employed for direct uses (EC, 1999). •

High temperature resources, used for power generation (with temperatures above 150 °C) are confined to areas geologically active, that is where movements of the earth crust bring the magma near the surface.

•

Low temperature resources which are mainly used for heat production (with temperatures below 150°C) can, on the other hand, be found in most countries. These are formed by the deep circulation o f meteoric water along faults and fractures, and by water residing in high porosity rocks, such as sandstone and limestone, at sufficient depths for the water to be heated by the Earth's geothermal gradient.

94   

A.1- Geothermal Resources in Foreland Environments: Geothermal resources are commonly confined where high heat flow (> 70 mW/m2) is recorded and extension controls the tectonic evolution, determining diffuse fracturing in rocks, underneath an impervious cover. Nevertheless, areas with low heat flow and located in foreland tectonic settings can be also affected by geothermal manifestations, although in spot‐areas and with low temperature (about 25‐28°C) geothermal fluids, as it is the case of the Santa Cesare Terme zone, located in the Apulia carbonate platform, the foreland of the southern

Apennines

(Cretaceous‐Pleistocene).

The

platform

is

constituted

of

a

Jurassic‐Cretaceous succession, thick more than 5 km in the study area, and believed to rest over the Late Triassic evaporite (Burano Fm). Oligocene‐Pleistocene calcareous and terrigeneous sediments rest unconformably over the Platform. The area is deformed by transtensional structures, thus determining extensional jogs and pull‐apart structures where the permeability is enhanced. It is therefore concluded that along these almost vertical structural channels the upflow of deep fluids, heated through the thermal gradient normally typifing foreland areas (Liotta, 2012). A.2- Thermal Coastal Springs: Carbonate aquifers represent important thermal water resources outside the volcanic areas, supplying spans or geothermal installations. The thermal springs constitute so the discharge areas of the deep groundwater flowing within these carbonate aquifers whose hydraulic conductivity and the relevant geothermal fluid migration are strictly controlled by both the discontinuities network and the karsification processes. An example of these springs occurs along the south‐easternmost portion of the Apulia region (Southern Italy) where some sulphurous and warm waters (25‐33°C) flow out in partially submerged caves located along the coast, supplying so the spas of Santa Cesarea Terme. These springs are known from ancient times (Aristotele in III century BC) and the physical‐chemical features of their thermal waters resulted to be partly influenced by the sea level variations (Polemio et al., 2012). In Morocco there are several geothermal anomalies and thermal clues, with occurrence of numerous hot springs and important deep aquifers; thus it could be considered as a real geothermal promising country. Measured temperature of hot springs ranges from 21 to 54°C and disharge rates from 2.5 to 40 l/s. Geothermometers applied are: silica, Na/K, Na-K-Ca, Na-K-Ca-Mg, Mg/Li and Na/Li (Zarhloule, 2003).

95   

VI. B- Geothermal Uses The direct use o f geothermal energy can involve a wide variety of applications including the geothermal heat pumps. In most industrialised countries, a significant percentage of the energy consumption is devoted to heat production at temperatures of 50-100 °C, which are common in low-enthalpy geothermal areas. Most of this energy is supplied by the burning o f oil, coal or gas at much higher temperatures. The scope for using geothermal water alone as well as in combination with other local sources of energy is therefore very large (EC, 1999). The direct use o f geothermal energy is at a relatively advanced stage in European countries compared with other parts of the world. It supplies a wide range of applications and uses due to the versatility and demand for base-load heat demand plus the availability o f the resource. European countries have been pioneers in the exploitation of geothermal resources. European experience and expertise in this sector has been duplicated by other countries world-wide. However, European operators should still be in a position to maintain their leading role in the development and utilisation o f geothermal energy for both direct use and for electricity production (EC, 1999).   B.1- Electrical production Italy has been the first country to exploit geothermal energy for electrical production, starting in early 20th century in Larderello, Tuscany. One century after, a large electrical production (about 6000 GWh per year, with an installed power of about 800 MW). A sound and careful program of geothermal exploitation of existing resources in Italy could lead, in a period of 10‐20 years, to increase geothermal‐electrical production of a factor 2 to 5, possibly increasing its contribution to the total budget from the present 2% up to 10% (De Natale et al., 2012). Germany, Geothermal power generation is done through the use of binary cycle technology. Since November 2003, a ca. 0,2 MWe pilot power plant using this process is exploited at Neustadt-Glewe and another twenty megawatts (4 or 5 power plants) is currently in the planning and construction stage, chiefly in Southern Bavaria. The most advanced project is that of Unterhaching (Antics and Sanner 2007). Geothermal electricity production in Iceland has increased significantly since 1999, with the installation of new plants in Svartsengi, Krafla and Nesjavellir, up to the present value of 202 MW. An additional 30 MW single flash unit at Nesjavellir is at an advanced stage of construction (Ragnarsson, 2005; Gunnlaugsson, 2003).

96   

In Portugal, exploitation of geothermal energy to produce electricity has been developed on the volcanic archipelago of the Azores, or more precisely on the Sao Miguel Island. This island has five geothermal power plants achieving a total capacity of 16 MWe (Antics and Sanner 2007). The electricity generation in Turkey has been increased to 30 MWe with the addition of the Aydin-Salavatli binary cycle geothermal power plant, adding a 10 MWe installed capacity to the existing Kizildere geothermal power plant (20 MWe installed capacity) (Mertoglu et al, 2007). France started up its second geothermal power plant in 2004 on the Bouillante site, i.e. an additional 10 MW (14.7 MWe in total), that could produce an additional 72 GWh per year. Furthermore, the Bouillante 3 feasibility study, launched in 2003, could result in a third power plant with more than 10 MWe capacity (Antics and Sanner 2007).

V- Mineral Resources in the Mediterranean Region V.1- Organic minerals (Oil – Natural Gas – Coal) 1. A- Oil and natural Gas resources A.1- Lavantine basin Assessment of undiscoverable gas resources of the Lavantine basin province (East Mediterranean) using current technology were estimated by the USGS (U.S. Geological Survey) to be about 3.5 tcm (trillion cubic meters) of gas. Already in the Israeli E.E.Z. (Exclusive Economic Zone) an amount of 800 bcm (billion cubic meters) has discovered in the fields of Marie B, Gaza Marine. In the Cypriot part of the Levant basin, the estimated amount of gas reserves around 300 bcm. In the Nile delta and the E.E.Z. of the Cyprus Republic USGS has estimated a natural gas potential of 6.3 tcm, besides the 2.2 tcm of gas and 1.7 Bbbl (Billion barrels) of oil already discovered in the Egyptian E.E.Z. Out of the 6.3 Tcm. These estimated resources are comparable to some other large gas provinces encountered in the world. In the same region, crude oil potential reserves of about 1.7 Bbbl of oil and about 6 Bbbl of gas condensate are also estimated by USGS to exist (Bruneton et al., 2011). A.2- Adriatic Sea During the geological period from Triassic until Upper Lias, Dinarides and Apulian platform formed one consistent unit. Since Upper Lias until the end of Upper Cretaceous, due to paleo-tectonic influence, this consistent platform was separated by Adriatic Basin called 97   

’Scaglia-Biancone Basin’ with pelagic and hemi-pelagic younger Mesozoic deposits and during Tertiary with clastic sediments of flysch and molasse type. Platforms are divided from the basins by steep offshore slopes where periplatform carbonates clastics and turbidites were sedimented. Due to the obvious analogy between Apulian and Dinarides slopes and their petroleum-geological characteristics (Grandic and Kolbah 2009). The separation episode on the Italian side is characterized by Rosso Ammonitico stratigraphic horizon which in the Dinarides corresponds to ’Spotted limestone’ formation in the top of Lithiotis deposits. However, the term Adriatic carbonate platform (Veli, et al. 2002; Vlahovi, et al. 2002; Vlahovi, et al. 2005) has been used lately for offshore and onshore part of carbonate sediments which were formed in the period from Triassic to Paleogene (Grandic and Kolbah 2009).

