The territory of Estonia has been inhabited at least ten thousand years. The first tribes to settle in the area were hunters and fishers who could already use local raw materials, such as crystalline erratic boulders, gravel, sand and clay. At about 6000 BP inhabitants learned to make earthenware from clay and around 5000–4000 BP to apply carbonate rocks to building of town lets and fortified settlements. Since 1230, lime has been widely used as a binder. Red bricks, made of local clays and used as building material for strongholds and churches, provide Estonia’s historical buildings and architectural monuments with a specific geological splendour.
The first geological studies were carried out in Estonia more than 150 years ago. The long tradition of geological research in the area is due to the large and representative bedrock exposures providing excellent conditions for the study of Lower Palaeozoic rocks, and making Estonia a key region for solving several principal stratigraphic problems. The Palaeozoic rocks in Estonia enclose extraordinarily rich communities of well-preserved fossils, and a great number of new species and higher taxa have been established here. Ancient coastal formations of the Baltic Sea and relief forms, left behind by the last glaciation, are represented here more completely than in other regions. Within Estonia are found excellent examples of meteorite craters and the largest erratic boulders in northern Europe. All this makes Estonia’s geology unique in several aspects.
For centuries Estonia has served as an economic, scientific and cultural bridge between the East and the West. Already in the Middle Ages it was an arena of serious ideological conflicts. At the end of the Livonian War (1558–1583), southern Estonia fell under Polish rule. In the interest of restoring Catolicism, Jesuits opened a gymnasium in Tartu in 1583. After Estonia was taken over by Sweden, the Swedes founded a Protestant gymnasium in Tartu in 1630 to counterbalance the Jesuit school. In 1632, the Protestant gymnasium was changed into a university, which is one of the oldest and most prominent higher educational establishments in northern Europe. In the 17th century, Tartu University became a principal centre of education, science and humanistic ideas in the region. After re-opening in 1802, it developed into an outstanding centre of geological education and science in the former Russian Empire. The corresponding topics were also advanced in the Tartu (later Estonian) Society of Naturalists, founded in 1853.
In 1920, Estonian became the medium of instruction at Tartu University. In 1937, the Geological Committee of Estonia was founded. After the occupation and incorporation of Estonia into the Soviet Union, the most prominent geologists and a lot of promising young scientists left homeland and a new generation of geologists was trained. In 1947, the Institute of Geology of the Estonian Academy of Sciences was established and ten years later the Geological Survey of Estonia was founded. Both these institutions developed into important centres for geological, geophysical and environmental research in the northwestern portion of the Soviet Union and neighbouring countries.
The essential results of the research carried out during more than two centuries were summarised in multi-volume issues showing the directions and level of geological studies in Estonia (Geological Studies of the USSR, 50, Estonian SSR, Tallinn, 1968, 1972, 1973, 1974, 1977, 1984, 1987; History of Geological Sciences in Estonia, 1986). These, like most monographs in the field of geology issued in Estonia during the last decades, are in the Russian language and practically unknown to our western colleagues. Due to intensive drilling programmes and medium- and large-scale geological mapping, a lot of new geological information has been obtained. As there are currently no published general surveys on the geology and mineral resources of Estonia, the present monograph attempts to fill this gap. Its main purpose is not only to impart scientific information about Estonia’s natural environment, but to serve also for industrial and agricultural purposes encouraging the sound use of mineral resources in the present-day Estonia.
Mining of mineral resources has inflicted incurable wounds on Estonia’s nature. Another task of the present issue is to assist in drawing up main outlines of the strategy addressing improvement of the environment.
Most distinguished specialists of the Republic have participated in the compilation of this monograph. Its publishing has been made possible by the financial help of the Estonian Science Foundation (grant No. 1661), which is gratefully acknowledged. Thanks are due to the authors and all persons who have contributed to finalizing of this book. Special thanks go to Mrs. Helle Kukk for the revision of the English text, to Mr. Jüri Nemliher for the layout of this book and to Mr. Paul Pärkma for the drawings.
Anto Raukas and Aada Teedumäe
A. Raukas
The Republic of Estonia, the northernmost of the three Baltic States, is situated in the North-East of Europe, on the east coast of the Baltic Sea. The name Estonia is probably derived from Aists, the name the ancient Germans used to denote the Baltic tribes, living to the northeast of the Vistula River. In a written record the Aists (Aesti, Aestorium gentes) were first mentioned by the Roman historian Tacitus in the first century AD. The first written reference to the land of Estonians dates from 1154. On the order of Roger II, the king of Sicily, the Arab geographer and traveller Abu Abdallah Muhammad al-Idrisi designed a map of places in the world known in those times including Qalewany (Tallinn) in Astlanda (Estonia).
The territory of Estonia in nowadays boundaries extends from 57º30’34" to 59º49’12"N and from 21º45’49" to 28º12’44"E (Fig. 1). The northernmost point of Estonia is on the Island of Vaindloo (the Cape of Purekkari on the mainland), the easternmost point in the Town of Narva, the southernmost point is the Naha farmstead at Mõniste, and the westernmost point is on the Island of Nootamaa (the Cape of Ramsi on the mainland). The extreme length of the Estonian territory is 350 km from west to east, and 240 km from north to south. The length of the Estonian coastline is 3,780 km; of this 1,242 km are on the mainland and 2,540 km are divided among the islands.
Estonia has an area of 45,215.4 sq km of which 9.2% is taken up by islands and 4.6% is under inland bodies of water. Climatically, Estonia belongs to the mixed-forest subregion of the Atlantic continental region of the temperate zone, which is characterized by warm summers and moderately mild winters.
Geologically, Estonia is situated in the northwestern part of the East-European Platform. Structurally, it lies for the most part within the boundaries of the southern slope of the Fennoscandian Shield with only its extreme southwestern and southern parts forming the wings of the Baltic Syneclise and the Valmiera-Lokno Uplift, respectively.
As part of the vast East-European Plain, Estonia is a generally flat country (Photo 1), where uplands and plateau-like areas alternate with lowlands, depressions and large valley-like forms. The average height above sea level is approximately 50 m, relative heights of landforms do not as a rule exceed 20 m, being only seldom 50 and more metres. About 40 per cent of Estonia’s territory is at an absolute height of 50 to 100 m, and only one tenth has an elevation over 100 m above sea level (Fig. 1). The highest point in Estonia, the Suur Munamägi Hill (nearly 318 m), is located in the Haanja Heights.
Estonia displays a large variety of landscapes (Fig. 2). The northern part of the country consists of an extensive limestone plateau (Fig. 2), the northern edge of which forms a steep escarpment (Photo 2), known as the North-Estonian Klint (relative height up to 56 m). The narrow Fore-Klint Coastal Plain is situated in front of the Klint. The highest areas in the northern part of Estonia are the Pandivere Upland (166 m a.s.l.) and the Jõhvi Upland (81 m a.s.l.). To the south of the Pandivere Upland lies the gently sloping Vooremaa watershed (the Saadjärv Drumlin Field, with elevations up to 144 m a.s.l.).
Relatively high areas of North Estonia border on the Kõrvemaa and Alutaguse lowlands. To the south-west of the Pandivere Upland lies the Central-Estonian Plain which, gently sloping, passes over into the Võrtsjärv Depression. The Alutaguse Lowland turns into the Peipsi Depression.
In western Estonia the absolute height seldom exceeds 20 m and large areas are entirely flat. This is the region of the lowlands of West Estonia and West-Estonian (Moonsund) Archipelago. Some small elevations are Kõpu (63 m a.s.l.), Middle-Saaremaa (54 m), Sõrve (36.6 m) and Tõstamaa–Varbla (44 m); the scarps of the islands of Saaremaa and Muhu and those in the western part of the mainland form the West-Estonian Klint (up to 21 m a.s.l.).
In South Estonia the topography is more varied and differences in the altitude are greater than elsewhere in Estonia. The area has four topographic highs (Fig. 2): Sakala (up to 146 m a.s.l.), Otepää (217 m), Karula (137 m) and Haanja (318 m). They are separated from one another by the Valga and Hargla depressions and Võru Valley. The South-Estonian medium-height terrain (Ugandi Plateau, 40–100 m a.s.l.) is occupied by the South-West Estonian Plain.
The largest relief forms — plateaus, uplands, lowlands, depressions, the North-Estonian and West-Estonian escarpments (Aaloe & Miidel 1967) were formed in Pre-Quaternary times as a result of the long-term continental erosion (Tavast & Raukas 1982). Monoclinal bedding of bedrock strata and their different resistance to erosion resulted in the questa-like ancient topography (Orviku 1955). During all ice ages glacial erosion prevailed in North and West Estonia. These areas are characterized by a thin Quaternary cover and wide distribution of alvars against the background of Estonia’s generally flat topography. The erosional relief forms here are represented by both small (glacial scratches, etc.) and large (rock drumlins, hollows and troughs of glacial ploughing) ones.
In the transitional zone between the prevailing glacial erosion and accumulation areas in central Estonia, the most characteristic relief forms are of the erosional-accumulative type. Among them are drumlins (Photo 3) and drumlin-like ridges, including megaflutings, which may reach 13 km in length, and 80 m in height (Saadjärv Drumlin Field). The accumulation area in southern Estonia features gently sloping and undulating till plains, and morainic hills, with the latter being especially common on accumulative insular heights (Otepää, Haanja). In places, dump and push moraines stretch some tens of kilometres in length with relative heights up to 50 m (West-Saaremaa Elevation).
Glaciofluvial accumulative relief forms are widely distributed in Estonia, with classic eskers and kame fields formed, as a rule, in passive or dead ice (Karukäpp & Raukas 1976). Radial eskers are most common on the Pandivere Upland (Photo 4) and marginal eskers on the West-Estonian Lowland (Raukas et al. 1971). Fluvio- and limnoglacial kames either form separate fields or are scattered in hilly topography. As for the genesis, the glaciofluvial gravel and sandy plains are for the most part glaciofluvial deltas or outwash deltas. Less frequent are kame and glaciofluvial terraces and outwash cones.
Genetically and morphologically, the valleys of glacial meltwater discharge are diverse. These relief forms are most typical of southern Estonia where in places they form orthogonal valley systems. They include both radial and marginal valleys, some of which are formed under the ice (e.g. rills of discharge), while others came into being by epigenetical superimposing on subglacial topography under the conditions of jointed passive or dead ice, but also due to glacial breaks and intense joining of ice-dammed lakes. Glacial meltwaters often flowed along ancient valleys which had developed before the last glaciation (Tavast & Raukas 1982).
Characteristic of glacial terrain are also funnel- and saucer-shaped closed depressions — kettle holes, formation of which is associated with the melting of buried dead ice blocks (glaciokarst). Undoubtedly, in many cases, the process came to an end in the Late-glacial or at the beginning of the Holocene. However, it seems that some of the kettle-holes formed considerably later, with the process having started in the Boreal and coming to an end only in the Atlantic climatic period (Raukas & Rõuk 1995). Quite often kettle-holes are filled with peat, the thickness of which may reach 17 m.
In all stages of deglaciation considerable areas in front of glacier margins were occupied by glaciolacustrine basins of different ages (Raukas 1992a). These bodies of water have left behind deposits (mainly varved clays) and coastal relief forms (abrasional scarps, beach ridges, etc.) which are traceable at different levels, such as those on the slopes of the Otepää and Haanja heights.
The extensive glaciolacustrine plains, which were of great landscape-forming significance, occur only in the lower parts of the territory, particularly on the Alutaguse, Võrtsjärv, Peipsi and Kõrvemaa lowlands (Raukas et al. 1971). Due to the wide distribution of clayey deposits on the surfaces, modest absolute height and unfavourable drainage conditions these plains have undergone paludification, and many of them have turned into bog plains.
The effects of a variety of coastal processes can be found along both modern (Orviku 1974) and ancient coasts of the Baltic Sea (Raukas et al. 1965) and large lakes (Raukas & Tavast 1989). These features include wave-cut notches, scarps, abrasional platforms and plains, accumulative terraces, spits, barrier beaches, arrow-shaped spits, tombolos, bars, beach ridges, etc.). The development of the largest ancient coastal formations is connected to the transgressive phases of the Baltic Sea development (Raukas 1966).
More prominent aeolian relief forms (ridge-like longitudinal dunes, parabolic dunes, etc.) are spread along the ancient transgressive coastlines of the Baltic Sea. The height of dunes seldom exceeds 20 m. Formation of higher dunes was hindered by the small supply of sand, humid climate, continuing uplift of the Earth’s crust and by several other circumstances (Raukas 1968). Beside contemporary and ancient coastal dunes there are also inland aeolian formations.
Karst topography and underground features (Photo 5), resulting from the solution of rocks and leaching processes, are common on the outcrops of carbonate rocks in northern, western and central Estonia, and of limited distribution in the southeasternmost part of the Republic (Heinsalu 1977). However, because of the relatively small thickness of soluble rocks, low absolute heights of the terrain, short duration of the post-glacial evolution of the territory and for several other reasons, karst relief forms have rather modest dimensions in Estonia. Nevertheless, several features have been identified here, including karren, karst holes, open jointings, karst relicts, funnel-sinks of absorption, angular subsidence sink-holes, underground and disappearing rivers, small caves, and other karst phenomena. In some places one may find peculiar temporary karst lakes.
Beside karst caves there are suffosion caves (Heinsalu 1987). For the most part, they are found along the slopes of the ancient South-Estonian valleys and scarps of the North-Estonian Klint. Calcium carbonate precipitates out of spring water forming dome-shaped travertine terraces on valley slopes.
Gravitational relief forms composed of debris are for the most part distributed at the foot of the North-Estonian and West-Estonian klints. In the hilly topography of southeastern Estonia, deluvial processes are ongoing due to human impact.
Biogenic (organogenic) relief forms were formed by the plant growth (phytogenic forms) and animal activities (zoogenic forms). The largest phytogenic relief forms having great landscape-forming significance are the bog or telmatogenic plains, with their hummocks, mochezhinas and other relatively small features (Masing 1977).
A. Rõõmusoks, V. Puura, A. Raukas & E. Mark-Kurik
The essential natural resources of Estonia have been used for centuries. The first geological studies were carried out in the 17th century by members of staff at Tartu University. Publications of that time include G. Mancelius’ paper on earthquakes (1619) and M. J. Herbinius’ survey of waterfalls (1678). After reopening of Tartu University in 1802, the foundation was laid for the systematic scientific research into Estonia’s geology.
Orviku and Viiding (1986) differentiated several stages, substages and periods in the history of geological research in Estonia. Within the first stage, which covered the time span from the beginning of human settlement in Estonia up to the reopening of Tartu University, they distinguished three substages: (1) gathering of elementary empiric knowledge about local geological monuments; (2) limited use of local mineral resources; (3) rapidly expanding use of local building materials after Estonia had been occupied by Germans, Danes and Swedes.
During the second stage, the most noteworthy events were the establishment of the Cabinet of Mineralogy at Tartu University in 1820, the foundation of the Tartu (later Estonian) Society of Naturalists in 1853, the University reform in 1892, the October Revolution in 1917, the birth of the independent Republic of Estonia in 1918, the incorporation of Estonia into the Soviet Union in 1940 and the restoration of the independent Republic of Estonia in 1991.
This chapter will deal only with the main outlines of the historical studies.
Organisation of geological education, research and exploration
After reopening in 1802, the old (1632) Tartu (Dorpat, Derpt, Jurjev) University developed into an outstanding centre of geological education in the former Russian Empire. Moritz von Engelhardt (1779–1842), who was born in Estonia but received special education in Germany (Leipzig, Freiberg), became the first professor of mineralogy at Tartu University in 1820 (Photo 6). Students came to Tartu beside Estonian and Livonian districts also from Poland, Lithuania and other countries. Lecturing in the Estonian language started in 1920. Hendrik Bekker (1891–1925), the first Estonian professor of geology, defended his Ph.D. degree at London University in 1921 (Photo 7).
During 1920–1940, only a few geologists were trained at Tartu University. A special governmental body - the Geological Committee (Survey) of Estonia was formed in 1937, but soon after Estonia was occupied by the Soviet Union in 1940, it ceased to exist. During the German occupation (1941–1944) it operated as the Department of Geology of the Institute of Industrial Research and afterwards, with the similar staff, as the Department of Mineral Resources of the Central Research Institute of Industry of the Estonian S.S.R. In 1947, the latter gave its staff to the new institution — the Institute of Geology of the Estonian Academy of Sciences (Jaanusson 1994). Up to the beginning of the fifties this institute carried out both geological survey and basic research. The small scientific staff of the geological department of Tartu University concentrated its efforts in some special fields of research.
Only a few Estonian geologists survived through World War II. Since the end of the war, some 350 geologists have graduated from Tartu University. Several tens of Estonian-born geologists were trained at the universities in Moscow, Leningrad (St. Petersburg), Irkutsk, Tomsk (all in Russia) and Vilnius (Lithuania).
