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 

I  

Contents

Subject P.No. SUMMARY 1 CHAPTER I: SEDIMENTARY BASINS 9

1- Sedimentary Basins Definition 9 2- Origin and Mechanisms of Basins Formation 10 Earth's Crust Components 12 A- The continental shelf 12 B- The continental slope 13 C- The continental 13 D- Deep-Ocean Trenches 3 E- Plate margin (plate boundary 13 E.1- Convergent boundaries (subduction zones) 14 E.2- Continental crust 14 E.3- Oceanic crust 15 E.4- Triple junctions 15 3- Sedimentary Basins evolutions 16 Divergent boundaries (ocean ridges) 17 Convergent boundaries 18 Transform boundaries 19 Passive continental margins 20 Active continental margins 21 Foreland basin Systems 22 Pripheral or Pro-foreland basin 23 Retroarc or Retro-foreland basin 23 Dynamic topography 24 4- Classification of Sedimentary Basins 25 I- Tectonic Basin Classification 25 I.1- Continental or interior sag basins 27 I. 2- Continental graben structures and rift zones form narrow elongate basins bounded by large faults

28

I. 3 - Failed rifts or aulacogens 28 I. 4- Passive margin basins 29 I. 5- Oceanic sag basins or nascent ocean basins 29 I. 6- Basins related to subduction 30 I. 7- Terrane-related basins 30 I. 8- Basins related to collision 31 Retroarc or intramontune basins 33 Pannonian-type basins 33 I. 9- Strike-slip and wrench basins 33 II- Pre-, Syn-, and Post-Depositional Basins 34 1- Post-depositional basins 34 2- Syn-depositional basins 34 3- Pre-depositional basins 35 5- Basins Morphology 36 6- Depositional Environments 38 1- Continental Sediment Environments 39 1.1- Glacial Environments 39 1.2- Aeolian Environments 39 1.3- Rivers and Alluvial Fans 40 1.4- Lakes and Lacustrine Environments 42 2- Marine Sediment Environments 44 2.1- Marine deltas 45

II  

2.2- Clastic Coasts and Estuaries 46 2.3- The beach 46 2.4- Coastal plains 47 2.5- Beach barriers 48 2.6- Shallow Marine Carbonate and Evaporite Environments 49 2.7- Adjacent sea basins and epicontinental seas 49 2.8- The shallow seas and continental shelf sediments 50 2.9- Deep-sea basins 50 CHAPTER II: GEOTHERMAL ENERGY IN THE SEDIMENTARY BASINS 52 1- Geothermal Gradient 54 2- Effect of the geothermal energy on hydrocarbon maturation 55 3- Geothermal energy utilizations 56 3.1- Hydrothermal Systems - Geothermal Aquifers 57 3.2- Hot Dry Rocks (HDR) Enhanced Geothermal Systems (EGS) 57 3.3- Geothermal energy in contemporary balneotherapeutics and Tourism

59

4- Healing and therapeutic value of geothermal waters 59 4.1- Therapeutic tourism 60 4.2- Geothermal Electricity Production around the world 60 CHAPTER III: MINERAL RESOURCES OF THE SEDIMENTARY BASINS 62 I- Organic Mineral Resources 63 I.1- Oil and Natural Gas Resources 63 1.1- Sedimentary basins and petroleum formation in the Middle East

63

1.2- Petroleum prospectivity of the principal sedimentary basins on the United Kingdom Continental Shelf

65

1.3- Prospectivity of the sedimentary basins of Irish Sea 66 I. 2- Coal bearing formations 66 2.1- Australia 66 2.2- India 66 II- Inorganic Mineral Resources 66 II.1- Volcanogenic massive sulphides (VMS) 67 II.2- Metaliferous Oxides 67 II.3- Metallic and Gem Minerals in Placer Deposits 67 II.4- Evolution of a Mineralized Geothermal System, Valles Caldera, New Mexico, USA

68

II.5- Mineral Resources of the Western Canada Sedimentary Basin 68 II.6- Mineral Resources of the Australian Sedimentary Basins 69 6.1- Heavy minerals 69 6.2- Bauxite 69 6.3- Sedimentary phosphate deposits 69 6.4- Other Metals 69 CHAPTER IV: MEDITERRANEAN SEA 70 I- Mediterranean Geosynclinal Belt 70 II- Origin and evolution of Mediterranean geosyncline 71 III- Paleoenvironmental analysis 73 IV- Mediterranean basins 74 IV. 1- Tectonic Settings of Eastern Mediterranean basin 75 IV.2- Tectonic Settings of the Western Mediterranean 76 V- Origin and Tectonic History of Mediterranean Sub-basins 79 V.1- The Levantine Basin 79 V.2- Aegean Sea basin 80 V.3- Adriatic Sea basin 82 V.4- Ionian Sea basin 84 V.5- The Tyrrhenian Sea 86