A.3- Neogene petroleum system at Alboran - Algerian Basins The Algerian offshore is part of the southern margin of the western Mediterranean Sea. The western part of this offshore area represents the transitional margin between the South Algero-Balearic Basin and the Alboran Basin. The Yusuf-Habibas Ridge is a major EW-striking structure of this complex plate boundary, separating the eastern and southern parts of the Alboran Basin from the South Algero-Balearic Basin. The ridge played an important role during the Neogene Alboran westward block migration between the Africa and Iberia plates, while the Kabylies blocks migrated southward and accreted to Africa. Three main reservoirs are recognized in the Habibas well sedimentary section: (1) sandstones in the Pliocene, above the Messinian evaporites; (2) sandstones in the Middle-to-Upper Miocene, below the Messinian evaporites; and (3) carbonates and sandstones in the older allochthonous units (Medaouri et al., 2012).

1.B- Coal Bearing Formations Turkey: During the period Neogene to Quaternary, several lacustrine basins developed in Turkey. These basins are generally characterised by volcanic-sedimentary successions. These basins are characterized by important fossil–fuel and industrial–mineral resources such as lignite, oil shale, clays, borates and zeolites (Sener et al 1995; Sener and Gundogdu 1996). The coal-bearing Hirka Formation was deposited over the Galatian Andesitic Complex and/or massive lagoonal environments during the Miocene (Sener, 2007). The lignite-bearing Yoncalı formation was found between Yozgat and Sorgun, in central Anatolia (Akkiraz et al., 98   

2008).The Middle–?Upper Eocene Yoncalı formation dominantly consists of continental and shallow marine sediments containing basal conglomerate and fine-grained sediments at the base. These fine-grained sediments are overlain by flysch-like sediments including fossiliferous reef limestone lenses. Coal seams occur in the lowest part of the unit and are intercalated with sandstone and mudstone layers (Akkiraz et al., 2008).

V.2- Inorganic mineral resources Turkey: During the period Neogene to Quaternary, several lacustrine basins developed in Turkey. The Neogene basins were filled by clayey, carbonaceous and sandy sediments, and also by explosive products of contemporaneous K-rich calc-alkaline volcanism with various degrees of crustal contamination (Yılmaz 1989; Gulec 1991; Inci 1991; Gundogdu et al., 1996).The mineral matter of the basins are mainly clay minerals (illite–smectite and kaolinite), plagioclase and quartz in Bolu coal field, clay minerals (illite–smectite, smectite and illite), quartz, calcite, plagioclase and gypsum in Seben coal field, quartz, K-feldspar, plagioclase and clay minerals (kaolinite and illite), dolomite, quartz, clinoptilolite, opal and gypsum (Sener, 2007). Italy: Supercritical Fluids in Geothermal Systems: Information from Fluid Inclusions Trapped in Minerals of the Larderello Geothermal Field and from the Study of Fossil, Magmatic‐Related, Hydrothermal Systems of Southern Tuscany: At Larderello, the possible occurrence of a high‐temperature fluid phase below the vapor-dominated reservoirs hosted within metamorphic and sedimentary rocks, is suggested by the occurrence at about 3‐5 km depth by the abundant fragments of quartz–tourmaline veins erupted by a geothermal well. Some of the fluids could escape from K‐horizon through crustal shear zones, which are interpreted to be generated by the extensional tectonic events (Early, Middle Miocene ‐ Present). These events affected the inner Northern Apennines after collisional tectonics (Cretaceous‐Late Oligocene, Early Miocene), and determined the thinning of the crust and the lithosphere to the present thickness values of about 22 and 30 km, respectively. Extensional tectonics was coeval with the emplacement of shallow‐level igneous intrusions since early Miocene (Corsica) and progressively shifted eastward. These intrusions, at Larderello are testified by the Pliocene‐Quaternary peraluminous leucogranites and monzogranites found in several deep wells (between 2.5 and 4.5 km depth) with significant fluorine and boron content, and their thermometamorphic aureoles (Rocchi et al., 2010). 99   

Fluid inclusion studies on granites and thermometamorphic rocks found in Larderello geothermal wells indicate the occurrence of three nearly coeval fluids which could record the conditions within a past K‐horizon: 1) Na‐Li‐rich brines, probably exsolved from granites during crystallization, 2) high‐saline fluids and aqueous vapours produced by boiling of the Na‐Li brines and 3) aqueous‐carbonic fluids (H2O+CO2±CH4±N2), and formed as a consequence of de‐hydration processes and graphite‐water interaction during heating of the Paleozoic metamorphic (sometimes C‐rich) rocks (Boiron et al., 2007). Interpretation of fluid inclusion data alone, or combined with contact metamorphic mineral equilibria, indicates that the early fluids were trapped at high temperatures (≥420°C) under infra‐lithostatic or lithostatic pressure conditions. Evidence for boron metasomatism and veining (tourmaline precipitation), occurring at the contact between the granites and schistose metamorphic basement in the eastern sector of Elba Island, are also present at Larderello, and in particular testified in the K‐horizon by the quartz‐tourmaline veins fragments erupted by a geothermal well (Dini et al., 2008). Moreover, eastern Elba is characterized by cataclastic level, associated to a significant network of mineralized Fe‐ and quartz veins (Ruggieri et al., 2012). Spain: Th- and U-bearing minerals, which were recently found in the SE Mediterranean margin of Spain. These minerals are REE phosphates (mainly monazite) which occur as amoeboidal-to-elongate inclusions, from around 10 μm to 120 μm, hosted in single garnet crystals from dacite lavas and metamorphic rocks from the El Hoyazo Volcanic Complex. Th and U contents are higher than 1 wt%, with 3.04 to 5.62 wt % for ThO2, and 0.7 to 1.75 wt% for UO2. Both elements are also found in xenotime (ThO2: 0.24, UO2: 0.27 wt%). Given that the erosion of the volcanic source rocks has generated a "placer-type" deposit of monazite sands and that garnets (main carriers of monazite) are being commercialised, an environmental monitoring and management plan should be urgently executed in the area (Martinez-Frias et al., 2004). Bulgaria and Greece: Hydrothermal ore deposits related to post-orogenic extensional magmatism and core complex formation The Rhodope Massif in southern Bulgaria and northern Greece hosts a range of Pb– Zn–Ag, Cu–Mo and Au–Ag deposits in high-grade metamorphic, continental sedimentary and igneous rocks. Following a protracted thrusting history as part of the Alpine–Himalayan collision, major late orogenic extension led to the formation of metamorphic core complexes, block faulting, sedimentary basin formation, acid to basic magmatism and hydrothermal activity within a relatively short period of time during the Early Tertiary. Large vein and 100   

carbonate replacement Pb–Zn deposits hosted by high-grade metamorphic rocks in the Central Rhodopean Dome (e.g., the Madan ore field) are spatially associated with low-angle detachment faults as well as local silicic dyke swarms and/or ignimbrites. Ore formation is essentially synchronous with post-extensional dome uplift and magmatism, which has a dominant crustal magma component according to Pb and Sr isotope data. Intermediate- and high sulphidation Pb–Zn–Ag–Au deposits and minor porphyry Cu–Mo mineralization in the Eastern Rhodopes are predominantly hosted by veins in shoshonitic to high-K calc-alkaline volcanic rocks of closely similar age. Base-metal-poor, high-grade gold deposits of low sulphidation character occurring in continental sedimentary rocks of synextensional basins show a close spatial and temporal relation to detachment faulting prior and during metamorphic core complex formation (Marchev et al., 2005).