In 1957, the Geological Survey of Estonia was founded and the Institute of Geology became mainly engaged in basic studies. Since then, the geological, hydrogeological, geophysical and other studies, initiated by branch establishments of central institutions of the former Soviet Union after World War II, were gradually turned over to local organisations. In the course of complex studies aimed at geological mapping, mineral prospecting and exploration, tens of thousands boreholes penetrating deep into the sedimentary cover and more than 500 boreholes (Fig. 3) passing right through the sedimentary cover to reach the crystalline basement were made. Large-scale prospecting and exploration studies on oil shale and phosphorite were performed.
The stratigraphic studies aimed at compiling regional legends for geological mapping have always been of high priority in Estonia. Based on the results of large-scale field and laboratory studies and drilling programmes, researchers from the Geological Survey of Estonia began compilation of basic maps separately for the Palaeozoic bedrock and the Quaternary sediments, on a scale of 1:200 000. These were completed in 1975. The mapping on a scale of 1:50 000 is under way. A large number of complementary maps dealing with the geomorphology, hydrogeology, engineering geology, distribution of mineral resources, geological structure and several other aspects, and provided with the relevant comments, have also been compiled.
During the late 1970s and the 1980s, as a result of joint efforts of scientists of Estonia, Latvia, Lithuania and the Kaliningrad District of the Russian Federation, a set of 15 general geological maps of the East Baltic (1:500 000) was compiled and published (Grigelis & Puura 1980).
Evolution of scientific ideas
At the dawn of biostratigraphical studies in Estonia, E. Eichwald (1795–1876), first and foremost in several papers from 1840, introduced the terms Cambrian, Silurian and Devonian (sensu Murchison 1839) in the area (Photo 8). In 1856–1858, Fr. Schmidt (1832–1908) proposed a reasonably detailed, palaeontologically and lithologically motivated subdivision of the northern Baltic Lower and Upper Silurian (now Cambrian, Ordovician and Silurian) rocks on stages level for the outcrop area, and an excellent geological sketch map.
In the last century it was already understood that the extraordinary survival of sediments, defined today as Vendian and Palaeozoic, in a nearly original low lithified level with perfectly preserved rich assemblages of fossils of calcitic, phosphatic and chitinous skeleton could be due to a very low tectonic compression and a small depth of burial which, according to recent estimation, never reached deeper than some 800 m during the whole geological history. In the middle of the last century, also the general feature of the regional geological structure — a homocline with a very gentle southward inclination of the weakly disturbed Palaeozoic strata overlying the crystalline (“granitic”) basement, was described by Fr. Schmidt (Photo 9), C. Grewingk (1819–1887) and several others.
Since the pioneer work of palaeontologists of several generations, tens of thousands specimens of fossils have been collected during systematically arranged field studies. The collections stored at the Institute of Geology and at the Geological Museum of Tartu University include about 1000 type-specimens of new species of fossils. A great number of new genera and families have been established. The most numerous groups of fossils, mostly Ordovician and Silurian, which have been studied in particular detail, are marine invertebrata: articulate and inarticulate brachiopods (studied by A. Rõõmusoks, M. Rubel, L. Hints, I. Puura during the last decades), molluscs of different classes, arthropoda like trilobites (Reet Männil), ostracods (L. Sarv, T. Meidla), merostomata and different echinodermata (Ralf Männil), tabulate (E. Klaamann) and rugose (D. Kaljo) corals and stromatoporoids (H. Nestor), bryozoans (Ralf Männil), graptolites (D. Kaljo), conodonts (V. Viira) etc. Besides, chitinozoans of problematic origin (Ralf Männil, J. Nõlvak, V. Nestor) as well as fossil elements and fragments of endo- and exoskeleton of early representatives of vertebrata from the Late Silurian and Devonian (E. Mark-Kurik, T. Märss) have also been studied.
Based on the data on the distribution of fossils in the Vendian – Middle Palaeozoic sedimentary record, a stratigraphic scheme for the Baltic States and the East-European Platform, as a whole (Ralf Männil, D. Kaljo, K. Mens, E. Mark-Kurik, H. Nestor, L. Hints a.o.) and a detailed biostratigraphic subdivision of the Ordovician and Silurian rocks in the North Baltic area were elaborated. The ecostratigraphic studies, taking into account peculiarities in the distribution of different fossils in lateral changes of lithologies, provided a basis for the correlation of the successions of different facial belts of the Early Palaeozoic pericratonic Baltic Basin (Reet Männnil, D. Kaljo a.o.). An impact of global and regional geologic/tectonic and palaeoenvironmental processes, like the continental drift of the Baltic Continental Plate from the south polar position to the northern tropics during the Late Vendian – Devonian, the continental glaciation in Gondvana in the Late Ordovician, causing ocean-level changes, and pericratonic tectonic activities along the continental palaeomargins with specific influence on sedimentation and fossil communities, has been fixed in the geological sedimentary record of the Baltic Palaeozoic marine basin.
Estonia was among the first regions where the theory of continental glaciation in the middle of the last century was applied (E. Eichwald, Fr. Schmidt) and later the structure and formation of different landforms and the evolution of the Baltic Sea were described in detail (Karl Orviku, H. Kessel, A. Raukas, E. Rähni a.o.).
Structural studies
Up till the 1950s, the area of Estonia was structurally classified as a simple, almost horizontal homocline. Detailed investigations of the sedimentary cover and the studies carried out within large-scale drilling and geophysical survey programmes revealed a typical fault-and-block tectonic pattern of both the sedimentary cover and the underlying crystalline basement (Vaher et al. 1962, Puura 1979). Classifications of the fault-related linear structures were suggested during the detailed structural studies of the oil shale (Puura 1986) and phosphorite basins (Puura 1987). The recent local crustal movements and weak earthquakes occurring against the background of the postglacial crustal uplift (Orviku 1960c) are, to a certain extent, related to faulting lines (Sildvee & Vaher 1995). The tectonic map with the accompanying explanatory text for the whole East Baltic territory was published by Suveizdis et al. (1979).
Research into the crystalline basement and the Vendian-Cambrian strata
The existence of the crystalline (“granite”) basement under the sedimentary cover was recognized in Estonia in the middle of the 19th century (Schmidt 1881), but systematic geological research into the basement was initiated not until the 1960s. The studies showed that in northern and north-eastern Estonia the basement consisted mainly of migmatized metamorphic rocks of the amphibolite facies and in southern and southwestern Estonia of granulite facies (Puura et al. 1976). The recent Sm-Nd dating of folded rock assemblages indicated Palaeoproterozoic age of the orogenic continental crust in southern Estonia (Puura & Huhma 1993) and in the whole Byelorussian – Baltic granulite province (Gorbatschev & Bogdanova 1993). Thus, the basement of Estonia was considered as a continuation to the Svecofennian Domain of the Fennoscandian Shield.
The first evidence about the Vendian rocks was obtained in 1842–1845 from a 90-m-deep borehole in Tallinn (Helmersen 1851). The first borehole, which reached the crystalline basement at a depth of 162.8 m, was drilled in 1898–1899 at Aseri. According to Sokolov (1953), the thickness of the Vendian complex was 92.5 m there (Rüger 1923). As a result of recent studies (Mens & Pirrus 1971, 1974, 1980, 1986), a great lateral and vertical facial variability has been established and a detail lithostratigraphical subdivision presented.
As early as the first half of the 19th century, the Cambrian sequence was divided into a lower, Blue Clay Unit and an upper, Sandstone Unit (Engelhardt 1820, Strangways 1822, Eichwald 1825). Following Murchison’s concept of the Silurian system, Eichwald (1840) included these two units into the Silurian, and Schmidt (1858), more specifically, into the Lower Silurian. Within the Cambrian sandstone upon the blue clay, Linnarsson (1873) suggested to distinguish units equivalent to the Swedish Euphyton Sandstone and Fucoid Sandstone units. These two terms were widely used in Estonia until denomination on the basis of geographical names (Öpik 1933).
A great step forward was A. Mickwitz’s discovery of malacofauna in the Eophyton Sandstone (Schmidt 1888) and the finds of an olenellid trilobite which proved Lower Cambrian age of these beds, and also the finds of the brachiopod Mickwitzia which indicated the same age with the correlatable beds in Sweden. Öpik (1925, 1926, 1929) contributed further important biostratigraphical data on the Lower Cambrian and revised the terminology (Öpik 1933). He (Öpik 1956) was convinced that in places the terrigenous beds dated as the basal Ordovician could be of Late Cambrian age. An increased access to numerous cores of borings all over Estonia enabled the subsurface Cambrian to be studied in more detail. Since the 1960s, many papers dealing with various aspects of the Lower Cambrian sequence, including a detailed lithostrati-graphic classification (Mens & Pirrus 1977), have been published. In southeastern Estonia, some parts of the terrigenous sequence were supposed to belong to the Upper Cambrian (Volkova et al. 1981). Intense biostratigraphical studies aimed at determining a convenient basis for the Ordovician System indicate that a comprehensive lower portion of the sandstone, previously regarded as basal Ordovician, is actually Late Cambrian in age.
History of Ordovician research
The study of the Ordovician strata in Estonia was commenced with the descriptions of the North-Estonian Klint (Severgin 1808, 1809). Engelhardt (1820), Strangways (1821, 1822) and Eichwald (1825) pointed out the similiarity of the North Estonian and Scandinavian sequences. Strangways (1822) produced the first geological-lithological map which included also Estonia.
In the history of Cambrian-Silurian stratigraphy of Estonia, the second half of the 19th century is known as the Schmidt’s Epoch (Männil 1986). Friedrich Schmidt (1832–1908), a descendant of Baltic Germans, published a comprehensive survey of the Cambro-Silurian outcrop area in northern Estonia (Schmidt 1858) which included the stratigraphic classification of the Ordovician sequence in the East Baltic area and a bedrock map. The units of Schmidt (Table 1) are neither strictly lithostratigraphical nor based on guide fossils, but reflect the stages of development of the regional benthic fauna without any preference for any particular group of fossils. In modern terms, they could be compared to topostratigraphical units. This approach, together with the use of geographical names for the units, was followed by subsequent generations of geologists in the research of the Estonian bedrock.
Based on a detailed study of the post-Tremadocian Lower Ordovician in Ingria, Lamansky (1905) introduced a stratigraphic classification. It was primarily based on ranges of trilobite species which he could follow into northern Estonia. An important result of his study was the recognition of numerous breaks in the correlatable North-Estonian sequence.
In the late 19th and early 20th century, a number of important taxonomic monographs were published by both native and foreign specialists (Rosen 1867, Schmidt 1874, Pahlen 1877, Dybowski 1877b, Holm 1886, Mickwitz 1896, Stolley 1896b, Koken 1897, 1925, Hoyningen-Huene 1899, Jaekel 1899, Bonnema 1909), and the data on the Ordovician of Estonia was included in several other publications as well. Of particular importance was the series of monographs on trilobites by Schmidt (1881, 1885, 1894, 1898, 1901, 1904, 1906, 1907). With Bassler’s (1911) monograph on bryozoans the former Lower Silurian Series was, as far as the Estonian sequence was concerned, definitely elevated to the rank of a system, termed Ordovician.
In 1914, the renowned American geologists P.E. Raymond and W.H. Twenhofel examined Cambro-Silurian exposures in Ingria, Estonia and Scandinavia with the purpose of attempting a correlation with eastern North America. Raymond (1916) proposed five new terms (Table 1) and divided the Ordovician System in the Baltoscandian Basin into three series: Lower, Middle and Upper Ordovician. According to current correlations, the boundaries of these series correspond very closely to those between the Ordovician series of North America. Bekker (1921) introduced the term Stage (Table 1) for the basic chronostratigraphic unit (= beds in Schmidt 1881) and revised the spelling of geographic names to accord modern maps (Bekker 1922, 1925). Armin Öpik (1898–1983), the successor of Bekker on the chair of geology at Tartu University (Photo 10), focused, as far as the Ordovician stratigraphy is concerned, on studying the Tremadocian sequence (Öpik 1929). However, his main contribution was in the field of palaeontology, documented by a number of monographs on brachiopods (Öpik 1930b, 1933, 1934, etc.), trilobites (Öpik 1937b), ostracodes (Öpik 1937a) and some other groups. In a manuscript in 1934, he proposed the terms Viru Series for the Middle Ordovician and Harju Series for the Upper Ordovician of the region, which were first published by Luha (1940a). Karl Orviku (1903–1981) made a noteworthy contribution to knowledge of the Ordovician in Estonia mainly in the field of detailed lithostratigraphy (Photo 11). His excellent monograph on the lower Middle Ordovician of northern Estonia (Orviku 1940), in which he distinguished the Lasnamägi and Uhaku stages (Table 1), exerted great influence on the succeeding generations of Estonian geologists. His analysis of discontinuity surfaces (hardgrounds), in particular, both with regard to their morphology and importance as markers of stratigraphic breaks, received wide attention. In subsequent papers Orviku (Orviku 1960a, b) gave a detailed lithostratigraphic classification of the Lower Ordovician Volkhov and Kunda stages of northern Estonia.
The scientific activity of the succeeding generation (Jaanusson, Männil, Rõõmusoks, Kaljo, Aaloe and several others) started in an organisation known as the Section of Geology of Gustavus Adolphus Natural Science Circle. As a result of field studies and later publications, the Schmidt’s hitherto poorly known Lyckholm beds (Table 1) were divided into three separate stages (Jaanusson 1944b, 1956) and the stratigraphy of the Viru Series (Middle Ordovician) of northern Estonia was revised (Jaanusson 1945). In the upper Middle Ordovician, the occurrence of K-bentonites was observed (Jaanusson & Martna 1948), and their importance for correlation and stratigraphic classification recognized (Jürgenson 1958a, Bergström et al. 1995).
The progress in knowledge of the Ordovician sequence in Estonia after the war was largely due to the availability of numerous cores of drill borings. Ralf Männil (1924–1990) solved the somewhat confused stratigraphic terminology of the Schmidt’s unit F1a by introducing the term Nabala Stage (Männil 1958b). He also showed (Männil 1958c, d, 1960) that information from borings indicated Schmidt’s Vasalemma beds to represent a lithostratigraphic unit which also includes the beds of Keila Age. For the chronostratigraphic unit above the Keila Stage Männil proposed the term Oandu Stage, borrowed from the poorly exposed lithostratigraphic unit Oandu beds in northeastern Estonia (Öpik 1934) which was unknown to Schmidt. Kaljo, Rõõmusoks and Männil (Kaljo et al. 1958) emphasized the need of regional terms for Ordovician series and proposed the Oeland Series for the Lower Ordovician in the accustomed Baltoscandian usage. The post-Tremadocian Lower Ordovician was renamed Ontika Subseries.
As a result of a thorough examination of the subsurface Ordovician in Estonia, Latvia and Lithuania, Ralf Männil (Männil 1966, preliminary report 1964) summarized the development of the Baltoscandian Basin in the Ordovician Period. He showed that in an extensive west-central area of the East Baltic region (Livonian Tongue after Jaanusson 1976), including southern Estonia, the post-Tremadocian Ordovician rocks are basically of the same type, both lithologically and faunistically, as on the Swedish mainland north of Scania, and that they belong to a single Swedish-Latvian Facies Zone (Central Baltoscandian Confacies, Jaanusson 1976). The Estonian Facies Zone (North Estonian Confacies Belt, Jaanusson 1976) was proved to be conspicuously different in many respects. Põlma (1967, 1972, 1982) presented a comparative analysis of the carbonate lithology of these main regions and distinguished a lithologically transitional belt, characterised by transitional lithologies and interlocking pattern of various lithofacies from the south and north. The relation of the boundaries of the transitional belt to those of confacies belts is still disputed (Jaanusson 1995, Meidla 1996). Of great palaeogeographical interest was the discovery that the Varangu beds (Männil 1958a), which occur in a limited area of northern Estonia, correlate with the upper Tremadocian Ceratopyge Shale in Scandinavia (Viira et al. 1970, Kaljo & Kivimägi 1970). The stratigraphy of the Middle Ordovician (Viru) of North Estonia (stratotype area) was summarised in a comprehensive monograph by Rõõmusoks (Rõõmusoks 1970), and that of the Ordovician of Estonia in general, in a chapter of a separate book by the same author (Rõõmusoks 1983).
The rich knowledge of the Ordovician fauna of Estonia has increased the precision of correlations not only within Estonia but also with the Ordovician sequences elsewhere (Nõlvak & Grahn 1993, Männil & Meidla 1994, Jaanusson 1995).
History of Silurian research
M. von Engelhardt (Engelhardt & Ulprecht 1830) was the first to draw attention to the differences between the rocks now classified as Ordovician and Silurian. He noted that on mainland Estonia limestones containing orthoceratite cephalopods and trilobites were succeeded southwards by younger rocks with corals and pentamerid brachiopods.
In the monograph by Murchison, Verneuil and Keyserling (1845) the Silurian sequence of Estonia is briefly but adequately summarised. Murchison distinguished within the Silurian of Estonia (in the current sense) the following three units in ascending order (1) Pentamerus Limestone, (2) Limestone with corals, and (3) Limestone with Terebratulas.