III  

V.6- The Alboran Sea 88 V.7- The Algerian Basin 90 VI- Geothermal Potentials and Uses of the Mediterranean 94 VI. A- Geothermal potentials 94 A.1- Geothermal Resources in Foreland Environments 95 A.2- Thermal Coastal Springs 95 VI. B- Geothermal Uses 96 B.1- Electrical production 96 V- Mineral Resources in the Mediterranean Region 97 V.1- Organic minerals (Oil – Natural Gas – Coal) 97 1. A- Oil and natural Gas resources 97 A.1- Lavantine basin 97 A.2- Adriatic Sea 97 A.3- Neogene petroleum system at Alboran - Algerian Basins

98

1.B- Coal Bearing Formations 98 V.2- Inorganic mineral resources 99 CHAPTER V: EGYPT (Genius of the Place) 102 1- Sedimentary Basins of Egypt 102 1.1- Nile Delta 102 1.2- Eastern Desert 104 1.3- Red Sea Rift Valley 105 1.4- Western Desert 106 2- Geothermal Regime of Egyptian Basins 108 2.1- Geothermal reservoirs in the Hammam Faraun and Hammam Musa regions

110

3- Mineral Resources 111 3.1- Organic minerals 111 A- Oil and Gas 111 B- Coal Bearing formations 112 3.2- Inorganic minerals 113 2.1- Talc deposits 113 2.2- Gold, magnetite and zircon 113 2.3- Platinum-group minerals 114 2.4- Uranium isotopes 115 2.5- Phosphate deposits 116 2.6- Gypsum deposits 117 2.7- Limestone deposits 118 2.8- Shale formations 119 References 120

1  

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 deep-

ocean’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

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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., Mid-

Atlantic 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

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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

4  

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.

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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;

6  

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 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.

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 supra-

subduction 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

7  

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

8  

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 carbonate-

dominated 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 proto-

rifting 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.

9  

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.

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11  

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 brittle-

ductile 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).

12  

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, 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

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

13  

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

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

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

14  

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 beneath the

continental mass of South

America (Gutscher, 2002).

Convergent plate boundaries are

defined by earthquake

hypocenters that lie in 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

Fig., (4) Subduction zone. Source: http://geology.com/nsta/ divergent-plate-boundaries.shtml

15  

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).

16  

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

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

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

17  

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 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 lithosphere is

produced at a ridge the

supercontinent splits and moves

apart on each of the ridge

flanks. Divergent boundaries

occur when plate are rifted apart

and begin to move 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

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

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

18  

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 fast-

spreading ridges. Model studies sag as oceanic lithosphere thickens with distance from a ridge

axis, horizontal extensional stresses can produce the axial topography found on slow-

spreading 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

(Condie, 2003). .

Convergent boundaries form

where plates converge. One

plate is usually subducted

beneath the other at a

convergent plate boundary.

Convergent boundaries may be

of different types, depending on

the types of lithosphere

involved. These results in a wide

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

Source: http://geology.com/nsta/convergent-plate-boundaries.shtml

Fig., (7) Convergent Tectonic motion between continent and ocean

plate. Source: http://rainforestgirl.edu.glogster.com/plate-boundaries/

19  

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 boundaries are 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 and reflect

inhomogeneous fracturing of the

lithosphere. Transform faults, like

ocean ridges, are characterized by

shallow earthquakes (< 50 km deep).

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

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

20  

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

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

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).

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

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

21  

Active continental margins are

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

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

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). 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):

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

Fig. (9d) Active continental margin. Source: http://elearning.stkc.go.th/lms/html/earth_science/LOcanada6/604/7_en.htm.

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

22  

• 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 –

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

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

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

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

23  

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 belt once steady- state is

achieved, and are consequently

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

 

 

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

24  

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 relative

contributions of shallow

asthenospheric flow versus

deeper flow 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, dynamic

displacement of the surface

topography, usually referred to

as dynamic 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).