101   

CHAPTER V EGYPT (Genius of the Place) 1- Sedimentary Basins of Egypt Egypt, located in the northeastern corner of the African continent, is bounded to the east by the Red Sea and by what has been interpreted as a median spreading center in the Red Sea and Gulf of Suez (McKenzie et al., 1970). Such a tectonic setting suggests that this area may be suitable for geothermal development. The far northern end of the Red Sea is divided into two parts by the Sinai Peninsula: the Gulf of Suez in the west, and the Gulf of Aqaba in the east. Egypt can be subdivided into five major morpho-structural units; 1) the Mediterranean Fault Zone, 2) a belt of linear uplifts and half-grabens, 3) the North Sinai Fold Belt “Syrian Arc”, 4) the Suez and Red Sea Graben, and 5) the intracratonic basins of southern Egypt (Sestini, 1995) (Fig., 59). Pre-Cambrian basement rocks outcrop in South Sinai, the Red Sea Mountains, at Aswan and near Sudan. The sedimentary section overlying the basement, 1-3 km in the south (Kharga Oasis), thickens northwards to 5-6.5 km near the Mediterranean, but with notable irregularities in the basinal areas (e.g. 10-13 km in Abu Gharadig Basin versus 3 km on the Ras Qattara Ridge at its north margin). The distribution of sedimentary facies follows a simple north-south trend: sands increase in percentage and grade from shallow marine to predominantly continental (including coals) towards the south, whereas carbonate rocks are more common in the north, except in the stratigraphic intervals that correspond to southwards transgressions (Sestini, 1995; El Diasty et al., 2012). 1.1- Nile Delta The Nile Deep-Sea Fan (NDSF) forms a thick sedimentary wedge covering about 100,000 km2, constructed, for the most part, since the late Miocene by influx of clastic sediments from the Nile River (Dolson et al., 2000). The present day NDSF covers a segment of an older passive margin thought to have formed during successive rifting episodes in Jurassic and early Cretaceous times (Hirsch et al., 1995). According to Aal et al. (2001) and Mascle et al. (2003), the total thickness of sediments on the Egyptian margin (including the post-Miocene NDSF) could exceed 9 km (Loncke et al., 2006). The geodynamic framework of the eastern Mediterranean and its surroundings is characterized by a complex pattern of active, thick-skinned, crustal-scale tectonics (Mascle et al., 2000; McClusky et al., 2000), resulting from interactions between various tectonic plates 102   

and

microplates.

Geodynamic

features

surrounding the region are (a) in the southeast, the almost-aborted Suez Rift; (b) in the east and northeast, the Dead Sea/Levant and East Anatolian Fault zones related to the motion of the Arabian plate with respect to Africa; (c)

northward,

along

the

eastern Hellenic and Cyprus arcs, the subduction/collision of Africa beneath Europe and the rapidly moving Aegean– Anatolian microplate; and (d) the

Egyptian

margin,

a

passive margin of Mesozoic age that may have been reactivated

partly

during

Miocene rifting of the Suez– Red Sea Rift system (Mascle et al., 2000). In this tectonic framework, sediments of the NDSF drape onto the former Egyptian passive margin and reach north to the subducting Fig., (59) Geological maps of Egypt.

Tethyan oceanic domain. The

distal parts of the NDSF, however, have not yet reached the Hellenic and Cyprus arcs: its western edge feeds the accretionary Mediterranean Ridge (MR), whereas its eastern corner is bounded by an almost flattopped, subcircular seamount (Eratosthenes Seamount, hereafter referred to as ‘ESM’). This bathymetric high is interpreted to have a continental origin, having been rifted away from the African/Levant domain during the Mesozoic. It is currently colliding against the island of Cyprus (Robertson et al., 1995; Guiraud and Bosworth, 1999). 103   

According to Abdel Aal et al. (2000) and Samuel et al. (2003), two major fault trends characterize the offshore NDSF. The Temsah trend (oriented NW–SE) and the Rosetta trend (oriented NE–SW to ENE–WSW) are both thought to be inherited from the Mesozoic rifting phase (Loncke et al., 2006). The structural elements affecting the northern margin of Egypt, including the Nile Delta, were formed during the tectonic evolution of the southern part of Eastern Mediterranean basin (Abdel Aal et al., 2001; Abd-Allah, 2008). This region represents the Northeast African continental margin that is covered by the Nile Delta sediments. Likes the other deltas in the World, the largest Nile Delta has attached attentions of several hydrocarbon companies. The pre-existing faults were reactivated during the evolution of the Nile Delta by two tectonic events. These events took place during the Late Miocene–Early Pliocene and Late Pliocene–Early Pleistocene times and were coeval with two falls in the sea level pattern. The thickness of the Pliocene-Recent sediments and the location of the pre-existing faults controlled these reactivations. The mechanical contrast of these sediments and fault displacements controlled the geometry of the reactivated faults. The northwest sinking (bending) of the outer part of the African continental margin under the Eurasian plate at the Hellenic subduction Arc has induced a tangential northwest trending extension (Abd-Allah et al., 2012). 1.2- Eastern Desert The Eastern Desert of Egypt constitutes the northwestern end of the Nubian segment of the Arabian-Nubian Shield. The ophiolitic rocks of the Arabian-Nubian Shield have suprasubduction geochemical signatures (Stern et al., 2004), but Zimmer et al. (1995) reported the occurrence of mid-ocean ridge (MOR) ophiolite in the Gerf area in the southern Eastern Desert. The supra-subduction signature of the ophiolites in the Eastern Desert led to further debate on whether they were formed in a back-arc setting (El-Sayed et al., 1999; Farahat et al., 2004) or in a forearc setting during subduction initiation (Azer and Stern, 2007; Khalil and Azer, 2007). Azer and Stern (2007) proposed that the Neoproterozoic ophiolites of the Eastern Desert were formed in a forearc setting based on the depleted nature of the serpentinized mantle rocks. Although their conclusion is consistent with other ArabianNubian Shield ophiolitic mantle units (Stern et al., 2004), alternative geodynamic settings have been proposed for the upper-mantle peridotites of the Central Eastern Desert. Khalil (2007) inferred a mid-ocean ridge tectonic setting for the mantle rocks of Wadi Ghadir ophiolite in the Eastern Desert. Ophiolitic gabbros and pillow lavas in the Central Eastern Desert were interpreted as remnants of oceanic crust formed in a back-arc basin (Farahat et al., 2004; Abd El-Naby and Frisch, 2006; Abd El-Rahman et al., 2009). 104   