Schrenk (1854) gave the first comprehensive lithological survey of the Silurian localities in Estonia. Schmidt (1858) continued the study in a much greater detail and with emphasis on fossil fauna. His classification of the Silurian sequence of Estonia into three groups is roughly comparable to Murchison’s tripartite subdivision and the distinguished groups correspond to the Llandoverian, Wenlockian and Ludlovian + Downtonian Series in the current sense (Table 2). The lower, Schmidt’s Group with Smooth Pentamerids (= Llandoverian), was subdivided into smaller units and the base of the Upper Silurian defined, in the outcrop area of northern Estonia, at the level of the present Ordovician-Silurian boundary in the area. Schmidt interpreted the Borealis Banks as local mass accumulation of shells (Muschelbänke) and regarded the unit to be somewhat artificial in a chronostratigraphic sense. Subsequently, Schmidt ( 1879, 1881) introduced capital Latin letters combined with Arabic numbers (G1, G2, G3, etc.) as symbols for this unit. He (Schmidt 1892) further refined his stratigraphic classification by distinguishing within the Upper Oesel Group Eurypterus beds and Ilionia (Didyma) beds which correspond to the Rootsiküla and Paadla stages of the current classification, respectively.
The Silurian fauna of Estonia received early attention, especially by Eichwald, in various papers. A contribution of great international importance was the monograph on Silurian (=Ordovician + Silurian) fishes by Pander (1856). The Silurian (s.str.) material was mostly derived from the Schmidt’s Upper Oesel beds of Saaremaa. The description of Pander’s new group Conodonta was partly based on specimens from Saaremaa.
The exquisitively preserved specimens of merostomes, fishes and some other uncommon groups from exposures at Viita in Rootsiküla on Saaremaa Island were of wide international interest. The exposures were located by Schrenk in 1852 and material from those beds formed the subject of many papers. The monographs by Rohon (1892, 1893) and Holm (1898) should be mentioned in the first place.
The Schmidt’s stratigraphic terminology was somewhat inconsistent which made Twenhofel (1916) to propose more adequate terms for some units (Table 2). Bekker (1922) revised the terminology to accord modern maps and started field work with an aim of studying the stratigraphy of the Silurian sequence of Saaremaa in more detail (Table 2). On account of his untimely death, only an outline of his stratigraphical results became published (Bekker 1925). The study was continued by Artur Luha (1892–1953) who assisted Bekker in the field. Regrettably, only a condensed version (Luha 1930) of his voluminous manuscript was published.
In 1929, Luha discovered a locality with a rich agnathan fauna at Himmiste-Kuigu on Saaremaa. Later studies (Aaloe 1963b) showed that the Himmiste beds were younger than previously believed (about middle of the Paadla Stage).
Luha (1933) improved the chronostratigraphic classification by restricting the term Jaani Stage to beds now known to correspond to the Lower Wenlockian (Table 2). Later Luha (1946) introduced the term Jaagarahu Stage for the unit comparable to the Upper Wenlockian. Teichert’s (1928) study on the Lower Llandoverian in the western part of mainland Estonia and on the Island of Hiiumaa is also worthy of mentioning.
Studies of the Silurian resumed in the mid-1950s. Numerous borings made it possible to extend examination of lithofacial relationship also south of the outcrop area, and facilitated establishing of a detailed lithostratigraphic classification (Aaloe 1958a, 1960, 1961, Aaloe & Kaljo 1962, Einasto 1962, etc.). Kaljo and Sarv (1966) specified equivalents to the Downtonian Series in the Estonian sequence. Examination of graptolites (Kaljo 1967, Kaljo & Vingisaar 1969) contributed to a more precise correlation of the sequence with other areas. As the Kaugatuma Stage proved to be a composite unit, Klaamann (1970a) distinguished its lower, Ludlovian part as a separate Kuressaare Stage. The discovery of Llandoverian K-bentonite beds in borings on Saaremaa and in the southwestern part of mainland Estonia by Jürgenson (1958a, 1964) contributed to correlations. Studies of lithology of the Silurian rocks received increased attention (Jürgenson 1966, 1974, etc.).
Biostratigraphic correlations were greatly facilitated by monographic descriptions of several groups of microfossils such as ostracodes (Sarv 1968), agnathan scales (Märss 1996), chitinozoans (V. Nestor 1994) and conodonts (Männik 1992b).
Silurian macrofauna was described in numerous papers, some of those monographic in character. Various groups were covered, including stromatoporoids (Nestor 1964, 1966), tabulate corals (Sokolov 1951b, 1952b, Klaamann 1961, 1962, 1964, 1966, 1970b), bryozoans (Astrova 1970), articulate brachiopods (Rubel 1970a, b), and trilobites (Männil 1982, 1992).
The impressive array of information from the Silurian of Estonia is systematised and discussed in several books, covering general features (Kaljo 1970c), facies and fauna (Kaljo 1977), communities and biozones (Kaljo & Klaamann 1982), ecology (Kaljo & Klaamann 1982) and lithology (Jürgenson 1988).
History of Devonian research
Engelhardt and Ulprecht (1830) provided the earliest information of the red sandstones and mentioned also finds of vertebrate teeth and bone fragments. Kutorga (1835, 1837) described the exposures at Tartu (Fig. 4). Eichwald (1840a, c) and Buch (1840) recognised very early that the terrigenous sequence was comparable to the Devonian Old Red rocks in Great Britain. Helmersen was the first to show the approximate outcrop area on a map. A similar map was provided by Murchison et al. (1845). The available information was summarised by Grewingk (1861, further specifications in 1878 and 1879, Photo 12).
Quenstedt (1838) suggested that the teeth and bones recovered from the sandstone belong to ancient fish-like forms. The fairly substantial material of fish remains was subject to monographic studies by Asmuss (1856) and Pander (1856, 1860).
Based on the field studies of the Upper Devonian carbonate rocks of the southeastern part of the Republic, Bekker (1924a) attempted a chronostratigraphic classification of the sequence of which the term Dubniki Stage is still in use. In the current stratigraphic practice the other terms are based on the sections in the neighbouring areas of Latvia and Russia where carbonate rocks of this age have a wider distribution.
The lithostratigraphy of the basal Old Red lithofacies in the outcrop area was subject to a detailed study by Orviku (1930c, 1932, 1935b). Obruchev (1933) distinguished two stratigraphic units, now regarded as the Pärnu and Narva stages. Subsequently, Orviku (1948) published an additional comprehensive paper on the stratigraphy of the Narva River Stage (=Narva Stage).
An important contribution to the knowledge of the Old Red sequence was the biostratigraphic classification, based on fishes from both Estonia and Latvia. The classification was developed by Gross (1933, 1940a, b, 1942, 1951) and later continued by E. Kurik (née Mark). Mark (1958) distinguished the Eifelian Aruküla Stage. The historically known Aruküla caves near Tartu (Photo 13) where, since 1831 during more than 20 years, H. Asmuss had excavated bones of large placoderms and other fishes, were selected for the type locality.
During several years between World War I and II, V. Paul (1934, 1939) excavated fossil fishes at Tamme near Lake Võrtsjärv and at Haaslava. The excavations, organized by E. Mark-Kurik, started in 1949 and lasted with a few intervals up to 1993 (more than 10 excavations were made at Karksi alone). Taphonomical study was provided during these excavations.
The overlying, lower Givetian unit was distinguished by Mark (1958) as the Burtnieki Stage with the stratotype in Latvia. The stratigraphy of the Devonian in Estonia was summarised by Mark and Paasikivi (1960), including new information obtained from examination of cores from numerous deep borings. Special attention was further paid on studying the mineral composition of terrigenous rocks, largely on the basis of core material. The studies of the kind were initiated by H. Viiding (1929–1988) and supplemented particularly by A. Kleesment (née Tamme). Numerous papers were published on the subject (Viiding 1962, 1964, 1965, 1976b; Tamme 1962, 1964; Mark & Tamme 1964). Based on lithological and mineralogical criteria, Kleesment (1966) and Kleesment et al. (1975) recognised in the borings of southeastern Estonia the presence of the Lower Devonian Tilžė and basal Middle Devonian Rēzekne stages, previously distinguished in Latvia and Lithuania. Mark-Kurik (1991a) has considered the Rēzekne Stage as a lower Devonian unit.
The Devonian stratigraphy in Estonia was subsequently summarised in the papers published by various authors in the book “Devonian and Carboniferous of the Baltic” (Sorokin 1981). Recently, Kleesment (1994, 1995) published important contributions to detailed lithostratigraphy of the Middle Devonian sequence in Estonia. Mark-Kurik correlated Estonian Devonian units with those in other areas, e.g. in Scotland and Latvia (Mark-Kurik 1981, 1991a, b, Kurik et al. 1989). Valiukevičius (1994) gave an acanthodian zonation for the Baltic Basin (including Estonia). The fossil fishes of the Devonian sequence continued to receive attention. Beside the above-mentioned papers by Gross, the contributions by Heintz (1930, 1934), Mark (1953a, b, 1956, 1963), Mark-Kurik (1968, 1973, 1993a, b, Mark-Kurik et al. 1991), Obruchev and Mark-Kurik (1965, 1968) deserve special mention. Mark-Kurik has also studied special problems of palaeoichtyology: palaeopatology (Mark-Kurik 1966), functional morphology (1984) and trophic relations (1995). Rich collections in Estonia have provided an excellent basis for studying of several fossil fish groups: antiarchs by Karatajūtė-Talimaa (1960, 1963), arthrodires by Obrucheva (1962, 1966), acanthodians by Valiukevičius (1985) and crossopterygians by Paul (1940), Vorobyeva (1977). Thomson (1940) pioneered in studying the Devonian plant remains (Photo 14). Later Vaitiekūnienė (Kleesment et al. 1975) and Kalamees (1988) studied spores and macroremains (including phytoleimma), respectively. A number of papers contain descriptions of the Devonian invertebrates from Estonia: ostracodes (Öpik 1935a, Polenova 1966), lingulates (Batrukova 1960, 1964, 1969, Gravitis 1981) and conchostracans (Mironova 1969).
History of Quaternary research
The first stratigraphical scheme of the Quaternary deposits was compiled by Schrenk (1854). Based on the then prevailing drift theory he, like Schmidt (1854) and Grewingk (1861), divided all Quaternary sediments into diluvial and alluvial sediments with several lithological varieties.
Estonia was among the first regions where the theory of continental glaciation was applied. Eichwald (1853) was the first in the Baltic provinces to consider the possibility that at least northern Estonia had once been covered by an ancient active glacier. Already in 1865, Schmidt (1865) clearly spoke about glacial sediments, and a bit later he (Schmidt 1869) described glacial and postglacial formations and differentiated four stages in the development of the territory, including the time of the invasion of big glaciers, the time of the melting of the glaciers, and the time of the final melting of the ice with a wide distribution of fresh-water lakes. Schmidt (1858) was the first who found the shells of Ancylus fluviatilis in beach deposits on Saaremaa Island and distinguished a fresh-water stage in the Holocene history of the Baltic Sea.
Helmersen (1869) explained the distribution of erratic boulders and formation of boulder clays and ice scratches with the joint action of the continental ice sheet, floating icebergs and erosional processes on land. Schmidt (1871) proved that in the glacial epoch a unitary glacier moved from Scandinavia over the Baltic Sea depression and Estonia. In 1879, Grewingk already spoke about several glaciations (Grewingk 1879), basing on the study of the different till beds in Tartu.
However, it was not until half a century later that Grewingk’s statement was confirmed by palaeontological data. In 1939, Orviku performed the first detailed studies on interglacial organogenous deposits at Rõngu (Orviku 1939) which, according to pollen zones (Thomson 1939a, 1941), were correlated with the typical Riss-Würmian (Eemian) interglacial deposits in Western Europe. By now, tills of five glaciations or big stadials have been identified in Estonia (Raukas 1978) and both Eemian and Holsteinian interglacial sediments described in detail (Liivrand 1991). Official stratigraphical schemes of the Quaternary (Raukas & Kajak 1995), Late-glacial (Pirrus & Raukas 1996) and Holocene (Raukas et al. 1995b) have been accepted and published.
Monographic studies of ice-marginal formations (Raukas et al. 1971), bedrock topography (Tavast & Raukas 1982), lithology of Quaternary deposits (Raukas 1978), modern (Orviku Kaarel 1974) and ancient (Kessel & Raukas 1967) coastal formations have been published. The overviews about the glacial history (Raukas 1995a) and the history of the development of geomorphology in Estonia (Raukas & Karukäpp 1993), imparting more information about the study history of different landforms and types of sediments, appeared recently in print.
The authors of the chapter would like to thank Prof. Valdar Jaanusson for his numerous valuable comments.
V. Puura, V. Klein, H. Koppelmaa & M. Niin
The early Precambrian crystalline rocks are covered by the Upper Vendian and Palaeozoic sedimentary rocks. The basic data for studies of the Precambrian has been obtained by means of boreholes and geophysical survey. Of about 500 boreholes passing right through the sedimentary cover, the deepest ones penetrate into the basement to a depth of up to 450 m.
The Precambrian basement in Estonia consists of two megaunits: the orogenic Svecofennian complex of metamorphic and plutonic rocks and the anorogenic complex of plutonic rapakivi granites and related rocks. Earlier views about the age of basement rocks (Puura et al. 1976, 1983) have considerably changed during the last decade due to the isotopic dating of the South-Estonian granulitic crust (Puura & Huhma 1993) and rapakivi granites of Fennoscandia (Rämö et al. 1996). The new standpoints, based on the results of these studies, underlie the recent joint publications on the Precambrian of the Gulf of Finland and surrounding area (Koistinen 1994, 1996; Laitakari et al. 1996). Based on recent results, a stratigraphic chart of Precambrian rocks of Estonia was compiled (Table 3).
According to the degree of metamorphism and the composition of the metamorphic sequences reflected in the geophysical patterns, Estonia’s basement is divided into structural regions (Figs. 5, 6) which differ from each other in the volume of sedimentary and felsic to mafic volcanic rocks.
Comparison of the basement-forming rocks in Estonia, Finland and Sweden has shown that the metamorphosed volcanic and sedimentary rocks in Estonia’s basement have many features in common with the rocks in the Svecofennian orogenic complexes. In the first instance, it was established that the rocks of the Tallinn and Alutaguse structural zones and the Svecofennian complex are similar in lithology and have the same stage of metamorphism (Puura et al. 1983, Klein 1986). The orogenic supracrustal rocks of southern and western Estonia differ from the bulk of Svecofennian rocks of Scandinavia by their prevailingly mafic to intermediate composition and high‑grade metamorphism. This is in good correlation with high anomalies of gravity and magnetic fields (Puura et al. 1976, 1983, Koppelmaa et al. 1978). The Palaeoproterozoic age of the granulite complex of Estonia was dated by Sm‑Nb isotopic studies (Puura & Huhma 1993). Petrological signatures of mafic rocks in southern and western Estonia are concordant with those in the northern and northeastern parts of the territory.
In the Alutaguse Zone (Fig. 5) gneisses containing biotite, cordierite, garnet and sillimanite intercalate with biotite gneisses and form a complex of the same name (Puura et al. 1976). Within the Alutaguse Zone, in the area of Uljaste, Haljala and Assamalla, the basement comprises sulphidic black schists, quartzites, amphibole and pyroxene gneisses, marbles and pyroxene skarns.The rocks in the Alutaguse Zone derive from clastic successions with minor sequences of volcanic, sandy and carbonate rocks in the above-mentioned areas. The local uplifts of the basement in the Uljaste and Assamalla area consist primarily of quartzites.
The Al-rich, sillimanite‑garnet‑cordierite gneisses are, for the most part, medium‑grained, banded and migmatized by plagioclase‑microcline granite or pegmatite. The mineral composition of gneisses varies. Light minerals are represented by quartz, plagioclase (An 25–55) and microcline; dark minerals by biotite, cordierite, garnet and sillimanite. Muscovite is rare, while andalusite is occasional and rare. Chemically, the sillimanite‑ garnet‑ cordierite gneisses are similar to pelites and originate, in all likelihood, from psammitic to pelitic sediments. Microgneisses, rich in quartz, form interlayers with the highest sand content in these sediments.
The Tallinn Zone is characterised by the stratified Jägala Complex of intercalating sillimanite‑cordierite and biotite gneisses, intermediate to mafic metavolcanics, and leucocrate gneisses. In the WNW‑ESE‑trending zones, the primarily psammitic to pelitic metasediments intercalate with metavolcanites. The acidic metavolcanites alternate with more abundant intermediate to mafic metavolcanites.
The intermediate metavolcanic rocks are represented by metamorphosed to biotite‑hornblende and biotite gneisses which are fine‑ to medium‑grained and migmatized by microcline‑ plagioclase granites. The main minerals of biotite gneisses are plagioclase (An 30–45), quartz, biotite and, in places, microcline. The basic minerals of biotite‑hornblende gneisses, which are slightly more mafic in composition, are plagioclase (An 35–50), quartz, hornblende and biotite. These gneisses are andesitic (SiO2 55–63%, Na2O + K2 = 4.5–6.5%).