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

25  

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

26  

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 Underlying

crust Style of tectonics

Basin Characteristic

1 Continenetal or interior sag basins

Epicontinental basins, infracratonic basins.

Continenetal Divergence Large areas, slow subsidence

2 Continenetal or interior fracture basins

Graben structures, rift valley, rift zones, aulacogens

Continenetal Divergence Relatively narrow basins, foult bounded, rapid subsidence during early rifting.

3 Basins on passitve continenetal margins, margin sag basins

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

Transitional Divergence + shear

Asymmetric basins partly outobuidling of sediments, moderate to low subsidence during later stages

4 Ocean sag basins Nascent ocean basin (growing oceanic basin)

Oceanic Divergence Large asymmetric, slow subsidence

5 Basins related subduction

Deep sea trenches, Forearc basins, backarc basins, interarc basins

Oceanic Transitional oceanic

Convergence Dominantly divergence

Partly asymmetric, greatly varying depth and subsidence

6 Basins related collision Remnant basins Forland basins (periheral), retroarc basins (intramontane), brocken foreland basins, Terrane- related basins

Oceanic Continenetal Oceanic

Convergence Crustal flexuring, local convergace or transform motions

Activated subsidence due to rapid sedimentary loading. Asymmetric basins, tend to increasing subsidence, uplift and subsidence Similar to backarc basins

7 Strike-slip/wrench basins

Pull-apart basins (transtensional) and transpressional basins

Continenetal and/or Oceanic

Transform motions ± divergance or convergance

Relatively small, elongate, rapid subsidence

27  

I.1- Continental or interior sag basins

According to Einsele

(1992), basins on continental crust

are commonly generated by

divergent plate 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 occurs

predominantly in response 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 was recently

suggested that 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

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

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.

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

28  

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

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).

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).

29  

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

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. They commonly

accumulate deep-sea fan or

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

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

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

30  

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

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

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

31  

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 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

crust ("A-subduction", after

Ampferer, Alpine-type)

under the load of the

overthrust mountain belt. The

extension of these

asymmetric basins tends to

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):

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

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

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

32  

Stage one of arc-continent collision

Initial collision is characterized by overthrusting of the south- and southeastward-

facing 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 half-

grabens; 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.

33  

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

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.

 

Fig. (22) Tectonic basin classification: Collision-related basins and

strike-slip/wrench basins. (Allen and Allen, 2006). 

34  

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

35  

subsiding area, a

morphological 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 basins.

Rapid tectonic 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

simplified basin types and

certain basins may show a complex history and therefore contain pretectonic as well as syn-

tectonic or post-tectonic sediments (Allen and Allen, 2006).

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). 

36  

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 from continental

graben 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

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

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). 

37  

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 water-

filled 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 sediment.

Consequently, 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

basins are little influenced by

sediment outbuilding (Fig., 25). For example, terrigenous sediments transported into high-

energy 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

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). 

38  

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 of

sedimentary basins are

predominantly classified

according to their

depositional environment

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.

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

39  

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 of these

sediments can provide

important clues about past

climates (Nichols, 2009).

Glacial deposits are

compositionally immature and

tills are typically composed of

detritus that simply represents

broken up and powdered

bedrock from beneath the

glacier. Reworked 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

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

40  

material associated with sandy beaches and periglacial sand flats. Sands deposited in these

desert areas are characteristically

both compositionally and

mineralogically mature with

large-scale cross-bedding formed

by the migration of dune

bedforms. Oxidising conditions

in deserts preclude the

preservation of much fossil

material, and sediments are

typically red–yellow colours

(Nichols, 2009). Aeolian dust

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).

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). 

41  

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 these

circumstances, 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). Three

geomorphological zones can

be recognized within fluvial

and alluvial systems (Einsele,

1992). In the erosional zone

the streams are actively

downcutting, removing

bedrock from the valley floor

and from the valley sides via

downslope movement of

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

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).

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). 

42  

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 fining-

upward successions from a sharp scoured base through beds of trough and planar cross-

bedded, 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

43  

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

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

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

44  

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

45  

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 a ‘discrete

shoreline 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,

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).

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). 

46  

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

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

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). 

47  

(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 low-

lying 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

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

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

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

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). 

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).

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

48  

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

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).

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

49  

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

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

 

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

50  

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

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

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

51  

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.