The geodynamic origin of the Neoproterozoic ophiolites of the Arabian-Nubian Shield exposed in the Eastern Desert of Egypt remains controversial. Fawakhir ophiolite and from some mélange blocks along the Qift-Qusier Road were used to constraint the tectonic evolution of this part of the Central Eastern Desert. Neoproterozoic crustal growth of the Arabian-Nubian Shield was accomplished mostly through the accretion of island arcs to continental margins (El-Shafei and Kusky, 2003; Jons and Schenk, 2007). The final collision between West and East Gondwana resulted in the Pan-African orogeny (Korِner et al., 1987). The crustal evolution of the Eastern Desert culminated in the eruption of the Dokhan Volcanics, deposition of molasse-type Hammamat sediments, and emplacement of younger granites (Eliwa et al., 2006; Abd

El-Rahman

et

al.,

2009). 1.3- Red Sea Rift Valley The

Red

Sea

occupies part of a large rift valley in the continental crust of Africa and Arabia. This break in the crust is part of a complex rift system that includes the East African Rift System (Said, 1962). To the north, the Red Sea

Fig. (60) shows a stratigraphic and structural cross section of the central Gulf of Suez. The stratigraphic record of the Gulf of Suez shows that the Gulf existed as a shallow embayment of Tethys as early as the Carboniferous and that a landmass lay at its southern end until the late Cretaceous (After Abdel Zaher et al., 2011)

bifurcates into the Gulfs of Suez and Aquaba, with the Sinai Peninsula in between. The Gulf of Suez is a failed intercontinental rift that forms the NW–SE trending continuation of the Red Sea rift system and was initiated during the late Oligocene to Early Miocene by the NE–SW separation of the African and Arabian plates (Patton et al., 1994). It extends more than 300 km in length and can be divided into three parts: the northern portion of the Gulf dips to the SW; the central part dips to the NE; and the southern part dips to the SW. The structure of the Gulf of Suez region is governed by normal faults and tilted blocks, of which the crests represent a major hydrocarbon exploration target (Fig., 60). The faults can be divided into two major sets based on trend. The first set is longitudinally parallel to the axis of the rift created in an extensional regime during the Neogene. The second consists of transverse faults with dominant N–S to

105   

NE–SW trends that inherited passive discontinuities in the Precambrian basement rock (Colletta et al., 1988). The predominantly clastic sediments that characterize its early history transitioned to calcareous marine sediments in the Cenomanian. Igneous rocks younger than Precambrianin the Sinai and neighboring areas are predominantly basaltic dikes and flows of Mesozoic (Meneisy and Kreuzer, 1974) and Oligocene to Lower Miocene age (Siedner, 1973). Their main direction is parallel to the Suez and Red Sea rifts. The rift stratigraphy and related tectonics are well documented (Evans, 1988; Schütz, 1994). The Gulf of Suez is a failed intercontinental rift that forms the NW–SE trending continuation of the Red Sea rift system. This rift is structurally controlled largely by extensional normal faults that strike northwest, forming a complex array of tilted half grabens and asymmetric horsts (Pivnik et al., 2003). 1.4- Western Desert The Western Desert of Egypt

consists

of

a

number of sedimentary basins that received a thick

succession

Mesozoic

of

sediments.

Fig., (61) Sedimentary basins of Nile Delta and Western Desert

Various geological studies have been carried out dealing with the stratigraphy, facies distribution, and tectonic framework of these sedimentary basins (Fig., 61). The sedimentary section in the northern part of the Western Desert can be divided into three sequences based on lithology, namely: the lower clastic unit from Cambrian to pre-Cenomanian, the middle carbonates from Cenomanian to Eocene and the upper clastic unit from Oligocene to Recent (Said, 1962). The stratigraphic sequence in the northern part of the Western Desert is characterized by a number of major transgressive/regressive cycles on the platform margin. The Mesozoic sequence unconformably overlies Paleozoic rocks. The Mesozoic stratigraphic succession is much better understood than the Paleozoic one as it is encountered in all studied wells, albeit in different thicknesses, as indicated by Moussa (1986), Barakat et al., (1987) and Shalaby et al., 2008). The intra-cratonic Abu Gharadig Basin is an eastewest trending half graben of Late Mesozoic age in which the depth-to-basement exceeds 10,000 m. Its northern margin is 106   

known as the Qattara Ridge (where the depth to basement is about 3300 m); while to the south it is bounded by the Sitra Platform. The basin is divided into northern and southern subbasins by an eastewest trending horst (EGPC, 1992). The structure at the Abu Gharadig Field is a faulted, asymmetric anticline,

which

was

formed during the Late Cretaceous-Early Tertiary (Abdel Aal and Moustafa, 1988). The

sedimentary

section of the Western Desert ranges from Early Paleozoic to Recent. Four major sedimentary cycles occurred, with maximum, southward transgression in Carboniferous,

Upper

Jurassic, Middle and Late Cretaceous,

Middle

Miocene and Pliocene time (Schlumberger, 1984). In the unstable shelf area, the lithostratigraphic

column

Fig., (62) geological sequences comparison in the different localities of Egypt

of the overlying series maybe subdivided into three sequences. First the lower clastic unit, from Cambrian to Cenomanian, second the middle carbonates, from Turonian to Eocene and finally the upper clastic unit, from Oligocene to Recent (Fig., 62). In the northwestern corner of the Western Desert along the Libyan border a NeS trending synclinorium (Siwa Oasis - Faghur) has been delineated with a Paleozoic section of some 3000 m thickness; mostly continental to shallow marine sandstones, siltstones and shales, with thin intercalations of carbonates (El Diasty et al., 2012). Throughout Mesozoic time, continental environments prevailed over the Western Desert south of Latitude 28⁰N. The Lower Jurassic, Wadi Natrun Formation consists of lagoonal deposits; alternating with dense limestone, green shales and dolomite. Middle-Late Jurassic 107   