Typical felsic quartz‑feldspar gneisses are fine‑grained, rather massive or schistose granoblastic rocks. In single boreholes, rocks with relicts of blasto‑porphyritic texture (phenocrysts of quartz and plagioclase), indicative of their volcanic origin, have been found. Quartz (25–40%), plagioclase (An 20–40) and potassium feldspar form 85‑95% of quartz‑feldspar or granite gneisses. In the chemical composition (SiO2 65–76%, Na2O + K2O = 5.5–8%) the quartz‑feldspar and granite gneisses are similar to acidic volcanites (dacites, rhyolites) and arcosic sandstones.
Petrographically, the Al-rich gneisses of the Jägala Complex are similar to those in the Alutaguse Zone.
The basement in the West-Estonian Zone consists predominantly of the same assemblage of rocks as in the Tallinn Zone, although the stratified structure is not so well reflected in geophysical anomalies. The rocks, characteristic of this zone, are rather uniformly medium‑ and fine‑grained biotite‑hornblende gneisses and amphibolites, which have been migmatized by microcline granites. The amphibolites mostly occur as layered bodies and are intercalated with gneisses. In the amphibolites, plagioclase (An 35–55) and hornblende are the main minerals, however, they may also contain biotite, clinopyroxene and quartz. The main minerals of the biotite‑hornblende gneisses are plagioclase (An 30–50), quartz, hornblende and biotite, rarely potassium feldspar. There are also gneisses, the mafic parts of which consist entirely of biotite. North of Haapasalu, on the Noarootsi Peninsula, the gneisses also contain hypersthene. According to the chemical composition, the amphibolites (SiO2 45–53%, Na2O + K 2O = 3–4.5%) are referred to basalt, and the gneisses (SiO2 55–63, Na2O + K2O = 5–6.5%) to andesite.
In the Tapa Zone, a rock association, analogous to that of the West-Estonian Zone (amphibolites, biotite‑hornblende gneisses, in places pyroxene gneisses) occurs. The Jõhvi Zone (Fig. 5) is composed of the rocks of the Vaivara Complex. Magnetite quartzites occur in a limited area together with Al-rich and pyroxene gneisses. The latter contain interlayers of quartz‑feldspar and biotite‑amphibole gneisses. The fine‑ and medium‑grained pyroxene gneisses with a variable mineral composition display charnockitic and granitic migmatization. Orthopyroxene and biotite are always present. The content of clinopyroxene and hornblende varies from 0 to 25%. Of light minerals, plagioclase (An 40–55%) predominates, while quartz and potassium feldspar are often absent. Within the Jõhvi magnetic anomaly area, rocks of almost ultramafic composition comprising orthopyroxene, clinopyroxene, hornblende, biotite and plagioclase are occasionally encountered (5–10%). Biotite‑hypersthene gneisses containing plagioclase and quartz are also widespread. Chemically, the gneisses correspond to andesite.
The magnetite quartzites, fine-grained banded rocks in the Jõhvi area, contain besides quartz and magnetite, also garnet, orthopyroxene, clinopyroxene, hornblende, cummingtonite and biotite in different quantities. The average content (by microsections) of quartz is 30–40%, with the proportion of magnetite reaching 25–30%. The magnetite quartzites are cut by veins of pegmatoid microcline granite.
In the South-Estonian Zone, the metamorphic complex consists of hypersthene, clinopyroxene and amphibole gneisses, originating from mafic to intermediate volcanites, and possibly from greywackes. It also contains Al‑rich and minor members of felsic gneisses.
Different fine‑ to medium‑grained pyroxene gneisses, which have undergone charnockitic and granitic migmatization, are characteristic of southern Estonia. The primary structures of these gneisses have been obscured or obliterated. The characteristic mineral assemblage of the hornblende‑pyroxene gneisses is orthopyroxene + clinopyroxene + hornblende + biotite + plagioclase +/‑ potassium feldspar +/‑ quartz. The plagioclase is mostly antiperthitic, mainly andesine‑labradorite, rarely bytownite. The potassium feldspar is orthoclase‑microperthite. Quartz is rare. The hornblende‑pyroxene gneisses have been found mostly in boreholes in the surroundings of Pärnu and Viljandi where they occur as interlayers in acidic gneisses. The chemical composition of the amphibole‑pyroxene gneisses corresponds to basalt or basaltic andesite (SiO2 47–54%, Na2O+K20 = 3–5%), but the content of iron, magnesium and calcium differs noticeably. The increased content of magnetite (3–4%) is a specific feature of these gneisses. The essential minerals of the biotite‑hypersthene gneisses are plagioclase (mainly An 35–45%, in some cases An 70–80%), hypersthene, biotite and quite often quartz and potassium feldspar. Gneisses of this type occur typically in the vicinity of Tartu, Otepää and Laeva. Among the biotite‑hypersthene gneisses, both melanocratic and leucocratic varieties occur (SiO2 48–60%). Compared to the hornblende‑pyroxene gneisses, the biotite‑hypersthene gneisses are generally poorer in calcium, but richer in potassium and magnesium, which is evidently due to the weathering of the source rock and mixing with pelitic matter. In the rather rare quartz‑feldspar gneisses of southern Estonia, garnet or hypersthene and hornblende occur as accessory minerals.
The gneisses, formed at granulite facies in the South-Estonian and Jõhvi zones, contain hypersthene and accessory spinel, garnet and cordierite porphyroblasts (also in granitic veins) and sillimanite, the latter occurring as inclusions in cordierite. The biotite gneisses occur together with sillimanite‑cordierite gneisses. They are medium‑ to fine‑grained, often foliated migmatitic rocks, the main minerals being quartz, andesine, biotite and potassium feldspar. Garnet, cordierite, sillimanite and muscovite occur in small quantities. The content of dark minerals averages 20–25%. Compared with other gneisses, biotite gneisses have the highest content of quartz.
Traditionally, the granitoid rocks of southern Finland have been classified into four groups, based on their relationship to orogenic movements (Koistinen 1994, 1996). These groups are synorogenic (synkinematic), late‑orogenic (late‑kinematic), post-orogenic and anorogenic rocks (rapakivi granites). Practically, this classification is often expanded on the whole variety of plutonic rocks.
Compared to Finland, the Estonian basement is rather poor in orogenic plutonic rocks. Synorogenic granitoid complex and associated mafics are rare in Estonia. The late‑orogenic potassium granites which form the extensive W‑E‑striking belt in southern Finland, occur as small bodies and migmatite veins in the basement of northern Estonia.
In Finland, the granite migmatites fall into two distinct age groups which are related to early (1.9–1.87 Ga) and late orogenic (1.84–1.83 Ga, Koistinen 1996) granitoids. Like in southern Finland, where the both age groups of migmatites occur in the same metamorphic complex around the Potassium Granite Belt (Koistinen 1996), age classification of migmatites in Estonia’s basement is extremely complicated and, therefore, they are treated as one orogenic group. Occasionally, the classification of other plutonic rocks into the early and late orogenic groups is possible. Small bodies of synorogenic gabbronorite and gabbro, or metagabbro, cut by granite veins, occur in the Tapa Zone. The mafic rocks contain abundant hornblende and biotite of later origin. Similar rocks in northeastern Estonia form the considerably large Pada Pluton, which contains also diorite.
Drilling in southern and western Estonia has revealed some mafic, probably synorogenic rocks. There are small gabbro-norite plutons including Võru, Laeva, Pärnu, Vanaküla and several others, some of those with structural orientation. Gabbro-norite is a massive rock of coarse or medium grain size, which contains plagioclase (An 50–70), ortho- and clinopyroxene, hornblende and biotite of secondary origin. In northern Estonia, small granodiorite and quartz-diorite bodies have been found in single boreholes (Aruküla, Letipea). The quartz-diorite is orientated, medium‑grained and cut by veins of microcline granite. The small Utria body in the northeastern coastal area consists of massive medium‑grained gabbrodiorite.
The granite rocks of the Estonian crystalline basement are mainly migmatite granites: plagioclase‑microcline granites in northern and western Estonia, and charnockites and plagioclase‑orthoclase granites in the granulite facies area of southern Estonia. The charnockites consist of potassium feldspar, plagioclase (andesine) and, to a lesser extent, of biotite, hypersthene and hornblende, the latter three forming 5‑10 % of the rock. Quartz, microcline, oligoclase‑andesine and biotite are the main minerals in the migmatite granites of northern and western Estonia.
In southern Finland, the 1.82–1.78 Ga tonalitic to monzonitic and granitic post‑orogenic intrusions are neither voluminous nor numerous (Koistinen 1996). However, they mark the final stage of the Svecofennian orogeny when the temperature of the crust was still high. Recently, a group of post‑orogenic granites was identified in the Estonian basement as well (M. Niin, unpublished report).
The Taadikvere body in Central Estonia consists of granodioritic – quartz-monzonitic rocks which are of preferred orientation and contain plagioclase and potassium feldspar phenocrysts. The medium-grained groundmass of the rock comprises quartz, plagioclase (An 32–36), potassium feldspar, biotite and hornblende.
The Virtsu body in western Estonia, consists of porphyritic rocks of quartz monzonite composition that are strongly crushed within the west-east oriented central Estonian cataclastic zone. The medium-grained groundmass of the rock consists of plagioclase (An 31–42), potassium feldspar, quartz and biotite with some admixture of hornblende, and numerous potassium feldspar and plagioclase phenocrysts, up to 2–3 cm in diameter.
The mineral paragenesis of metamorphic rocks and, partly, of early orogenic plutonic rocks is due to regional metamorphism which in the Svecofennian orogen in Finland occurred in several stages during 1.885–1.81 Ga as dated by isotopic studies (Koistinen 1996).
Metamorphic zoning (Fig. 7) is typical of the Svecofennian orogen. In the province as a whole, andalusite‑muscovite mica schists prograde into potassium feldspar‑sillimanite gneisses and migmatitic garnet‑cordierite gneisses. The neosomes in the migmatites differ markedly in type. In central Finland the granitoid area is surrounded by tonalitic and trondhjemitic migmatites, while potassium granitic neosomes occur in migmatites of southern Finland (Koistinen 1996). Potassium granitic migmatites extend into the North-Estonian zones of amphibolite metamorphism. In the southern part of Estonia, charnockite and enderbite migmatites characteristic for granulite facies area have been described.
The Precambrian basement of Estonia consists of rocks which have been subjected to high-grade metamorphism. In northern Estonia (Figs. 8, 9), amphibolite facies gneisses are most abundant, while granulite facies mineral assemblages occur locally as in the Jõhvi and Tapa zones. Assemblages marking transition from amphibolite to granulite facies occur in the vicinity of Uljaste and Assamalla (Klein 1986). The amphibolite facies gneisses in Estonia serve as an extension to those spread in southern Finland. The dominant metamorphic grade here is a high- temperature amphibolite facies with a local PT zoning from sillimanite‑potassium feldspar subfacies to granulite facies (Uljaste, Haljala). Geothermobarometry, mostly of the biotite + garnet + /‑ sillimanite assemblage and of cordierite, estimates prograde metamorphism at 600–700º C and 3‑5 kbar.
In the Jõhvi Zone, there are characteristic granulite mineral assamblages in cordierite‑garnet gneisses (hypersthene) and in mafic gneisses (two pyroxenes and spinel). In the Tapa Zone, the traces of granulite metamorphism have probably been partly removed by high-temperature retrograde metamorphism.
In southern Estonia, granulite facies gneisses form a large domain (Fig. 7), which extends from the Middle-Estonian fault zone (Saaremaa‑Peipsi Zone) to northern Latvia and further south through the beltiform Belarussian‑Baltic granulite domain. The conditions of metamorphism have been mainly studied from two drillcores — Kõnnu 300 and Varbla 502 (Koppelmaa et al. 1978, Hölttä & Klein 1991). The mineralogy of granulites varies. The widespread garnet and cordierite, formed by breakdown of biotite and sillimanite, indicate prograde metamorphism. Hypersthene coexists with garnet and cordierite although, so far, the sillimanite‑hypersthene assemblage has not been found. The PT‑conditions have been calculated using several geother-mometers and geobarometers (Hölttä & Klein 1991). These give temperature estimates for the prograde stage of metamorphism of 700–800ºC and pressure estimates of 5–6 kbar or more.
The age relations between the described zones are still ambiguous. The southern Estonian granulite area has been correlated with those in Finland. In the Haukivesi‑Kiuruvesi Complex metamorphism has been dated at 1.88 Ga (Korsman et al. 1984). High-temperature metamorphism in southern Finland, 1.83–1.81 Ga ago (Korsman et al. 1984), may be correlated with similar metamorphism in northern Estonia.
It is emphasized that the southern Estonian granulites (Fig. 10) formed under higher pressure than is characteristic of Svecofennian metamorphism. This suggests that the South-Estonian region represents a deeper crustal section (Koistinen 1996).
A large time span for the Fennoscandian rapakivi and related rocks’ plutonism at 1.65–1.54 Ga was established by isotopic studies (Rämö et al. 1996). Recently, it was stated that the province consists of four subprovinces separated areally and differing in age (Puura & Flodén 1996). The large plutons have a central position in the subprovinces, while the stocks and dike swarms occur in peripheries of the subprovinces. Volcanic rocks have survived as remnants in the vicinity of the main plutons. In the basement of Estonia, plutonic and related rocks of two, Vyborg and Riga‑Åland subprovinces occur (Table 4).
The Vyborg Pluton in southeastern Finland and adjacent offshore area has a central position in the oldest rapakivi subprovince. Its southern satellite spreads in the bottom of the Gulf of Finland, near the northeast coast of Estonia (Koistinen 1994).
Characteristic structures of this subprovince in Estonia are stocks of porphyritic potassium granites, chemically only little differing from the proper rapakivi (Kuuspalu 1975, Puura & Flodén 1996). Presuming that also the westernmost and smallest known but undated Taebla Stock (Fig. 5) belongs to the Vyborg Subprovince, then practically the whole mainland Estonia was influenced by rapakivi magmatism at about 1.65–1.64 Ga.
The Naissaar Pluton (55km × 25 km), the northern part of which extends under the Gulf of Finland, is composed of porphyritic granites cut by aplites. According to the chemical and mineral composition, this pluton is divided into two phases (Soesoo & Niin 1992). The more melanocratic granites of the first phase form the periphery of the pluton. The central part of the pluton is composed of leucocratic granites of the second phase, which have some similarities with the second and third phases of the Märjamaa Stock. The structure of the rocks is massive, in places (second phase) trachytoidic.
There are two generations of quartz (25–35%): crystals within microcline phenocrysts and anhedral crystals in the groundmass. In places, tabular microcline (35–45%) contains inclusions of plagioclase. Euhedral to subhedral zoned plagioclase (20–25%) is of oligoclase composition. Biotite forms small flakes containing euhedral crystals of zircon and apatite as inclusions. Muscovite and fluorite of post-magmatic origin replace plagioclase in the second phase of the intrusion. Hornblende occurs sporadically in granites of the first phase. The other accessory minerals are apatite, titanite, zircon and epidote. The opaque minerals are represented by magnetite and ilmenite.
The Neeme Pluton (25km × 20 km), the northern part of which is under the Gulf of Finland, is composed of coarse- and medium-grained pinkish-grey porphyritic granites cut by aplites (Soesoo & Niin 1992). By chemical and mineral composition, the rocks form two groups, possibly two phases. Two small bodies in the central and northeastern parts of the intrusion are more melanocratic, their chemical composition being partially similar to granodiorite. Partially assimilated xenoliths of surrounding gneisses with a diameter of about 20–30 cm have been found in some drill cores. The structure of rocks is massive, in places trachytoidic.
Quartz (20–25%) is generally euhedral, smaller grains are anhedral. Microcline (35–45%) occurs in groundmass in the form of rare phenocryst, up to 3–5 cm in diameter. Plagioclase (15–25%) is represented by euhedral crystals of oligoclase-andesine composition. Biotite (2–10%) forms unhedral flakes that contain small crystals of apatite, titanite, fluorite and zircon as inclusions. Hornblende occurs sporadically. Muscovite, apatite, fluorite, titanite, zircon, epidote and opaques are accessory minerals.
The Ereda Pluton (5 km × 15 km) is composed of homogeneous pinkish‑grey coarse‑grained porphyritic granites (Soesoo & Niin 1992). As there are only two drill cores available, it is difficult to correlate the Ereda rocks with those of other stocks. The mineral composition and structure of the Ereda granites and the Märjamaa and Neeme leucocratic type of granites have some similar features.
Two generations of quartz (30–35%) have been distinguished. Microcline (35–45%) is present as tabular crystals; large phenocrysts with a dimater of 3–4 cm are zoned. Plagioclase (20–30%) forms various euhedral, tabular and prismatic crystals of andesine composition. Biotite (5–10%) is altered. The accessory minerals are fluorite, apatite, zircon, epidote and rutile. Magnetite and hematite opaques occur.