52  

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 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 near-

average heat flows and geothermal

gradients (Allen and Allen, 2006).

Changes in physical and

chemical conditions during basin

evolution control the interaction

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

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). 

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

53  

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

54  

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)

temperature resources. The low

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

55  

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).

56  

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 circulating

through permeable

rock, to drive steam

powered electricity

generators. 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

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)

57  

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

58  

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).

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

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

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

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

59  

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

60  

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 2008a).

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

61  

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-energy-

choices/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).

62  

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

63  

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

64  

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 and flexures

reflecting deep-seated 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

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

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

65  

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

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

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

66  

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 fault-

bounded traps of Lower Triassic,

principally aeolian Sherwood Sandstone

reservoir, top-sealed by younger Triassic

continental mudstones and evaporates

(Fig., 45). Future exploration 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)

Fig. (45). Hydrocarbon fields and discoveries.

67  

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 volcanic-

associated, 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

68  

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

69  

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).

 

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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 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. 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.

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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.

72  

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

73  

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

74  

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.

75  

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). The

origin of the Eastern

Mediterranean basin

(EMB) by rifting along its

passive margins is

reevaluated. Evidence from

these margins shows that

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

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

76  

between the Tethys and

the Indian Oceans. This

process determined

profound changes in the

oceanic circulation

patterns, which shifted

global climates towards

colder conditions. The

Hellenic Arc, which has a

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 sub-

basins 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

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

77  

crust is composed of Paleozoic and pre-

Paleozoic rocks deformed by the

Caledonian and Variscan orogenic

cycles (Carminati et al., 1998a, b).

The geological evolution of the

western Mediterranean exhibits

complicated interactions between

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 of these basins,

simultaneously with 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

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,

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

78  

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 oceanic

lithosphere. Back-arc

extension occurred in the

overriding plates as a result

of slow convergence rates

combined with rapid

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

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

79  

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 Neo-

Tethys 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

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

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

80  

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 (Ben-

Avraham 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).

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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 and

Papanikolaou, 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 experiences

considerable 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

with a south or southwest

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).

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

82  

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).

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

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

83  

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

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 sub-

parallel 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).

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

84  

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

85  

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

occupies the central part of the

southern Mediterranean. Here 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 Ionian Sea perhaps

experiences the major 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

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

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 fold-and-thrust belt and Gargano-Murge foreland (Puglia, southeastern Italy) (After Graziano, 2013). 

86  

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

87  

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 Tyrrhenian

Sea and migrated

southward to the

southern 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 the

Southern Apennines

and in Sicily

accompanied by

counterclockwise block

rotations in the former

and clockwise rotations

in the latter. These

processes have been

controlled by rapid

rollback of oceanic

Ionian lithosphere

beneath the Calabrian

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

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).

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). 

88  

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

89  

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 of this

arcuate belt. The

region as a whole is

bounded to the north

and south by the

Iberian and 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

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

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

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) .

90  

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, high-

resolution 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

91  

depths over 2,000 m and is

characterized by a number

of submarine canyons that

cut across it. These

canyons 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)

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 with the

Tethysian slab roll-back

(Fig., 57b) and possibly

with slab break-off

(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

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).

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)

92  

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

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

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: south-verging Tellian fold and thrust belt After (Domzig et al., 2010). 

93  

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 slow-

rate 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).

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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.

95  

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).

96  

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).

97  

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

98  

’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.,

99  

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).

100  

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

101  

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).

102  

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

103  

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

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).

Fig., (59) Geological maps of Egypt.

104  

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 supra-

subduction 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 Arabian-

Nubian 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).

105  

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 (Kornِer 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

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

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)

106  

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 of

Mesozoic sediments.

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

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

107  

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 sub-

basins 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

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

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

108  

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 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. 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

109  

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

a high of 175 mW m-2

measured in 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 150-

200 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 bottom-

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).

110  

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

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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

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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 northeast-

southwest 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 cross-

bedded, 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

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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 arc-

related 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, gold-

bearing tuffaceous sediments and gold-bearing volcanogenic massive sulphide deposits. Non-

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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 high-

Al 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 234U/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 234U/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).

116  

According to Baioumy and Tada (2005), the Duwi Formation in the Red Sea, Nile

Valley, and Abu-Tartur areas overlies non-

marine, 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,

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

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).

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

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

117  

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 present-

day sabkha. This is evidenced by the sharp termination of the evaporate exposure toward the

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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

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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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