rocks are represented by the Khatatba Formation, a thick carbonaceous shale sequence, with interbedded porous sandstone, coal seams and limestone streaks (Jenkins, 1990; Keeley and Wallis, 1991). Basinwards, the Khatatba Formation grades into the time equivalent Masajid Formation, made up of platform carbonates, including oolitic, reefal and dolomitic limestones with cherty intervals (Schlumberger, 1984). A widespread unconformity is recorded at the Jurassic-Cretaceous boundary. The Lower Cretaceous clastic series correspond to a transgressive cycle, with fluvio-continental sediments at the beginning (Neocomian) and at the end (Late Albian-Early Cenomanian) with a transitional near-shore to deltaic depositional environment during Early Aptian and Albian. The Late Cenomanian-Turonian Abu Roash Formation consists of an alternation of dolomitized calcarenites, shale and sandstones; the carbonates become more abundant and thicker northwards. The top of the Western Desert sequence is mostly formed by terrigenous clastics, the Late Eocene-Oligocene Dabaa Formation (200e400 m, max. 825 m) marine shales, and the Late Oligocene to Early Miocene Moghra Formation, 200e970 m, mainly sandstones, fluvio-marine, lagoonal to shallow marine upwards (Sestini, 1995; El Diasty et al., 2012). 2- Geothermal Regime of Egyptian Basins The thermal data at the eastern part of Egypt indicate that the geothermal situation of the Red Sea is more complex and broader than the Gulf of Suez. Observations near to the axial trough of the Red Sea have a mean of 470mWm2 that typical associated with an active spreading center. Whereas a mean of 116mWm2 was recorded near the coast of the northern Red Sea (Boulos, 1990) that is appropriate with the estimated values at the Gulf of Suez. According to Hosney (2000), two heat flow provinces were distinguished: 1- the west of Nilenorth of Egypt normal province with low heat flow about 46 mWm-2 and reduced heat flow of 20 mWm-2 typical of Precambrian platform tectonic setting and 2- the eastern Egypt tectonically active province with heat flow up to 80-130 mWm-2 including the Gulf of Suez and the northern Red Sea Rift System with reduced heat flow of > 30-40 mWm-2 , at the transition between the two provinces. Chemical and isotopic analyses of thermal waters of the main hot springs in the areas around the Gulf of Suez were performed by Sturchio and Arehart (1996). The preliminary heat flow values ranging from 42 to 175 mW m-2 have been estimated for Egypt from numerous geothermal gradient determinations with a reasonably good geographical distribution, and a limited number of thermal conductivity determinations. For northern Egypt and the Gulf of Suez, gradients were calculated from oil well bottom hole temperature data; east of the Nile, and at three sites west of the Nile, gradients were calculated from detailed temperature logs in shallow boreholes. With one exception, the heat flow west 108   

of the Nile and in northern Egypt is estimated to be low, 40~45 mW m-2, typical of a Precambrian

Platform

province. A local high, 175 mW m-2, is probably due to local oxidational heating or water movement associated with a phosphate mineralized zone.

East

of

the

Nile,

however, including the Gulf of Suez, elevated heat flow is indicated at several sites, with -2

a high of 175 mW m measured

in

Fig. (63) Temperature distribution with depth in the area of the Gulf of Suez, showing increasing temperature with depth, up to more than 300⁰C at 5000 m deep(After Abdel Zaher et al., 2011).

Precambrian

granitic gneiss approximately 2 km from the Red Sea coast. These data indicate potential for development of geothermal resources along the Red Sea and Gulf of Suez coasts. Water geochemistry data confirm the high heat flow, but do not indicate any deep hot aquifers. Microearthquake monitoring and gravity data indicate that the high heat flow is associated with the opening of the Red Sea (Morgan and Swanberg 1978/79). The most abundant solutes in all of the thermal waters are Na and Cl, while Mg, Ca, and SO4 are also prominent and the pH values are near neutral, which indicate that the solutes were mainly derived from regional marine sedimentary rocks and windblown deposits (marine aerosol and evaporate dust). Additionally, the ratio of 3He/4He in the gases emitted from the Hammam Faraun hot spring was found to be 0.256 times the atmospheric ratio (Ratm). The 3He/4He ratio in the mantle is eight times Ratm. Hence, this ratio indicates that there is excess of helium (3.2%) which may be attributed to a deeper source in the mantle (Sano et al., 1988). Sturchio and Arehart (1996) related such mantle He to the subsurface alteration due to late Tertiary volcanic eruptions. The high heat flow of the Gulf of Suez-Red Sea Rift, which is due to anomalous heated upper mantle, falls down laterally to reach the characteristic value of 46 mWm-2 at about 90 km away from the Gulf of Suez axes and 150200 km away from the northern Red Sea coast (Fig., 63). The extensional rifting in the Gulf of Suez augmented the heating and produced the broad uplifts flanking the rift (Steckler, 1985; Feinstein et al., 1996). Geothermal studies were made on the basis of the collected bottom109   

hole temperature logs of 103 deep oil wells in the Gulf of Suez, with depths ranging from 2000 to 4500 m. The blanketing effect of the overburden allows sediments to heat by conduction that causing to increase the pressure and temperature of the deeper parts of the earth. Thus, there is big difference in temperature between the surface and subsurface strata which leads to produce high value of temperature gradient (Abdel Zaher et al., 2011). The thermal data at the eastern part of Egypt indicate that the geothermal situation of the Red Sea is more complex and broader than the Gulf of Suez. Observations near to the axial trough of the Red Sea have a mean of 470mWm2 that typical associated with an active spreading center. Whereas a mean of 116mWm2 was recorded near the coast of the northern Red Sea (Boulos, 1990) that is appropriate with the estimated values at the Gulf of Suez.

2.1- Geothermal reservoirs in the Hammam Faraun and Hammam Musa regions A heat source, a reservoir, and a fluid represent the main elements in any geothermal system. The reservoir is a volume of permeable rocks from which the circulating fluids extract the heat. In the majority of cases the geothermal fluid is meteoric water, though systems near the coast may be fed by both meteoric water and seawater. It is possible that the magmatic heat source adds some water and dissolved constituents (Abdel Zaher et al., 2011). This geothermal fluid is the carrier that transfers the heat. The geothermal systems in the Gulf of Suez region represent low-temperature systems that occur in a variety of geologic units (Abdel Zaher et al., 2011). The Hammam Faraun tilted block is one of the main fault blocks in the central dip province of the Suez rift that is bounded to the east and west by major normal fault zones. These major border fault zones are in excess of 25 km long, dip steeply to the west, and have displacements up to 2–5 km (Moustafa and Abdeen, 1992; Sharp et al., 2000). The Hammam Faraun hot spring (70⁰C) flows from faulted Eocene dolomitic limestone. These geological characteristics, combined with geochemical and geophysical information, indicates that the source of the hot springs is the tectonic uplift of hotter rocks, causing deep fluid circulation through faults on the surface of the basement rock (Abdel Zaher et al., 2011). These faults allow the formation of discharge conduits for water ascending from depth after being heated and mixed with other water types. The Hammam Musa hot spring is located to the south of the Hammam Faraun hot spring, where the temperature of the emerging thermal water reaches 37⁰C and flows from faulted Miocene rocks. The geophysical interpretation of the Gulf of Suez reflects that the fault below the Hammam Musa hot spring is not only due to vertical displacement. During the early Miocene, NNE–SSW extension, oblique to the trend of the 110   

Gulf, took place. Faults in these trends frequently show a component of left lateral strike slip motion (Angelier, 1985; Moustafa and Abdeen, 1992). The conceptual model of the hydrothermal system in Hammam Musa assumes that the origin of the hot spring is this kind of oblique-slip fault (Abdel Zaher et al., 2011). Furthermore, tectonic uplift of deeper, hotter rocks below the Hammam Musa hot spring causes deep fluid circulation through faults at the surface of the basement rock. The hot water flows upward through lateral fractures and oblique-slip faults, and the main recharge of the deep water comes from two main sources: meteoric water and the intrusion of sea water. 3- Mineral Resources 3.1- Organic minerals A- Oil and Gas Three distinct oil and gas provinces were well known in Egypt; the Gulf of Suez, the Nile delta and Western Desert. The largest part of the production and reserves drives from prolific area of the Gulf of Suez. Egypt's hydrocarbons are accumulated in formations ranging in age from Carboniferous to Pliocene. The reservoirs are formed essentially by sands and sandstones and to a lesser extent by carbonates. The source rocks of hydrocarbons in the Gulf of Suez District are generally classified according to the amount and type of organic matter, the degree of maturation and the thermal alteration. The sedimentary section contains six intervals which exhibit source rock characteristics. These intervals consist of fine clastics and carbonates and are present in Carboniferous formation (Nubia B), in Upper Cretaceous carbonates (Sudr Formation), in Paleocene-Eocene deposits (Esna Shale) and in lower and Middle Miocene fine clastics (Kareem, Rudies and Belayim shales). Their content in organic, oil born matter ranges between 1.04% to 1.44% which classified as a good content, (Anon, Geology of Egypt). Abu Madi/El Qar’a is a giant field located in the north eastern part of Nile Delta and is an important hydrocarbon province in Egypt, but the origin of hydrocarbons and their migration are not fully understood. In this paper, organic matter content, type, and maturity of source rocks have been evaluated and integrated with the results of basin modeling to improve our understanding of burial history and timing of hydrocarbon generation. Modeling of the empirical data of source rock suggests that the Abu Madi formation entered the oil in the middle to upper Miocene, while the Sidi Salem formation entered the oil window in the Lower Miocene. Charge risks increase in the deeper basin megasequences in which migration hydrocarbons must traverse the basin updip. The migration pathways were principally lateral