The Märjamaa Pluton (40 km × 25 km) is composed of coarse-grained pink-grey porphyritic granitoids, sometimes cut by aplites (Soesoo & Niin 1992). According to the geophysical and drilling data, the pluton has features of ring structure and is accompanied by a smaller satellite. The contacts between the granites and surrounding Palaeoproterozoic gneisses are sharp. The round central part of the Märjamaa composite stock is represented by the most melanocratic and basic type of porphyritic granodiorites (first phase). In places it contains gneiss xenoliths, up to 20 cm in diameter. The second, intrusive phase, is represented by biotite and hornblende-bearing granites. The granites of the third phase (possibly the small individual Kloostri Massif in the northwestern part of the Märjamaa Intrusive) are more leucocratic in composition and, in places, have a trachytoidic structure.
Two generations of quartz (20–30 %) have been distinguished. The first generation consists partly of anhedral crystals between potassium feldspar individuals and partly of euhedral inclusions within microcline. The second generation occurs as anhedral crystals in the groundmass. Microcline (20–40%) is present as phenocrysts (diameter about 2–3 cm) and in the groundmass. The phenocrysts are often perthitic and contain inclusions of quartz, biotite and rare, titanite. Plagioclase (20–40%), which forms euhedral tabular or prismatic crystals, is represented by oligo-clase‑andesine. Anhedral crystals of biotite (2–10%) are often clustered together as swarms of small or large flakes. Hornblende (mainly in the first and, partly, in the second intrusive phase) and muscovite (in the third phase) occur sporadically. The content of the main opaque minerals, magnetite and ilmenite may reach 3–5%. Apatite, fluorite, zircon, titanite and epidote are accessory minerals.
The Taebla body, 6–7 km in diameter, is the smallest one distinguished by drilling of two wells. It is composed of homogeneous leucocratic porphyritic granites (Soesoo & Niin 1992). In terms of the mineral and chemical composition, the Taebla granites are similar to the rocks of the third phase of the Märjamaa Pluton and to the second phase of the Naissaar and Neeme plutons.
Geophysical data indicate that the Abja body (8 km × 5 km), which occurs in the structural zone of South-Estonian granulites, is ellipse-shaped. In the only drill core available, in the depth interval 550-635 m, medium-grained greenish-grey gabbrodiorites (SiO2 49–52%) with massive texture are cut by fine- and medium-grained pinkish-red potassium granites. The isotopic age (U-Pb, zircon of the gabbrodiorites is 1635 ± 7 Ma (Kirs & Petersell 1994). The age of the intersecting granites is 1622 ± 7 Ma. The gabbrodiorites comprise relatively euhedral plagioclase (An 33–39, 40–50%), hornblende (10–20%) and biotite (10–20%) with minor quartz and potassium feldspar. The content of accessory and opaque minerals is quite high; the most important being apatite (2–5%), titanite (1%) and titanomagnetite (2–6%).
Geophysical data suggests that the Sigula body of mafic rocks is a NE-trending dike (1.5km × 4 km). The one drill core available to date (depth interval 223.2-316.6 m) reveals that the pluton consists of inequigranular dark-grey massive diabase (SiO2 47–49%) with ophitic structure (Puura et al. 1983). The relatively large prismatic plagioclase crystals provide the rock with a slightly porphyritic outlook. The amount of plagioclase (An 55–63) is remarkable reaching 50–60%. Other minerals are hornblende (8–10%) and clinopyroxene (8–10%), biotite, orthopyroxene, quartz and potassium feldspar (all < 5%). The content of apatite (2–5%) and, especially, titanomagnetite (7–10%) is noticeably high.
The largest, Riga complex rapakivi‑anorthosite pluton (Bogatikov & Birkis 1973, Kuuspalu 1975) spread in the basement of western Latvia, southwestern Estonian Archipelago in the Gulf of Riga and Central Baltic proper, and the Åland Pluton belong to the Åland-Riga Subprovince. Defined and supposed rapakivi bodies in the northern Baltic seabed occur near the West-Estonian Archipelago (Puura et al. 1992, Koistinen 1994). It has been mentioned that the most intense rapakivi plutonism area coincides with the junction of the Baltic proper with the gulfs of Bothnia, Finland and Riga (Puura & Flodén 1996).
The rapakivi plutons and stocks are characterised by changeable magnetic and stable density properties and, as a whole, they have a considerably massive internal structure. Thus, they are easy to identify and contour by geophysical mapping and drilling.
The Riga Pluton is an essential representative of bimodal rapakivi-anorthosite complexes. The mafic part of the pluton locates in its southern part, in southwestern Latvia. In the central part, on the Kurzeme Peninsula, a variety of both typical vyborgite- and pyterlite-like and even-grained granites occurs. In the central and southern parts of the Riga Pluton, also quartz mangerites, mangeritic granosyenites and quartz‑syenites occur among mangeritic granitoids (Bogatikov & Birkis 1973). All these rocks have been formed in deep levels of the crust. However, in cores of two wells penetrating into the Riga Pluton in southwestern Estonia, on the islands of Ruhnu (3.4 m of crystalline rocks) and Saaremaa (at Kuressaare Town, 28.4 m), the rocks are represented by subvolcanic granite porphyries (granophyres) with micropegmatite matrix (Kuuspalu 1975, Puura et al. 1983).
Exterior to the northern part of the Riga Pluton (Fig. 5), the Undva well penetrates into a suite of rapakivi‑related volcanic rocks (Puura et al. 1983). The lower part of the sequence consists of plagioclase porphyrites, and the upper part of quartz porphyries.
Plagioclase porphyrites have some similarity with the rapakivi-related, more felsic porphyrites on Hogland, and with Dala porphyrites from the Transscandinavian Igneous Belt in central Sweden. The rocks are dark grey or black, in places, with pink-shaded massive and dense rocks. Their groundmass is fine- or very fine-grained, with microophitic texture, and consists of plagioclase (An 50–65, 65–75%), clino- and orthopyroxene (15–25%), hornblende (<5%), biotite, titanomagnetite, hematite and apatite. Prismatic and tabular phenocrysts of plagioclase (An 45–70), with an average size of 4mm × 5 mm (occasionally 30-40 mm), are rare forming about 3–10%. In terms of the chemical composition, plagioclase porphyrites are similar to andesites.
The quartz porphyries are brownish-red or pink massive rocks. Their fine-grained groundmass consists of quartz (30–40%), feldspar (40–50%), opaque minerals (hematite and magnetite up to 15%), chlorite, apatite and glass, and they have granophyric, radially fibrous and spherolitic texture. Small rounded phenocrysts of dark grey quartz are about 3–4 mm in diameter and make up 3–10% of the rock. Prismatic phenocrysts of plagioclase (An 1–7) and microcline-perthite are a bit larger (diameter about 5–8 mm, rarely 10–20 mm); their quantity varies between 20 and 30%.
The quartz porphyries of the Undva Member (Table 3) differ from those on Suursaari (Hogland) Island in colour and texture of groundmass. They comprise less and smaller phenocrysts, but the content of opaque minerals and apatite is higher which makes them more similar to some Dala porphyries in the Transscandiavian Igneous Belt in central Sweden.
K. Mens & E. Pirrus
The Vendian (Vendian Complex) as an independent stratigraphic unit, probably in the category of system, was distinguished in the early 1950s by B. Sokolov (1952a, 1953). Its stratotype area is in the western part of the East-European Platform.
The Vendian of the stratotype area includes three subdivisions in the rank of regional series, which in ascending order are Vilchan, Volyn and Valdai (Keller & Rozanov 1979b). The Valdai regional series consists of the Redkino (below) and Kotlin (above) stages, of which only the latter occurs in Estonia and forms the lowermost part of the sedimentary cover overlying the Proterozoic crystalline basement.
The current stratigraphic scheme of the Estonian Vendian was accepted in 1976 at the Baltic Stratigraphic Conference in Vilnius (Table 5).
The unit in the rank of stage was defined as the upper part of the Valdai Series corresponding to “Laminarites” Clay on the East-European Platform (Resheniya… 1965). The name Kotlin was proposed by Sokolov (Männil 1958a) after Kotlin Island in the eastern part of the Gulf of Finland. Mens and Pirrus (1974, 1980, Gnilovskaya et al. 1979, Keller & Rozanov 1979b) determined the present stratigraphic extent of the stage and worked out its classification for the East Baltic area.
The Kotlin Stage is widespread in mainland Estonia, lacking only in its southwestern part and in some local structures, including Assamalla and Uljaste (Fig. 11). Stratigraphically, the most representative and thickest sections are situated in northeastern Estonia. In a westerly direction, the sections thin out rather rapidly and change in lithology.
The lower boundary of the stage coincides with the base of the sedimentary cover in Estonia, and is easy to determine. Some complications occur if the core yield is low or the core is distorted. The upper boundary is clear in eastern and central Estonia where the overlying rocks contain glauconite and mineralized skeletal fossils. In the northwestern part of mainland Estonia and on Hiiumaa Island the boundary is less distinct due to the lithological similarity with the overlying Lower Cambrian rocks. Identification of the Kotlin Stage is most complicated in the sections west of Keila (Fig. 12) where the lower part of the sedimentary cover comprises light-coloured loose quartzose sandstones with occasional lenses or interbeds of compacted multicoloured argillaceous rocks. As the latter rock type is lacking in the overlying Cambrian beds, this part of the sequence is conditionally regarded as the Kotlin Stage.
The Kotlin Stage is represented by siliciclastic rocks which accumulated under cool and humid climatic conditions (Pirrus 1992). This extensive, high-order cycle of deposition covered three shorter cycles divided as successive Gdov, Kotlin and Voronka formations. Multicoloured sandy-silty sediments consisting of low maturity and poorly sorted detrital material accumulated at the beginning of the Kotlin depositional cycle (Gdov Formation). Upwards in the section, the coarse-grained red-coloured deposits change into grey-coloured clayey sediments which accumulated during the stable phase of the Kotlin depositional cycle (Kotlin Formation). The cycle ends with the reappearance of multicoloured sediments of high maturity (Voronka Formation).
In recent years, acritarchs as the most abundant and widespread fossil group in the deposits of the Kotlin Stage have underlain the subdivision and correlation of the Vendian rocks (Volkova et al. 1983). Acritarchs are represented by a taxonomically simple assemblage consisting mainly of representatives of the genus Leiosphaeridia (Paškevičiene 1980). Microfossils are accompanied by vendotaenids, of which Vendotaenia antiqua Gn. is most common, while Aataenia reticularis Gn. is rare (Gnilovskaya et al. 1979). Besides microfossils and vendotaenids, fragments of shapeless organic matter occur on the bedding surface. All the above-listed palaeontological finds occur in the rocks of the Kotlin Formation which have promoted their accumulation and preservation. In some sections in the northeasternmost part of Estonia (Meriküla, Sinimäe), the grey argillaceous rocks of the upper member of the Gdov Formation comprise acritarchs and organic matter of irregular form.
The Gdov Formation. Asatkin (1937) derived the name from the Gdov beds used as a division to denote the sandy strata between the “Laminarites” Clay and the crystalline basement in the northwestern part of the East-European Platform. The Gdov Formation is considered as the lower part of the Kotlin Stage accumulated during the initial phase of the Late Valdaian transgression over the northwesternmost part of the East-European Platform, including the present-day Estonia, Latvia, and the western part of the Leningrad Region.
In Estonia, the Gdov Formation rests immediately upon the crystalline basement and spreads in subsurface lying in the northern, eastern and central parts of the Republic. Its thickness ranges from 0.2 to 58.3 m (Fig. 11-203, 11-102). The Venevere (Fig. 11-86) drill core in the interval of 287–322.8 m has been selected as a hypostratotype for Estonia (Fig. 12). The formation prevalently consists of multicoloured sandstones of various grain-size. The uppermost and, locally, also the lowermost part comprises a considerable quantity of reddish and purplish argillaceous rocks. The sandstones are represented by arkose and feldspatic varieties comprising besides quartz up to 50% of feldspars. Micas, both muscovite and green altered biotite, are occasionally found. The mineral composition of the clay fraction is rather stable throughout the formation, being characterised by illite-kaolinite suite (Pirrus 1970). On the basis of lithological features, the Gdov Formation is subdivided into three members which in ascending order are Oru, Moldova and Uusküla (Table 5).
The Oru Member occurs locally in the base of the Gdov Formation. It accumulated in depressions of the crystalline basement on account of its weathering crust. The greatest thickness of the Oru Member (6.7 m) has been recorded in the Jaama borehole (Fig. 11-104). The member consists of red unsorted clayey-sandy-gravely deposits (mixtite). Unlike the overlying members, the deposits of the Oru Member differ both structurally and mineralogically. Quartz is the prevailing mineral in sand and gravel (up to 90%); its grains are angular or subangular with nonsorted size distribution. Feldspar is not common (less than 10%). The clayey matrix consists mostly of kaolinite. All this suggests that these deposits were formed in deluvial fans as a result of intense weathering of acid crystalline rocks.
The Moldova Member overlies either the Oru Member or the crystalline basement. Its thickness is usually 30–40 m, maximum 50 m. The member consists of yellowish or pinkish arkose and/or feldspatic sandstones of various grain size and a few interlayers of multicoloured, frequently reddish, argillaceous rocks. Hence, the clastic material transported into the sedimentary basin originates from a close-lying area with a low degree of weathering of rocks.
The Uusküla Member is the most fine-grained and multicoloured part of the Gdov Formation. It consists of silty claystones intercalated with silt- and sandstones. The number and thickness of claystone layers increases eastwards. In the same direction the rocks gradually loose the red colour until in the easternmost sections they are predominantly grey. The proportion of sandstones is low and they are mainly represented by arkose and feldspatic types. Micas occur in remarkable quantities, but rock fragments are uncommon.
The composition and structure of the rocks suggest a rather low hydrodynamic energy of the sedimentation basin.
The Kotlin Formation. The name was introduced by Sokolov to designate the “Laminarites” Clay (Männil 1958a). The formation is spread in eastern Estonia, in a more typical form in its northeastern part attaining a thickness of 52.6 m in the Narva borehole (Fig. 11-33). In the west direction the thickness decreases quickly pinching out on the Tapa - Ellavere line (Figs. 12, 13). The formation is known only from core sections, and the interval of 109–150 m of the Meriküla core (Fig. 11-32) has been defined as the hypostratotype. The lower boundary is drawn at the level where multicoloured deposits turn grey. At the base of the formation, gravel and coarse-grained sand occur locally.
The dominant components of the formation are thinly laminated grey claystones with intercalating light-coloured very fine-grained sandstones or siltstones, or both. The lamination is complicated by the occurrence of dark-brown films of organic matter.
The rocks of the formation are low in sand and silt, the content of which in the upper- and lowermost parts only locally exceeds 50%. Quartz and feldspars (particularly K-feldspar) are the main detrital minerals of sand and gravel grain-size. The content of micas, including biotite and muscovite, is also notable. Their ratios depend on the type of rock. The rocks are characterised by a small content of both opaque and transparent allogenic minerals. Heavy minerals are dominated by siderite and pyrite of authigenic origin. Illite is a dominant clay mineral, the content of kaolinite ranges from 15 to 40%, the latter value being fixed in the lowermost part. Chlorite is common, in the middle part of the formation its average content is 15–20% (Pirrus 1970).
On the basis of the lithological composition, the Kotlin Formation is divided into three members (Table 5).
The Jaama Member is made up of alternating grey-coloured massive siltstones and thinly laminated claystones. A few siderite nodules and organic matter films occur. This is the first stage in the large-scale clay accumulation in the eastern part of Estonia.
The Meriküla Member is the most typical unit of the Kotlin Formation. It is represented by the “Laminarites” Clay consisting of rhytmically alternating 0.5–0.8 mm thick pairs of dark-grey fine-dispersed clay layers and light-grey laminae higher in silt. The bedding plane is covered by dark-brown shapeless organic films. Vendotaenides, small flakes of mica and siderite nodules are common.
The fairly uniform mineral composition shows that the source areas must have been located relatively far from the depositional basin.
The Laagna Member has the most restricted distribution area, compared to other members of the formation. In the northeasternmost part of Estonia it is up to 6 m thick. The member consists of grey clayey and silty-clayey argillaceous rocks with many up-to-20-cm-thick intercalations of siltstone. The typical “Laminariates” Clay layers are absent. Scarce organic matter films and small siderite nodules occur suggesting the terminal phase of clay accumulation under weak hydrodynamic conditions.
The Voronka Formation was established by Mens and Pirrus (1971). Earlier, this part of the sequence was treated as two lower units of the post-Laminarites Sandstone or as the lower and middle parts of the Lomonossov Formation (Mardla et al. 1968). The type section of the formation is an outcrop on the lower reaches of the Voronka River, Russia (Mens & Pirrus 1971). Beyond the stratotype area, the formation is of subsurface occurrence being known in eastern and northern Estonia and in eastern Latvia. The Meriküla (Fig. 11-32) drill core in the interval of 90–109 m serves as a hypostratotype for the Voronka Formation (Mens & Pirrus 1980). The formation occurs between the overlying Lontova Formation and the weathering crust of the underlying Kotlin Formation (Mens & Pirrus 1969, 1970). In Estonia, the thickness of the formation ranges from 10 to 40 m. The Voronka Formation consists of variable siliciclastic rocks and represents a single upwards coarsening cycle from argillaceous rocks to well-sorted sandstones. The lower boundary of the formation is drawn on the basis of the change in colour. Based on lithological evidence, the formation is divided into the Sirgala and Kannuka members (Table 5).
The Sirgala Member consists of alternating multicoloured clays and siltstones with interlayers and lenses of light-coloured sandstones, the share of which increases upward the section. Most of detrital grains are subrounded quartz with a small quantity of feldspars (up to 10%) and micas (mainly muscovite). In the clay fraction, kaolinite slightly prevails over illite. Chlorite is uncommon.
The mineral composition suggests that these deposits derived from the weathered zone of sedimentary rocks.
The Kannuka Member consists entirely of light weakly cemented fine- to medium-grained quartzose sandstones with a few thin interlayers of multicoloured clayey siltstones, which are similar to the underlying deposits of the Sirgala Member. The clay fraction is dominated by kaolinite. Increase in the maturity in minerals upward the section is indicative of the redeposition of older sediments.
K. Mens & E. Pirrus
Cambrian rocks are widespread in Estonia. They are missing on the crest of the Valmiera-Lokno swell and on some peninsulas on the southern coast of the Gulf of Finland. Exposed Cambrian rocks are encountered in outcrops along the Baltic Klint, but mostly they are overlain by younger rocks and the basic data for studies has been obtained by means of boreholes.
The main pioneering work towards the subdivision of the Estonian Cambrian (then the Lower Silurian) was done by Eichwald (1854a) and Schmidt (1888) who worked out the lithostratigraphical subdivision and described the first fossils found in these rocks. A zonal division and modern nomenclature were introduced by Öpik (1933, 1956).
Up to the middle of this century, the Cambrian stratigraphy was based on outcrop sections and embraced the lowermost part of the Lower Cambrian succession of the studied area and the problematic Acrotreta Zone of the Upper Cambrian (Öpik 1956).
During the last fifty years, numerous deep borings were made which revealed the full thickness of the Cambrian rocks. Elaboration of Cambrian stratigraphy was greatly promoted by identification of plant microfossils (now known as acritarchs) by Naumova, Timofejev and Volkova on the East- European Platform. Palynological studies provided valuable data for establishing distinct acritarch assemblages, their ranges in the sequence and relationship with trilobite zones.
Elaboration of the present-day stratigraphical subdivision of the Estonian Cambrian (Table 6) was favourably influenced by international cooperation with researchers from neighbouring countries, and supported by IGCP projects No. 29 and 86. The results obtained were summarized in the stratigraphic scheme accepted in 1976 at the Baltic Stratigraphic Conference in Vilnius (Resheniya… 1978) and improved in 1983 by the Stratigraphic Conference on the Cambrian of the East-European Platform (Resheniye… 1986).
As the global standard is still under preparation, the boundaries between the Lower/Middle and Middle/Upper Cambrian are not strictly formal. There are no generally agreed names for the stages and their boundaries are unclear.
The Cambrian stratigraphic scale in the East-European Platform from the base of the Sabellidites cambriensis Zone to the top of the Acerocare Zone is based mainly upon the succession of trilobites, except the lowermost part where trilobites are lacking (Mens et al. 1990). Only part of the Cambrian is present in Estonia (Table 6). The lower boundary of the system is distinct in the studied area, and coincides with regional changes in the sedimentary conditions which led to the accumulation of normal marine sediments (Pirrus 1993). This level is marked by the appearance of primitive skeleton-forming organisms and changes in the composition of ichno- and phytofossils.
The upper boundary of the system is not obvious although, biostratigraphically, the Cambrian‑Ordovician transition in Estonia is relatively well studied (Mens et al. 1993). This is due to the circumstance that the IUGS has not yet passed the final decision on the Ordovician lower boundary.
Conventionally, the Cambrian is subdivided into three subsystems: Lower, Middle and Upper. The Lower/Middle Cambrian boundary is at the base of beds with Paradoxides, more exactly Eccaparadoxides insularis, and the Middle/Upper Cambrian boundary is at the base of the Agnostus pisiformis Zone for the East-European Platform (Table 6).
Rocks of all three subsystems are encountered in Estonia, but the degree of completeness varies. Compared to other subsystems, the Lower Cambrian rocks are most widespread and thickest. Their two-folded structure results from remolding of the basin prior to the Liivi transgression. Of the six regional stages established on the ground of the succession of acritarch assemblages in the Lower Cambrian on the platform, four are present in Estonia (Table 6).
The Middle Cambrian sequence in Estonia is entirely devoid of fossils and the regional stages have been established on the basis of lithological criteria.
The Upper Cambrian is documented on the basis of palaeontological evidence, derived from both the shelly fauna and acritarchs.
No regional stages (except the Kybartai and Deimena in the lowermost Middle Cambrian) have yet been differentiated in the rest of the Middle and through the Upper Cambrian in Estonia. The relevant rocks have been treated only in general lines by lithostratigraphic units.
Based on the stratigraphical completeness of the sections and facies conditions, Estonia’s territory is subdivided into the northern, western and southeastern regions (Brangulis et al. 1974, 1975, Table 6), with their characterisitic formations and members.
The oldest Cambrian rocks in Estonia were formed during the Baltic evolutionary stage in the pre-trilobite Early Cambrian (Mens 1981c). The onset of sedimentation in the Early Cambrian in Estonia corresponds to the Platysolenites antiquissimus Zone defined as the Lontova Stage. The Rovno Stage, composed of the lowermost rocks of the Baltic evolutionary stage, is lacking in Estonia (Table 6).
The name Lontova was introduced by Öpik (1933) in the rank of beds to designate the “Blue Clay” proper. It corresponds to the upper part of the blue clays by Schmidt (1888) and Mickwitz (1911), the Lontova beds by Öpik without the uppermost layers with Volborthella tenuis (Öpik 1933, 1956) and to the Platysolenites antiquissimus Zone in the current use (Mens et al. 1990).
The rocks of the Lontova Stage crop out at the foot of the Klint and extend as a narrow belt from Tallinn to Narva. The main localities are the quarries at Kopli, Tammneeme, Kolgaküla, Kunda and Aseri (Fig. 14).
The stratotype of the stage is the Kunda quarry (Öpik 1933), subsequently complemented by the Lontova drill core in the interval of 14.0 to 88.3 m (Mens & Pirrus 1977). The stage in the stratotype section is incomplete since the deposits of the regressive phase of the Baltic sedimentary cycle are lacking. The stratigraphical completeness and thickness of the stage varies with regions: the rocks are at their thickest (ca 90 m) in northeastern Estonia (Fig. 14) and thin in a southerly direction due to post-sedimentary denudation. In Estonia, the Lontova Stage overlies, with a break in sedimentation, the Kotlin Stage. Its lower boundary, known only from core sections, coincides with the appearance of typical marine sediments in the succession (Mens & Pirrus 1977, Pirrus 1993).
The Lontova Stage is represented by siliciclastic rocks with a clear lateral variation of the ratio of rock types. Argillaceous rocks are prevailing in eastern and central Estonia, while sandstones dominate west of the Vihterpalu ‑ Häädemeeste line (Fig. 14).
A relatively diverse assemblage of skeletal fossils containing Sabellidites cambriensis Yan., S. sp., Platysolenites antiquissimus Eichw., P. lontova Öpik, P. spiralis Posti, Yanichevskyites petropolitanus (Yan.), Aldanella kunda (Öpik) together with pyritized casts of hyolithids and hyolitelmintes, hornlike chitinous (?) sklerits, fragments of brachiopods and agglutinated tubes of Onuphionella has been identified from the Lontova Stage (Mens & Pirrus 1977, Mens & Posti 1984). Some of the above-mentioned species like P. antiquissimus, Y. petropolitanus and casts of hyolithids occur throughout the stage, whereas the vertical range of the rest is more limited. On the basis of the earliest appearances of the index taxa, the Lontova Stage is subdivided into four parts, which in ascending order are the Sabellitides cambriensis beds, Platysolenites lontova beds, Aldanella kunda beds and P. spiralis beds (Mens & Posti 1984).
The distribution of some species shows a distinct facies control. Thus, the hornlike chitinous (?) sklerites have been found in the argillaceous rocks of the eastern part of the territory only. The tubes of Platysolenites and Yanichevskyites are rather rare in the well-sorted sandstones in the western part of the studied area.
Acritarchs from the Lontova Stage have been described by several investigators (Naumova 1960, Volkova 1968, 1973; Jankauskas & Posti 1973, a.o.), who all agree that the stage has an acritarch assemblage of its own, which contains besides leiosphaerides and tasmanites also marginats forms. The frequency and diversity of acritarchs in the assemblage clearly depend on palaeoenvironmental conditions and facies changes in the basin of sedimentation (Mens & Paškevičiene 1981).
Trace fossils from the Lontova Stage are diverse and comprise numerous ichnospecies, among them Phycodes pedum Seilacher (Palij et al. 1983).
In conformity with the ratio of rock types in the succession, two formations lateratelly replacing each other, have been distinguished in the Lontova Stage (Kala et al. 1981b).
The Lontova Formation was identified in the rank of formation by Männil in 1958 (Männil 1958a). Its type section in the Kunda quarry has been selected for the stratotype of the Lontova Stage (see above). The formation occurs in northern, eastern and central Estonia (Fig. 14), and is westwards laterally replaced by the Voosi Formation. The Lontova Formation is represented by greenish-grey and variegated argillaceous rocks with interbeds of coarse- to fine-grained sandstone in the lowermost and fine-grained sandstones in the uppermost part. The formation is subdivided in ascending order into the Sämi, Mahu, Kestla and Tammneeme members (Kala et al. 1970, Mens & Pirrus 1977).
The Sämi Member consists of alternating sandstones and argillaceous rocks containing glauconite and, occasionally, also flattened phosphatized pebbles.
The Mahu Member is made up of greenish-grey sandy or silty claystones with thin interlayers of sandstone.
The Kestla Member is characterized by homogenous multicoloured claystones with greenish‑grey, reddish‑brown and purplish interbeds, strips and spots. The admixture of sandy material is limited.
The Tammneeme Member consists of fine‑grained sandstones and greenish‑grey claystones and occurs in a limited area (Fig.15).
The Voosi Formation was defined on the basis of lithological evidence. Sandstones account for more than 50% of its composition (Kala et al. 1981b). Its stratotype is the interval of 237.5 to 300 m in the Haapsalu‑3 drill core (Fig. 14-124)
The formation is distributed in the northwestern part of Estonia; in a easterly direction it is gradually replaced by the Lontova Formation. Its lower part extends farther to the east (Fig.16). The thickness of the formation ranges from 72 to 14.6 m.
The formation consists mostly of quartzose sandstones which is the dominant type of rock on the islands of the West-Estonian Archipelago. Claystones are of minor importance and associate mostly with the upper part of the formation in mainland Estonia.
The formation is subdivided in ascending order into the Taebla, Kasari and Paralepa members.
The Taebla Member consists of light fine-grained sandstones with a few thin interbeds of silty claystones. Glauconite is not common.
The Kasari Member is represented by sandstones of various grain-size and with limited claystone content. The sandstones are rather rich in glauconite and in some places flat phosphatized pebbles occur. The sandstones of the Kasari Member are quite similar to those in the Sämi Member of the Lontova Formation.
The Paralepa Member consists of interbedded greenish‑grey (with a few purplish spots) argillaceous rocks, mainly silty claystones, and of fine‑grained sandstones.
According to the currently accepted correlation (Resheniye… 1986), the lowermost part of the trilobite-bearing Cambrian on the East-European Platform, deposited during the Liivi evolutionary stage in the East Baltic area (Mens 1981c), is defined as the Dominopol’ Stage.
Previously, this unit in the same stratigraphical extent was referred to as the Lükati Stage (Aren et al. 1975, Keller & Rozanov 1979a), as the Talsy (= Lükati) Stage (Keller & Rozanov 1979b) or as the Talsy Stage (Birkis et al. 1970, Brangulis et al. 1981). It should be pointed out that the name Lükati in the rank of stage is also used in a more restricted stratigraphical extent (Mardla et al. 1968, Mens 1986).
The Dominopol’ Stage was distinguished by Kirjanov (Kirjanov 1969) with the stratotype section in the interval of 617.2 to 747 m of the Berezhki 2944 drill core situated in Volyn, the western Ukraine.
In Estonia, the Dominopol’ Stage is represented by three succeeding formations: Sõru, Lükati and Tiskre (Table 6) well recognizable on the basis of lithological and palaeontological evidence. The Dominopol’ Stage occurs on the islands of the West- Estonian Archipelago (except Ruhnu) and in the western, northern and central parts of mainland Estonia (Fig. 17). Only the upper part of the stage (Lükati and Tiskre formations) is exposed along the North-Estonian Klint. The outcrop extends from the Pakri Peninsula in the west up to the Narva River in the east. The main localities are Türisalu, Rannamõisa, Kakumägi, Lükati (Photo 15), Saviranna, Kunda, and Utria (Fig 17). The maximum thickness of the stage (76.6 m) has been fixed in the Kalana borehole where all three formations occur.
The stage consists of siliciclastic rocks, mainly sandstones. The lower boundary of the stage is lithologically distinct, accompanied by a change in the faunal composition and marked, in some places, by lenses of conglomerate.
The palaeontological finds in the lower part of the stage (Sõru Formation) are scarce. These are mostly trace fossils, rare shells of agglutinated foraminifers and a few acritarchs. The latter are represented by leiosphaerides and rare Globosphaeridium cf. cerinum (Volk.), Asteridium pallidum (Volk.), Loposphaeridium tentativum Volk. and Tasmanites bobrowskae Waz. (Mens 1986). This part of the stage corresponds to the Rusophycus parallelum Zone.
The middle part of the stage (Lükati Formation) is palaeontologically well characterized. The most typical species include Volborthella tenuis Schm., V. conica Schindewolf, Schmidtiellus mickwitzi (Schm.), Mickwitzia monilifera (Linnars.), torellellids, hyolithids, agglutinated foraminifers (“Lycatiella”), and in the west Platysolenites antiquissimus Eichw. has been identified in the basal beds. The acritarch assemblage in this part of the stage is abundant and diverse, with baltisphaerids dominating (Mens & Pirrus 1977). The middle and upper parts of the stage correspond to the Schmidtiellus mickwitzi Zone.
Of the upper part of the stage (Tiskre Formation), only its lower part (Kakumägi Member) is palaeontologically well characterized. It contains, particularly in the conglomerate lenses, Mickwitzia monilifera (Linnars.), M. formosa Wiman, M. concentrica Gorjansky, Paterina rara Gorjansky, Scenella discinoides Schm., S. tuberculata Schm., Bradoria? estonica Melnikova, Konicekion kundaensis Melnikova and fragments of trilobites. In the uppermost part (Rannamõisa Member), only occasional indeterminable fragments of brachiopods and trace fossils occur. Acritarchs have been identified in the Kakumägi Member, and only once they were found in the drill core of the Rannamõisa Member (Muraste-2 borehole). Its assemblage is much the same as in the Lükati Formation, except the appearance of Tasmanites piritaensis Posti et Jank.
Based on the differences in the palaeontological composition and distribution area as well as clear contacts between the three formations of the Dominopol‘ Stage, these parts of the sequence are considered independent stages (Mens 1986).
The Sõru Formation, resting transgressively either on the Lontova Formation or on the crystalline basement, occurs in the northwestern part of mainland Estonia and on the islands of the West-Estonian Archipelago, except Ruhnu (Fig.18).
Rocks of this formation are known only by core sections. The thickness of the formation ranges from 6.2 m at Vihterpalu to 58.2 m at Eikla (Fig.17-40, 178). The Tahkuna drill core (Fig. 17-34) in northern Hiiumaa in the interval of 100.5 to 147 m has been selected as the type section of the Sõru Formation (Resheniya… 1978, Kala et al. 1984a).
The lower part of the formation consists mostly of massive fine-grained quartzose-feldspatic sandstone with thin clay interbeds and films. The upper part is represented by a complex of interbedded argillaceous rocks and sandstones. In both parts the rocks are light-grey with greenish-grey shade, but red and purplish-red patches also occur .
The Lükati Formation, the most widespread division of the Dominopol’ Stage (Figs. 18, 19), lies transgressively on the Sõru Formation in the west and on the Lontova Formation in the east. It is often separated from the underlying units by conglomerate lenses containing pebbles of phosphatized sandstones (Mens & Pirrus 1975). Outside the distribution area of the Tiskre Formation, the upper part of the Lükati Formation is often weathered. The Lükati Formation, formed during the stable phase of the Liivi evolutionary stage, corresponds to the whole stratigraphical extent of the Dominopol’ Stage along the southern margin of its distribution area in Estonia.