111   

ramps and faults which enabled migration into the shallower middle to upper Miocene reservoirs (Keshta et al., 2012). Western Desert of Egypt forms the major part of the unstable shelf, located northeastsouthwest trending basin. This basin characterizes by its high oil and gas accumulations and its oil productivity about 45,000 BOPD from 150 producing wells in 16 oilfields, which represents more than one third of the oil production from the northern Western Desert of Egypt (Younes, 2012). The upper clastic and regressive series referred to as the Albian Kharita Formation and Early Cenomanian Bahariya Formation (Schlumberger, 1984). They were respectively deposited in shallow marine and fluvio-deltaic environment. They consist of consolidated sandstone with occasional coal seams and minor intercalations of shales. Oil and gas is produced from sandstones of the Bahariya Formation in the Alamein and Abu Gharadig fields. In the Upper Cretaceous section, the thickness of which may exceed 2300 m, two main rock units have been recognized, the Abu Roash Group and the Khoman Formation (El Diasty et al., 2012). B- Coal Bearing formations Lithostratigraphically, the Safa Formation belongs to the upper clastic unit of Middle Jurassic age. It consists of 215 m thick carbonaceous, banded, silty sandstones with a few earthy grey limestones. The ratio of lime/clay/sand is 29:37:34. The sandstones are crossbedded, ripple-marked, concretionary and with occasional iron sulphides. The lower part of this formation includes the economic coal beds of Gebel Maghara. The Jurassic coal deposit in the Maghara area, Sinai, Egypt contains at least 11 coal seams of lenticular shape (Issawi et al., 1999; Baioumy, 2009). These coals are interpreted as having been deposited in lakes or lagoons adjacent to the coastline (Jenkins, 1990). Hassaan et al., (1992) concluded that the Safa Formation was deposited from acid tropical soils in continental estuarine to very shallow marine environments interrupted by fluviomarine or continental phases. The structure of the Maghara area is an asymmetric doubly plunging anticline; its direction is concordant with the Syrian Arc Structural trend through Northern Sinai (Mostafa and Younes 2001). The thickness of the main coal seams ranges from 130 cm to 2 m and are underlain and overlain by thin black shale beds. Mineralogical analysis indicated that this coal is characterized by low mineral matter with traces of quartz in some samples. However, coal ash is made up of quartz with traces of calcite, anhydrite, and hematite. Analysis of coal rank parameters indicated that the Maghara coal can be classified as medium volatile bituminous coal. The high sulfur contents and the relatively high proportion of pyritic sulfur suggest a possible marine transgression after the deposition of precursor peat. This interpretation is supported by 112   

the relatively high B contents. The relatively high Ge in the Maghara coal could be attributed to an infiltration of Ge enriched water from the surrounding siliceous sediments probably during diagenesis. The high Au contents were contributed to an Au-rich provenance of the ash contents of this coal (Baioumy, 2009). 3.2- Inorganic minerals 2.1- Talc deposits The Atshan, Abu Gurdi, Darhib and Kashira talc deposits of the Eastern Desert of Egypt are located 18–60 km west of the Red Sea and occur within a 60 km radius. Of the 35 reported small talc occurrences in the Eastern Desert and Sinai, the Atshan and Darhib mines were the main talc producers. Atshan mine, the largest producer, was in operation intermittently from 1962 to 1992 and has an estimated reserve of approximately 60,000 tones of talc (Schandl et al., 2002). All four talc deposits occur within mafic, intermediate and felsic volcanic rocks and the talc ore bodies represent a distinct lithological unit within the volcanics. Shear zones and intrusive rocks are common at all four locations and the deposits had a protracted, complex metamorphic history. The talc crystallized from the replacement of siliceous carbonate beds locally intercalated with clastic sediments. The talc deposits may represent relict fragments of an ancient, regionally extensive carbonate horizon within the arcrelated metavolcanics. The talc-rich rocks, which contain relict carbonate, serpentinized olivine and tremolite, precluding mafic or felsic igneous protoliths. The deposits were locally affected by contact metamorphism, giving rise to pyroxene-hornfels and granulite facies assemblages, and by regional metamorphism which produced greenschist-amphibolite grade assemblages. Disseminated sulfides commonly occur in the talc-tremolite-rich rocks (having low Al2O3 concentrations), suggesting that the metals were probably present in the original carbonate beds, but were remobilized and reconcentrated during the various metamorphic events (Schandl et al., 2002). 2.2- Gold, magnetite and zircon: Gold deposits and occurrences located in the Nubian Shield have been known in Egypt since Predynastic times. These are stratabound deposits and non-stratabound deposits hosted in igneous and metamorphic rocks, as well as placer gold deposits. The stratabound deposits are hosted in island arc volcanic and volcaniclastic rocks of comparable composition formed in ensimatic island arcs. They are thought to have formed by exhalative hydrothermal processes during the waning phases of sub-marine volcanic activity. Stratabound deposits are sub-divided into three main types: gold-bearing Algoma-type Banded Iron Formation, goldbearing tuffaceous sediments and gold-bearing volcanogenic massive sulphide deposits. Non113   