In the type area in the vicinity of Tallinn, the formation reaches 20 m in thickness. The type section of the formation is an outcrop on the left bank of the Pirita River (Photo 15) in the eastern outskirts of Tallinn (Öpik 1933), complemented today by a drill core from the lower part of the formation (Mens & Pirrus 1977).
The lithologically monotonous formation consists of interbedded greenish-grey argillaceous rocks and very fine-grained sandstones. The latter form mostly 0.1—0.3-m-thick layers, some of which are hard, in places tightly cemented by poikilotopic carbonate.The upper surfaces of the sandstone layers are covered with ripple marks, the lower surfaces with casts of various trace fossils and mud cracks. Pyrite and glauconite are very common. In the lower part, glauconite often forms 1 - 3-cm-thick laminae.
The Tiskre Formation in the current use is interpreted as a unit containing the Tiskre beds (= Diplocraterion Sandstone) by Öpik (1933) or the Tiskre Formation by Männil (Männil 1958a) together with the Kakumägi beds (Scenella Zone) by Öpik (1933) or the Kakumägi Member of the Pirita Formation by Männil (Männil 1958a).
The Tiskre Formation is distributed in northern and western Estonia (Figs. 18, 19) where it overlies the Lükati Formation. The lower boundary is lithologically abrupt and marked by a change from argillaceous rocks to sandstones. Between Muraste and Aseri, lenses of Mickwitzia conglomerate occur at that level. Its type section is an exposure at the southern end of the Rannamõisa Cliff, 14 km west of Tallinn (Mens & Pirrus 1977). The formation is at its thickest (20 m) in the stratotype area.
The Tiskre Formation consists of light-coloured massive or thick-bedded sandstones with thin interbeds of greenish-grey argillaceous rocks (Photo 16). Within the outcrop belt the Tiskre Formation is divided into the Kakumägi (lower) and Rannamõisa (upper) members.
The Kakumägi Member is represented by poorly sorted sandstones with an admixture of clayey material. Conglomerate lenses within the sandy sequence are locally present. In the basal part, sandstones are often well cemented with dolomitic cement. Bedding is mostly lenticular, casts of mud cracks, slump-rolls and ripple marks are common.
The Rannamõisa Member consists of horizontally bedded winnowed fine-grained or very fine-grained sandstones with thin interlayers of argillaceous rocks. Glauconite is present. Ripple marks and convolute bedding occur throughout the member (Pirrus 1978).
The succeeding Aisčiai evolutionary stage terminates the Early Cambrian sedimentation in Estonia, and embraces deposits of the Ljuboml’ and Vērgale stages (Table 6). The former was earlier treated as a part of the Vērgale Stage (Mens et al. 1990) or as the Lower Vērgale Substage (Resheniye… 1986). Kirjanov (Kirjanov 1969) distinguished it in the rank of an independent stage and also as a formation with the stratotype section in the interval of 543.5 to 617.2 m of the Berezhki‑2944 drill core in Volyn, the western Ukraine.
The Ljuboml’ Stage is represented in western Estonia by the Soela Formation and in central and eastern Estonia by the Vaki Formation (Resheniya… 1978, Mens 1979, Kala et al. 1984a), and is known only from core sections. Since the boundary between the Ljuboml’ Stage and the overlying Vērgale Stage is difficult to determine, a map showing the distribution of the late Lower Cambrian rocks in Estonia was compiled jointly for these units (Fig. 20).
The Ljuboml’ Stage lies with a stratigraphic unconformity on the rocks of the Dominopol’ or Lontova Stage or on the crystalline basement (Ruhnu Island, borehole 257). Accordingly, the lower boundary is expressed variously. It is well recognizable in the areas where the underlying strata consist of crystalline or argillaceous rocks. In the latter case, the topmost part of the Lontova or Lükati Formation is often weathered. Identification of the lower boundary is most complicated in the sections where the Tiskre Formation of the Dominopol’ Stage and the Soela or Vaki Formation of the Ljuboml’ Stage occur simultaneously, because they are lithologically very similar and extremely poor in fossils. In that case, the lower boundary of the stage is tentatively drawn at the level of the essential change in the composition of detrital minerals (Mens 1979). Due to the general lack of fossils, the boundary between lithostratigraphical units is conditionally taken for the upper limit of the stage which is placed at the base of the Irben Formation. Since the boundaries of the stage have not been firmly fixed, its thickness is difficult to determine, but it is mostly less than 40 m.
Fossils are extraordinarily rare. Occasionally, casts of Volborthella, undeterminable fragments of inarticulate brachiopods, valves of agglutinated foraminifers and some ichnofossils, mostly of the genus Skolithos are encountered. Acritarchs have been found from the Soela Formation in the Varbla (Fig 20-188) drill core (440.1 m) and from the Vaki Formation in the Oostriku‑700 (Fig. 20-155) drill core (256.2 m). The Soela Formation contains an abundant and diverse acritarch assemblage prevailed by leiosphaerids. The occurrence of Goniosphaeridium varium (Volk.) and Skiagia ciliosa (Volk.) among acantomorphids suggests late Early Cambrian age (Keller & Rozanov 1979a). The acritarch assemblage from the Vaki Formation is pauperated, beside leiosphaerida it contains a few Goniosphaeridium volkovae Hagenfeldt and Comasphaeridium latviense (Volk.), and this unit can have a wider biostratigraphic bracketing than the Soela Formation.
The Ljuboml’ Stage is regarded as corresponding to the Holmia inusitata Zone.
The Soela Formation as an independent stratigraphic unit was distinguished recently (Mens 1979, Kala et al. 1984a). Earlier, this part of the sequence was treated either as the upper part of the Tiskre Formation (Kala 1972, Keller & Rozanov 1979b) or the lower part of the Kurzeme (now Irben) Formation (Mens & Pirrus 1972).
The type section is in the interval of 230.7 to 263.7 m of the Emmaste drill core (Fig 20-117), Hiiumaa Island. The formation occurs on the islands of the West - Estonian Archipelago and in the western part of mainland Estonia (Fig. 21). It is known only from core sections, and its lower boundary coincides with the lower boundary of the Ljuboml’ Stage (see above). The upper boundary is lithologically clear, only east of the Koluvere ‑ Rumba ‑ Pärnu line it is somewhat debatable. The formation lies transgressively on the Lower Cambrian rocks or the crystalline basement (Fig. 20-257). The thickness varies from 7.9 to 41.8 m, increasing from north to south.
The formation consists of weakly cemented light‑coloured fine‑grained feldspatic (subarkose) sandstones, containing up to 10% of coarse sand and gravel grains. In the lowermost part of the section, mainly beyond the area of distribution of the Tiskre Formation, interbeds of greenish‑grey siltstones and argillaceous rocks occur. Pebbles of greenish‑grey clayballs and cross‑bedding marked by mica flakes and glauconite grains are characteristic of the lower part of the formation.
The Vērgale Stage, embracing the topmost Lower Cambrian in Estonia, was established in the present stratigraphical extent by Birkis et al. (1970). Previously, it had been used in a wider extent (Keller & Rozanov 1979b, Resheniye… 1986). The Vērgale‑46 drill core in the interval of 1293 to 1318 m has been selected as the stratotype section of the stage (Birkis et al. 1970, Brangulis 1989). Only the lowermost part of the stage occurs in Estonia. The Vērgale Stage is of subsurface distribution on the islands of Hiiumaa and Saaremaa and in the western part of mainland Estonia (Fig. 20). Its thickness decreases northwards being about 40 m at Seliste and less than 1 m at Tahkuna (Fig. 20-225, 34).
The lower boundary of the stage is tentatively drawn at the level of the base of the Irben Formation, near the level of the appearance of the Vērgale acritarch assemblage. It is less expressed beyond the distribution area of the Irben Formation where the whole post‑Liivi sandy Lower Cambrian succession has been distinguished as a joint Ljuboml’‑Vērgale unit or as the Aisčiai Group.
The faunal record of the Vērgale Stage includes Volborthella, agglutinated foraminifers, fragments of trilobites and inarticulate brachiopods, also ichnites are very common. The so-called Vērgale acritarch assemblage comprises Estiastra minima Volk., Skiagia ciliosa (Volk.), S. insigne (Fridr.), S. compressa (Volk.), S. orbiculare (Volk.), Comasphaeridium strigosum (Jank.), Asteridium spinosum (Volk.), A.lanatum (Volk.), A. tornatum (Volk.), Tasmanites volkovae Kirjanov, T. bobrowskae Waz., T. tenellus Volk., Dictyotidium priscum Kirjanov & Volk., Leiovalia tenera Kirjanov, Pterospermella solida Volk., Lophosphaeriduim truncatum Volk., Alliumella baltica Vanderflit, etc. with several variations in the species composition from section to section. On the ground of the palaeontological evidence, the stage corresponds to the Holmia kjerulfi Zone.
The Irben Formation (Kala et al. 1984a) was earlier termed the Kurzeme Formation (Mens & Pirrus 1972). The stratotype of the formation is the type section of the previous Kurzeme Formation represented by the interval of 1329 to 1387 m of the Pavilosta drill core in Latvia (Lieldiena & Fridrichsone 1968).
The Irben Formation is distributed on the islands of the West-Estonian Archipelago and in the western part of mainland Estonia (Figs. 19, 20), and is known only from core sections. It rests conformably on the Soela Formation having the maximum thickness (42.4 m) in the Seliste drill core (Fig. 20-225). The lower boundary of the formation is marked by the appearance of argillaceous rocks. The upper boundary of the formation is erosional throughout its distribution area in Estonia.
The formation consists of interbedded clay‑ and siltstones with interlayers of fine‑grained sandstones, the number and thickness of which increase eastwards. Brown ferruginous oolith interbeds of goethite oolites are characteristic of the formation; in Estonia, these have been found only in the westernmost sections (Figs 19, 21). Owing to the unique lithology and wide distribution of argillaceous rocks, it is an important level for regional lithostratigraphical correlation (Pirrus 1986). Argillaceous rocks are greenish‑grey, but in the upper part they are locally dark-grey with a purplish or brownish shade of colour. Besides the wavy horizontal bedding, there are abundant ichnites in clay intervals responsible for a bioturbated structure of the “kråksten” type.
Silt- and sand-size fractions contain quartz (up to 85%), feldspars (up to 20%) and micas (usually 2‑10%, rarely up to 25%). Like in the Soela Formation, the content of heavy minerals is low. Clay minerals are dominated by illite.
The Vaki Formation occurs in central and eastern Estonia (Table 6), being probably a shallow‑water equivalent of the Soela and Irben formations. It is known only from core sections (Fig. 21). The formation rests with a stratigraphic unconformity on rocks of the Dominopol’ or the Lontova Stage. In the latter case, or if the Tiskre Formation is absent, the topmost layers of the underlying units are weathered (Mens et al. 1984).
The Vaki‑67 (Fig. 21 - 148) drill core in the interval of 284.4 to 322.0 m has been selected as a type section for the formation (Resheniya… 1978). The maximum thickness of the formation (more than 35 m) has been registered in its type area (Fig. 20).
The Vaki Formation consists of weakly cemented glauconite‑bearing light-coloured fine‑grained or very fine‑grained sandstones, or of both, with thin interlayers of greenish‑grey and bleached‑purplish argillaceous rocks. On some levels the latter are highly micaceous and often contain ichnofossils of Skolithos affinites filled with fine sandy material. Besides ichnites, the formation contains scarce fragments of inarticulate brachiopods and in the Oostriku‑700 (Fig. 20-155) drill core at the depth of 256.2 m also a pauperized acritarch assemblage (further above‑ the Vērgale Stage).
Compared to the Lower Cambrian, the Middle Cambrian is of more limited distribution and known only in the subsurface occurrence (Pirrus 1991).
Due to the lack of fossils and a very low core yield, the Middle Cambrian deposits are stratigraphically and sedimentologically poorly studied. The Middle Cambrian age of rocks was proposed according to their geological setting between the palaeontologically characterized Lower and Upper Cambrian rocks, and justification has been derived from the correlations with adjacent areas where the corresponding rocks contain fossils. Under such conditions, it is not always possible to identify the Middle Cambrian boundaries with a certainty. The lower limit of the Middle Cambrian is taken at the base of a thin band formed of coarse‑grained non-glauconitic sandstones with gravel and quartz pebbles overlapping the Lower Cambrian strata or the crystalline basement. Notable is an essential change in the mineral composition of the rocks, marked by the disappearance of glauconite, increase in the degree of maturity, prevalence of allothigenous minerals in the group of heavy minerals, etc. In southearn Estonia, the lower boundary is drawn on top of the whethering crust of the Lontova Stage. The identification of the Middle Cambrian is most complicated in the sections situated on the northern margin of the Middle Cambrian distribution area where it is underlain by lithologically similar Lower Cambrian light‑coloured quartzose sandstones (Mens & Pirrus 1992).
Throughout the distribution area, the Middle Cambrian consists of siliciclastic rocks dominated by mature light‑coloured well‑sorted non-glauconitic quartzose sandstones.
The Middle Cambrian succession in Estonia is subdivided into the Ruhnu and Paala formations (Table 6). The former is a local equivalent of the Deimena Formation in western Latvia (Resheniye… 1986, Mens et al. 1984, Pirrus 1991) and, according to fossil evidence, corresponds to the Ptychagnostus praecurrens Zone (Mens et al. 1987, 1990, Hagenfeldt 1989b).
The Paala Formation is distributed mainly in the southeast of Estonia and is tentatively interpreted as an equivalent of the Sablinka Formation of the Leningrad and Pskov regions (Mens et al. 1990). On the basis of palaeontological data, the Sablinka Formation is referred to the Paradoxides paradoxissimus and P. forchhammeri zones Judging from the mineral composition, the Middle Cambrian succession in some sections, including Abja and Otepää (Fig. 22-233, 249), is not clear.
The Ruhnu Formation is distributed in the southwestern part of Estonia, while its range in the middle of southern Estonia is plausible. It lies transgressively on the Lower Cambrian and is overlain by the Upper Cambrian or Lower Ordovician rocks.
The Ruhnu drill core in the interval of 706.8 to 748 m (Fig. 22-257) has been selected as the type section for the Ruhnu Formation (Kala et al. 1984a). The maximum thickness of the formation (41.2 m) has been established in this borehole and it decreases towards the north.
The Ruhnu Formation is represented by light well‑sorted, fine‑ to very fine‑grained quartzose sandstones with only a few thin interbeds of dark‑grey, in the lower part sometimes variegated, argillaceous rocks. The basal part is often marked by a layer of very coarse‑grained sandstone containing over 10% of quartz with gravel or pebble grain-size, or with both. The top of the formation is usually cemented by carbonates and pyrite. Often an admixture of feldspars and micas (muscovite) occurs. Glauconite is lacking. Heavy mineral assemblage is characterized by the prevalence of transparent minerals dominated by zircon and tourmaline. Among opaque minerals, detrital leucoxene and ilmenite occur in almost equal quantities or the former prevails slightly. Clay minerals are dominated by illite, the content of kaolinite often reaches 30‑35%. Chlorite is rare reaching occasionally 10% in the lower part of the formation.
The Paala Formation occurs in central and southeastern Estonia transgressively on the crystalline basement or on the Lower Cambrian rocks the type section is the Viljandi drill core (Fig. 22-203) in the interval of 409.8 to 434m (Mens et al. 1984). The thickness of the formation varies greatly. Due to the low core yield, the determination of the boundaries is complicated and the thickness of the formation is unclear.
The formation consists of light quartzose non‑glauconitic middle‑ to fine‑grained sandstones with pellets of white kaolinitic clay. The grain-size of the rocks in the lower and upper parts of the formation is coarser than in its middle part where very fine- to fine‑grained varieties dominate. Feldspar and muscovite are not common. The heavy mineral assemblage is mainly composed of allothigenic minerals dominated by ilmenite. In the group of heavy transparent minerals, zircon is always the index mineral.
The character of the basal bed depends on the type of underlying rocks. On the crystalline basement, it consists of conglomeratic sandstone, locally with phosphatized pebbles (Laanemetsa, Fig. 22-269). Resting upon the argillaceous Lower Cambrian rocks, which are often weathered, the basal bed occurs like a variegated kaolinitic clay comprising coarse grains of quartz redeposited from the weathering crust.
During the last two decades, the stratigraphic extent of the Estonian Upper Cambrian and its biostratigraphical subdivision has been established more precisely due to the progress in research on acritarchs, conodonts and lingulates (Volkova et al. 1981, Kaljo et al. 1986, Popov et al. 1989, Mens et al. 1993 etc.).
The Upper Cambrian rocks have been palaeontologically documented in two isolated areas - in northern and south-eastern Estonia (Fig. 23). They crop out along the Baltic-Ladoga Klint and in the river valleys crossing it. The main localities in Estonia are Ülgase, Valkla, Turjekelder and Suurjõgi.
In Estonia, the Upper Cambrian strata are distributed sporadically and dominated by sandstones, less than 20 m in thickness. Argillaceous rocks are of limited distribution forming grey to greenish-grey interlayers within light-coloured sandstones in the lower part of the succession and brownish-grey varieties in the upper part.