stratabound deposits occur in a wide range of igneous and metamorphic rocks. They were formed during orogenic and post-cratonization periods by mineralizing fluids of different sources. Non-stratabound deposits are divided into veintype mineralization, which constituted the main target for gold in Egypt since Pharaonic times, and disseminated-type mineralization hosted in hydrothermally altered rocks (alteration zones) which are taken recently into consideration as a new target for gold in Egypt. Placer gold deposits are divided into modern placers and lithified placers. The former are sub-divided into alluvial placers and beach placers. Conglomerates occurring on or near ancient eroded surfaces represent lithified placers (Botros, 2004). The stream sediments of Dahab area, southeastern Sinai, Egypt, are immature as indicated by poor sorting and other mechanical parameters. They are derived from Precambrian basement rocks, which are mostly represented by granitic rocks in addition to lesser amounts of volcanics and gabbros. The mineralogical investigation revealed that these sediments contain considerable amounts of placer gold, Fe–Ti oxides and zircon. The concentrated Fe–Ti oxides comprise homogeneous magnetite and ilmenite in addition to ilmeno-magnetite, hemoilmenite and rutile–hematite intergrowths. Isodynamic separation of some raw samples of size = 1 mm revealed that up to 15.12% magnetic minerals can be recovered. Zircon shows remarkable variations in morphology, colour, chemistry and provenance. U-poor and U-rich varieties of zircon were discriminated containing UO2 in the ranges of 0.04–1.19 and 3.05– 3.68 wt.%, respectively. REE-bearing minerals comprise monazite, allanite and La-cerianite (Surour et al., 2003). 2.3- Platinum-group minerals Serpentinites are the predominant components in the ophiolitic mélange, either as matrix or as variably sized blocks, and are derived from harzburgite and subordinate dunite. The central Eastern Desert chromitites have a wide compositional range from high-Cr to highAl varieties, whereas those of the southern Eastern Desert have a very restricted compositional range. The Cr of spinel ranges from 0.5 up to 0.8 in the former, while it is around 0.8 in the latter. Platinum-group element (PGE) mineralization has been recently reported in podiform chromitites from the late Proterozoic Pan–African ophiolite of the Eastern Desert of Egypt. The populations of platinum-group minerals (PGM) in the Eastern Desert chromitites are quite distinguishable; they are mainly sulfides (Os-rich laurite) in the former, and Os–Ir alloy in the latter (Ahmed, 2007).

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2.4- Uranium isotopes The economic iron ore deposits of Egypt are located at Bahariya Oasis in the Lower Middle Eocene limestone. The main iron minerals are goethite, hematite, siderite, pyrite, and jarosite. Manganese minerals are pyrolusite and manganite. Gangue minerals are barite, glauconite, gibbsite, alunite, quartz, halite, kaolinite, illite, smectite, palygorskite, and halloysite. Geochemical comparison between the ore and the Nubia sandstone showed that the ore is depleted in the residual elements (Al, Ti, V, and Ni) and enriched in the mobile elements (Fe, Mn, Zn, Ba, and U) which indicates that the Bahariya iron ore is not a lateritic deposit despite the deep weathering in this area. On the other hand, the Nubia sandstone showed depletion in the mobile elements, which demonstrates the leaching process in the Nubia Aquifer. The presence of such indicator minerals as jarosite, alunite, glauconite, gibbsite, palygorskite, and halloysite indicate that the ore was deposited under strong acidic conditions in fresh water. Isotopic analyses of the uranium in the amorphous and crystalline phases of the ore, in the country rocks, and dissolved in the Nubia Aquifer water, all support the conclusion that U and Fe were precipitated together from warm ascending groundwater. U and Fe display strong co-variation in the ore, and the

234

U/238U activity ratio of the newly

precipitated U in the country rock and the leached component of U in the groundwater are identical. There is only slightly more uranium in the amorphous phase than in the crystalline and only a slightly lower

234

U/238U activity ratio, suggesting that the iron in the two phases

have a similar origin (Dabous, 2002). 2.5- Phosphate deposits Phosphorite deposits in Egypt, known as the Duwi Formation, are a part of the Middle East to North Africa phosphogenic province of Late Cretaceous to Paleogene age. Phosphatic grains in these deposits are classified into phosphatic mudclasts and phosphatic bioclasts (fig., 64a). Phosphatic bioclasts are subdivided into fish bone fragments and shark tooth fragments. All phosphatic grains are composed of francolite (Baioumy et al., 2007). The Duwi Formation overlies a fluvial shale sequence of the middle Campanian Qusseir Formation, and is overlain by the deeper marine shales and marls of the middle Maastrichtian Dakhla Formation. Thus, deposition of the Duwi Formation represents an initial stage of the late Cretaceous marine transgression in Egypt (Fig., 64b). The precise age of the Duwi Formation is poorly known, and generally considered as either late Campanian to early Maastrichtian based on paleontological evidences (Glenn and Arthur, 1990).

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According to Baioumy and Tada (2005), the Duwi Formation in the Red Sea, Nile Valley, and Abu-Tartur areas overlies nonmarine, varicolored shale of the middle Campanian

Qusseir

Formation,

and

is

comformably overlain by marine, laminated, gray, foraminefera-rich shale of the middle Masstrichtian Dakhla Formation. The Duwi Formation is subdivided into four members based on its lithology. The lower member is composed of coarse phosphatic sandstone in the Abu-Tartur area whereas it is composed of quartzose sandstone and siliceous shale in the Nile Valley and Red Sea areas. The middle

member

is

composed

of

soft,

Fig. (64a) Geological map of Egypt with the localities of phosphate areas (Baioumy et al., 2007).

laminated, organicrich, black shale in the three areas. The upper member is composed of coarse glauconitic sandstone at Abu-Tartur area, phosphatic sandstone in the Nile Valley area, and phosphatic sandstone and oyster fragment-rich calcarenite in the Red Sea area, respectively. The uppermost member is composed of hard, massive grayish brown to gray shale in the three areas. Individual phosphorite beds in the Duwi Formation range in thickness from a few millimeters to tens of centimeters. Thicker phosphorite beds are formed by amalgamation of thinner

Fig. (64b) Correlation of columnar sections of the Duwi Formation and equivalent phosphate-bearing formations in the studied localities (Baioumy and Tada, 2005).

individual beds. The thickest accumulation of minable phosphorites occurs in the lower member in Abu-Tartur area where the phosphorite beds locally amalgamate to form a single seam averaging approximately 12 m thick. One common feature of nearly all Duwi phosphorites is extensive bioturbation. As a result, most of the phosphatic beds appear massive and internally structureless (Baioumy et al., 2007).

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2.6- Gypsum deposits Gypsum crusts are recorded only capping the Middle Eocene carbonate rocks that are interbedded with thick gypsiferous shale beds in the north central part of Egypt. In Girza area, the gypsum crusts are capping different stratigraphic formations, the oldest of which is the Middle Eocene Ravine beds (Gehannam Formation) that consist of gypsiferous shale, marl, limestone and sandstone (Strougo and Haggag, 1984). The Ravine beds form the inselberg of Girza (Gebel Gerzah) that reach a height of 99 m and overlooking the Fayum Depression that reach a depth up to 45 m below sea level at Lake Qarun (Aref, 2003). The Quaternary littoral plain of the Red Sea between Ras Shukeir and Ras Banas comprises a narrow pediment of gently sloping alluvial fans, fringing the Neogene hills and the raised edge of the Precambrian range. This part of the African Shield, of up to 2000 m, constitutes the so-called Red Sea Hills, separating the Eastern Desert from the Red Sea coastal area Quaternary evaporite sites examined on the western side of the Red Sea contain two contrasting sedimentary series, of Late Pleistocene age, respectively composed of reefal carbonate and salina gypsum deposits (Orszag-Sperber et al., 2001). The Late Pleistocene (MIS 5.5) reefs constitute the lowest subcontinuous carbonate cliff fringing the present reefal shoreline, whereas the evaporitic unit is located a few hundred meters behind the eroded back of the MIS 5.5 reef-and-beach relief and is interpreted as subaqueous gypsum deposited in salinas (Orszag-Sperber et al., 2001). The Ras Shukeir Holocene evaporites are located on the western shoreline of the Gulf of Suez (about 3 km west of Ras Shukeir, 35 km southwest of Ras Gharib city. They are separated from the sea by a 1-km-wide barrier ridge of sandy, bioclastic limestones and mudstones probably of Plio-Pleistocene age (Purser et al., 1987). The dry sabkha plain is covered with gypsum and halite crusts that exhibit tepee structures. Near its southern extremity several small salinas occur; these have an average depth of 1 m and were formed along NW-SE fault lines (Wali et al., 1986). The Holocene evaporite sequence in the Ras Shukeir area conformably overlies marine shell banks and cross-bedded to graded-bedded beach sands and gravels. The evaporite sequence is represented by gypsum-anhydrite layers that are interbedded with mudstone layers. Field and petrographic investigations of the evaporite deposits revealed two facies types, laminated evaporite facies (primary) and nodular to enterolithic anhydrite facies (diagenetic) (Aref, 2003). The studied evaporite rocks crop out in three isolated hills separated by small wadis. The hills are located on the upthrown side of a WNW-ESE fault line, whose downthrown side is believed to be located beneath the presentday sabkha. This is evidenced by the sharp termination of the evaporate exposure toward the 117   

sabkha and the presence of 50 cm low-lying evaporite hills north of the fault line. The evaporite rocks overlie uplifted beach terraces which contain rock fragments from the surrounding Quaternary gravels (thus the evaporites are younger in age than the Quaternary gravels) (Aref et al., 1997). 2.7- Limestone deposits The upper Oligocene Wadi Arish Formation is composed of a carbonate-dominated succession at Gebel Risan Aneiza (Sinai). The 77-m-thick unit disconformably overlies Jurassic to lower Cretaceous carbonates and is subdivided into three members, comprising six lithofacies units. The lower Wadi Arish member contains three units, a gypsiferous sandstone unit, overlain by two limestone units. The middle Wadi Arish member is represented by a conspicuous marl unit that is overlain by two upper limestone units of the upper Wadi Arish member (Kuss and Boukhary 2008). The Mio-Pliocene sedimentary sequence is widely distributed in the Cairo–Suez road district and along the western coast of the Gulf of Suez. It unconformably overlies either the Middle or upper Eocene rocks and unconformably underlies the Pleistocene rocks. this sequence was divided (exposed in a district between Gabal Ataqa and El Galala El Bahariya plateau, located at the northwestern side of the Gulf of Suez) into three series: the Lower, Middle and Upper Miocene. This division was differentiated it into three formations; these are from older to younger, Sadat, Hommath and Hagul, respectively. They established the Hagul formation to represent an Upper Miocene clastic/limestone sequence of about 22 m thick measured near the entrance of Wadi Hagul (Khalaf and Gaber 2008). Gebel Umm Hammad in the Red Sea Mountains east of Quseir, Egypt, enabling the widening of joint controlled openings in the Thebes Limestone. The valley and Gebel, together somewhat more than 5 km wide, run parallel to the Egyptian Red Sea coast for a distance of about 30 km, 25 km inland from Quseir. The hogback consists of Thebes Limestone, the last marine deposit before Red Sea proto-rifting began in Oligocene times. The Thebes Limestone formation contains several beds, 0.5 to 2 m thick, with intercalated ‘conglomerates’ of rounded chert nodules (Moeyersons et al., 1999). Thebes Formation forms an extensive carbonate platform on the southern margin of Tethys, outcropping along the Nile Valley and over large areas of the Western Desert of Upper Egypt. It has an extensive literature, but its biostratigraphy, depositional environments and sequence stratigraphy are still not well integrated on a regional scale. The type section of the Thebes Formation is Gebel Gurnah, on the west bank of the Nile opposite Luxor. The 118   

maximum thickness is preserved in the area of the high peak EI Qurn, which overlooks the Valley of the Kings. A 350 m section was logged in this area; samples were collected for XRD analysis, microfauna and nannofossils. The Thebes Formation comprises mainly chalk and chalkstone (chalk with secondary interstitial cement). Layers of chert nodules are common, and siliceous limestones become increasingly common in the upper part. Lithofacies range from calcareous and dolomitic claystone through chalk to nodular limestone. These largely reflect relative water depths. Thin bioclastic limestones with larger foraminiferids represent episodes of reduced sedimentation (King et al., 2011). 2.8- Shale formations Carbonaceous shales have a wide distribution on the Egyptian surface and in subsurface sedimentary sequences e.g. in sediments of predominantly Carboniferous, Jurassic, Cretaceous, Paleocene and Eocene age. The carbonaceous and black shales in Egypt gained interest since five decades when the phosphorite deposits were discovered and exploited. The phosphorites are intercalated with and capped by black shales that contain considerable amounts of organic matter and are enriched in trace elements, which may be of economic potential. Black and carbonaceous shales of Duwi Formation, Dakhla Shale and Esna Shale of Upper Cretaceous and Lower Tertiary age in Abu Tartur, Nile Valley and Quseir, detrital smectite is the dominant clay mineral in addition to minor kaolinite and chlorite contents. The smectite content gives evidence of considerable marine influence during the sedimentary processes in South Egypt in comparison to Sinai. The Duwi Formation in Abu Tartur was deposited in a shallow and restricted marine environment under prevailing reducing conditions. There, the Campanian/Maastrichtian transgression was interrupted by multiple regressive phases which caused intensive reworking of sediments and enrichment of the phosphate layers in this formation. The Duwi Formation is characterized by absence of foraminifera, compared to the abundance of foraminifera's assemblages in Dakhla and Esna shales which suggest open marine environments and prograding marine transgression during the deposition of these formations (Temraz, 2005).

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Internet Sites 1- http://www.classroomatsea.net/general_science/plate_tectonics/tectonics_intro.html. 2- http://geophysics.ou.edu/geol1114/notes/structure/structure.html 3- http://www.ga.gov.au/oceans/sa_krgln_Sdmnt.jsp 4- http://www.classroomatsea.net/general_science/plate_tectonics/tectonics_intro.html 5- http://ehsgeowiki.wikispaces.com/Ocean+Trench 6- http://geology.com/nsta/divergent-plate-boundaries.shtml 7- http://www.indiana.edu/~geol116/week7/week7.htm 8- http://geology.com/nsta/convergent-plate-boundaries.shtml 9- http://rainforestgirl.edu.glogster.com/plate-boundaries/ 10- http://www.kidsgeo.com/images/transform-boundary.jpg 11- ttp://www.visionlearning.com/library/module_viewer.php?mid=66 12- http://pubs.usgs.gov/fs/1999/fs110-99/ 13- http://sio.ucsd.edu/png/science/ 14- http://www1.fccj.org/pacrews/oceans2.htm 15- http://elearning.stkc.go.th/lms/html/earth_science/LOcanada6/604/7_en.htm 16- http://www.earth.northwestern.edu/people/seth/202/lectures/Platetect/Continentalevl/Ima ge130.gif 17- http://www.earth.northwestern.edu/people/seth/202/lectures/Platetect/Continentalevl/Ima ge131.gif 18- http://www.eia.gov/cneaf/solar.renewables/renewable.energy.annual/backgrnd/fig19.htm 19- http://www.mpoweruk.com/geothermal_energy.htm 20- http://my.opera.com/talatkm/blog/index.dml/tag/renewable%20energy. 21- http://www.princeton.edu/~achaney/tmve/wiki100k/docs/Mediterranean_Sea.html. 22- http://www.mpoweruk.com/geothermal_energy.htm. 23- www.nrm.qld.gov.au/ factsheets/pdf/mines/m4.pdf.

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