The Upper Cambrian succession is condensed and interrupted by several minor and major hiatuses. The Upper Cambrian biostratigraphy and interregional correlation are usually based on trilobites. In the Estonian Upper Cambrian sections, where trilobites have not been found, it is based on acritarchs, conodonts and lingulates (Popov et al. 1989, Mens et al. 1993). Altogether five acritarch-, four conodont-, and at least three lingulate-based biostratigraphic units are distinguished (see Mens et al. 1993).
The oldest Estonian Upper Cambrian acritarch assemblage recorded from the Petseri Formation (SE Estonia) is similar to the acritarch assemblage BK1 b by Volkova (1990, Volkova & Kirjanov 1995), except the occurrence of Leiofusa stoumonensis Vang, which has been found in the base of the acritarch-bearing part of the Petseri Formation. Based on the occurrence of the genera Stelliferidium, Cymatiogalea, Leiofusa and Veryhachium, which appeared in Olenus time (Potter 1974, Downie 1984), and the absence of the Impluviculus species, the rocks comprising this acritarch assemblage are considered as a stratigraphic equivalent of the lower or middle, or of the both parts of the Olenus & Agnostus Zone.
The next acritarch assemblage in the Ülgase Formation (Table 6) is similar to that described above, but differs in the presence of Veryhachium dumontii and representatives of the genus Impluviculus and in the lack or restricted distribution of typical Middle Cambrian species. As the appearance of the genus Impliviculus has been correlated with the uppermost part of the Olenus & Agnostus Zone (Downie 1984) and with the lowermost part of the P. spinulosa Zone (Martin & Dean 1988), we regard the Ülgase Formation tentatively as a stratigraphic equivalent of the uppermost Olenus & Agnostus and the lowermost Parabolina zones (Table 6).
The third Upper Cambrian acritarch assemblage in the Estonian succession is distinguished by the appearance of Trunculumarinum revinium (Vang.) Loeblich et Tappan, Dasydiacrodium caudatum Vang. corresponds to the microflora A4, i.e. T. revinium - D. caudatum assemblage sensu Martin and Dean (1988), to the assemblage BK-3 sensu Volkova (1990), to the uppermost P. spinulosa Zone and to the Leptoplastus Zone, as a whole (Table 6).
The next acritarch assemblage, showing a high taxonomic diversity and variation in different sections, contains a significant amount of diacroids and sometimes also endemic forms. On the East-European Platform, this assemblage was first described from the upper part of the Ladoga Formation distributed in the Leningrad Region (Volkova & Golub 1985) and is referred to as BK-4. In Estonia, the assemblage occurs together with conodonts of the Proconodontus Subzone in the Tsitre Formation, but these acritarchs have also been found in the Cordylodus andresi Zone in the Kallavere Formation. A relatively similar assemblage together with the trilobites of the Peltura scarabaeoides Zone has been determined from the Degerhamn section of southern Öland (Di Milia et al. 1989).
The youngest Upper Cambrian acritarch assemblage, which also contains an abundance of diacroids (Volkova 1989, 1990), can be distinguished by the appearance of Acanthodiacrodium angustum (Downie) Combaz and Dicrodiacrodium ramusculosum (Combaz) Volkova. It occurs together with conodonts of the C. proavus Zone, and may be also characteristic of the C. intermedium Zone (Volkova & Mens 1988).
Conodonts have been studied from a number of outcrops and drill cores. The number of specimens is small, representing mostly the genera Phakelodus, Furnishina, Prooneotodus and Westergaardodina. Eoconodonts (Proconodontus, Eoconodontus and Cordylodus) have been found only in the topmost Upper Cambrian. According to the conodont zonation worked out by Sergeyeva and Viira for the Baltic-Ladoga Klint area, the Upper Cambrian sequence is subdivided as follows (from below upwards): the Westergaardodina Zone with W. bicuspidata, W. moessebergensis and Proconodontus subzones, and the Cordylodus andresi and C. proavus zones (Kaljo et al. 1986).
Lingulate brachiopods are represented in the Upper Cambrian of Estonia by lingulids and acrotretids. Following the brachiopod zonation suggested by Popov and Khazanovitch (Popov et al. 1989) and adopted by Puura and Holmer (1993), four brachiopod zones can be distinguished. From below upwards these are the Ungula inornata Zone, the Ungula convexa Zone, the Ungula ingrica Zone and the Obolus apollinis Zone. The first three zones belong to the Upper Cambrian, while the latter one can belong partly to the Ordovician, depending on the position of the lower boundary of the Ordovician System.
The Estonian Upper Cambrian includes three succeeding lithological units in the rank of formation (Petseri, Ülgase, Tsitre; Table 6), and the lower part of the overlying Kallavere Formation.
The Petseri Formation introduced by Kajak (1967) in the rank of beds is known only from core sections in southeastern Estonia, from where it spreads east- and southwards. In the Petseri borehole in Russia (Fig. 23), serving as the type section, the formation is 10.7 m thick. In complete sections the Petseri Formation can be subdivided into three parts. The lower and upper parts are represented by light-coloured weakly cemented quartzose sandstones, whereas the middle part is predominantly composed of grey argillaceous rocks (Volkova et al. 1981). Sandy parts of the formation contain some glauconite and debris of inarticulate brachiopods. In the argillaceous part, shells and fragments of lingulates of the genera Ungula and Oepikites and the Dasydiacrodium setuensis - Leiofusa stoumonensis assemblage of acritarchs occur (Volkova et al. 1981, Volkova 1990, Paalits 1992a). On the ground of this acritarch assemblage, the Petseri Formation, or at least its middle part, is considered as a time equivalent of the lowermost part of the Olenus & Agnostus Zone (Table 6).
The Ülgase Formation was referred by Öpik (1929) to the Acrotreta Sandstone and assigned to the local Acrotreta-Lingulella Zone. Subsequently, this part of the succession was defined as the Ülgase Member in the limits of the Pakerort Stage (Müürisepp 1958). In the rank of formation it was first considered by Khazanovich and Missarzhevsky (1982). The formation with a thickness of about 10 m is better fixed in the vicinity of Tallinn and within some 50 km east and south of it. The Ülgase Formation consists of light-coloured very fine- to fine-grained sandstones with interbeds and lenses of greenish-grey clay in the lower part and brownish-grey thin films in the upper part. Its upper boundary is transitional and in the earlier papers the lower part of the overlying Tsitre Formation was regarded as belonging also within this formation (Mens 1984, Popov et al. 1989). The formation contains numerous lingulates of the genera Ungula (including U. inornata), Oepikites, Angulotreta and Cerotreta. The occasional conodonts belong to the genera Phakelodus, Furnishina and Prooneotodus (Kaljo et al. 1986, Mens et al. 1993). Torellella sulcata Missarzhevsky abound. In the argillaceous interlayers the acritarchs, forming the Impluviculus multiangularis - Veryhachium dumontii assemblage are numerous (Volkova 1982, 1990, Volkova & Mens 1988). On the basis of fossil evidence, the Ülgase Formation is assigned to the uppermost part of the Olenus & Agnostus Zone and to the lower part of the Parabolina spinulosa Zone (Table 6).
The Tsitre Formation was introduced by Popov and Khazanovich (1985) with the stratotype in the Turjekelder section. Earlier, this part of the succession belonged to the Kallavere Formation (Kaljo et al. 1986, Resheniye… 1986). Currently, the Tsitre Formation (Table 6) includes also the underlying beds containing kerogen-bearing argillaceous interlayers and differing in fossil record from the Ülgase Formation (Mens et al. 1993).
The Tsitre Formation expands as a narrow belt from Tallinn to Kohtla-Järve (Fig. 23). Its thickness in the outcrop sections is a bit more than 3 m. In the drill core sections its thickness is unclear due to the low core yield, but probably it is less than 10 m.
The formation is typically represented by light-grey weakly cemented fine-grained quartzose sandstones, with a few thin interbeds of variegated, dominantly brownish-grey clayey rocks. These interlayers are often accompanied by bedding planes covered with convex-up lingulate shells. This has been considered in drawing the boundary between the Ülgase and Tsitre formations.
The co-occurrence of Trunculumarinum revinium and Dasydiacrodium caudatum in the lower part of the Tsitre Formation suggests that these deposits belong to the upper part of the Parabolina spinulosa Zone (Martin & Dean 1988, Paalits 1992b). The upper part of the Tsitre Formation contains a rather rich fossil record and its relationship with the trilobite zones is shown in Table 6.
The uppermost part of the Cambrian is considered to be represented by the lower part of the Kallavere Formation. The latter is discussed with the rest of this formation within the Ordovician.
J. Nõlvak
In the Ordovician, epicontinental seas with extensive distribution of carbonate sediments had a greater extent than in any other period. The marine flora and fauna changed markedly in the course of the Ordovician. A number of major taxonomic groups (bryozoans, brachiopods, echinoderms, trilobites, ostracodes, chitinozoans and others) appeared or became common. In this respect, the Ordovician is one of the most interesting periods in the history of marine faunas, and Estonia is among the areas in the world where this fauna is well preserved and studied.
The Ordovician was characterised by an extreme biogeographical differentiation of both planktic and benthic faunas, but with different degree. This makes the worldwide correlation of the Ordovician rocks difficult and has resulted in numerous regional stratigraphic schemes. A series of detailed stratigraphical charts compiled for the East Baltic (see Resheniya… 1978, 1981, Männil & Meidla 1994, and literature cited in these papers), gave a relatively stable detailed local classification for Ordovician rocks and afterwards obtained the status of a regional standard for most of the East-European Platform (Männil 1990).
The large-scale biogeographical and facies differentiation within the Ordovician Palaeobasin of Baltoscandia is well expressed in the concept of confacies belts (Jaanusson 1976, Fig. 24). The territory of Estonia is divided between the North Estonia and Central Baltoscandian confacies.
An emended version of the correlation chart, presented for the Estonian succession ranging from 70 to 180 m in thickness (Table 7), is based mainly on the above-cited schemes. For practical reasons the present version is simplified and many smaller subdivisions have been omitted. Some subdivisions, defined earlier as formations due to overestimations or difficulties in their specifications, are treated as members. Often the lithounits (Table 7) have diachronous (wavy line) or topical (discontinuous line) boundaries or the unit serves as “topostratigraphic” unit (sensu Jaanusson 1976, p. 310) with their boundaries coinciding with the stage boundaries.
H. Heinsalu & V. Viira
The lowermost Ordovician Pakerort Stage (Table 7) distinguished by Raymond (1916) consists of two different lithotypes - the Obolus Sandstone (the Kallavere Formation, Männil & Rõõmusoks 1984) and Dictyonema Shale (Türisalu Formation, Müürisepp 1958, 1960b). During the last decades, both lithotypes have been studied in particular detail in terms of their industrial use and potential environmental impact. In northern Estonia, the so-called Obolus-Conglomerate (brachiopod coquina) occurs usually at the base of the Obolus Sandstone. The coquina was used as a good lithological marker in fixing the lower boundary of the Pakerort Stage and the Cambrian/Ordovician boundary in Estonia.
The Cambrian/Ordovician boundary (Photo 17) became an object of special international studies in the 1970s, and since then several different stratigraphical levels have been proposed as the lower boundary of the Ordovician system. According to most of stratigraphers, the definition of the Cambrian/Ordovician boundary should be based on conodonts and the horizon chosen should be close to, but below the lowest planktic (nematophorous) graptolites. Three biostratigraphic horizons of conodonts were considered as possible guides for marking the boundary level. These are the base of Cordylodus proavus, of C. intermedius and of C. lindstromi zones. Currently, the attention is focused on the first appearance of the conodont Iapetognathus n. sp. in the lower part of the C. lindstromi Zone just above the first appearance of the planktic graptolites R. praeparabola and R. parabola in the Dayangcha section (China). In Estonia, these conodont zones have been identified in the Kallavere Formation considered preliminarily (Männil & Rõõmusoks 1984) as the oldest part of the Ordovician sequence. Somewhat later the lower boundary of the system and of the Pakerort Stage was tentatively drawn at the level of the first appearance of Cordylodus (base of C. andresi Zone, Fig. 25) in the lower part of the Kallavere Formation (Kaljo et al. 1986, Resheniya… 1987, Männil 1990). But if a higher stratigraphical level (e.g. the base of the C. lindstromi Zone) will be accepted internationally for the boundary between the Cambrian and Ordovician systems, most of the Kallavere Formation must be excluded from the Ordovician (Norford 1991, Miller & Taylor 1995, Fig. 25).
As there are no distinct lithological changes on the boundaries of the conodont zones in most of the sequences, the Obolus Sandstone of the Kallavere Formation will be treated below as an entity (Fig. 26). Within that formation the main attention focuses on the distribution of conodonts, zones of which could serve as guides for the boundary between the systems. The following succession of conodont zones (in ascending order) has been established in the Obolus Sandstone and Dictyonema Shale (Kaljo et al. 1986): Cordylodus andersi, C. proavus, C. intermedius, C. lindstromi and C. angulatus. The Cordylodus andresi Zone has been established only in a few sequences in northern Estonia (Kidaste core in Hiiumaa, outcrops at Turjekelder, Vihula and Toolse). The C. proavus Zone occurs in all sections (except Turjekelder) where conodonts have been studied, and in some sections it is rather thick (Fig. 25). The zone is absent in most of the brachiopod coquina (Obolus-Conglomerate). The morphological variability of the zonal species of the C. proavus Zone suggests that the zone is discontinuous: sometimes the lower, sometimes the upper part is missing. Compared to other conodont zones in the Estonian sequences, the C. proavus Zone has the most distinct lower boundary which coincides or occurs close to the lower boundary of the Kallavere Formation.
The C. intermedius Zone is of limited distribution in Estonia. The index species has been established only in three sections (Mäekalda in Tallinn, Ülgase, Toolse). The C. lindstromi Zone, vice versa, is widespread. It is missing only on the Pakri Cape and possibly also in the Vihula section (Fig. 25). The specimens of C. lindstromi Druce et Jones found in Estonia and Australia (Nicoll 1991), are morphologically very similar (low base, small cavity with one or more pointed secondary tips).
The C. angulatus Zone occurs in all northern Estonian sections, except the Pakri Cape. Its lower boundary is well defined by the appearance of numerous specimens of C. angulatus Pander apparatus. In the western part of Estonia, it coincides with the lower boundary of the Suurjõgi Member, which is a good lithological marker (Fig. 26).
The co-occurrence of conodonts and graptolites differs considerably from area to area (Kaljo & Viira 1989). The Rhabdinopora flabelliformis group makes its first appearance at different levels in relation to the conodont zones. In the vicinity of Tallinn, it occurs in the top of the C. proavus Zone or in the C. intermedius Zone, east of Tallinn in the C. lindstromi Zone, and to the west of it in the C. angulatus Zone. This complicates the use of graptolites in the correlation of sections. Nevertheless, the graptolites serve as the most important group of fossils in establishing the upper boundary of the Pakerort Stage which falls into the organic-rich argillites of the Türisalu Formation.
Besides the above-mentioned conodonts and graptolites, lingulate brachiopods and acritarchs can be used for the subdivision of the Cambrian-Ordovician boundary beds (Kaljo et al. 1986, Puura & Holmer 1993, Mens et al. 1993, Paalits 1995).
Kallavere Formation
The quartzose sandstone with interbeds of dark argillite of the Kallavere Formation are distributed almost all over Estonia. It is missing only in a belt running from southwestern to eastern Estonia (Fig. 27). The formation is at its thickest (more than 17 m) in central Estonia. It consists of the Maardu, Rannu, Katela, Orasoja and Suurjõgi members, replacing each other in space or in time (Heinsalu 1981, 1987, Fig. 26).
Lithologically, the Kallavere Formation is dominated by quartzose sandstone, commonly weakly cemented, with the grain-size of 0.05-0.25 mm. It contains phosphatic brachiopod valves and their fragments, forming a distinct brachiopod coquina layer at the base of the Kallavere Formation (in the Maardu and Rannu members). The thickness of the coquina is only a few cm, except the Maardu (up to 1 m) and Rakvere (up to 4-6 m, sometimes even more) areas. The Kallavere Formation comprises lingulate brachiopods of the Ungula ingrica and Obolus apollinis zones (Heinsalu et al. 1987, Mens et al. 1993). In the conglomerate bed the most common species is Ungula ingrica, accompanied by species of the genera Schmidtites, Keyserlingia and Oepikites.
Dark argillite interbeds, 0.1 mm to 15 cm in thickness, occur generally above the brachiopod coquina. In northwestern Estonia, a 10–30-cm-thick dark argillite (Dictyonema Shale) bed with very thin interbeds of sandstone lies immediately on the lower boundary of the formation. In northwestern Estonia, west of the Kunda - Rakvere line, the uppermost part of the formation is represented by the so-called skeletal detritus layer of the Suurjõgi Member which consists of cross-bedded quartzose sandstone comprising brachiopod fragments with a size of 1-3 mm.The thickness of the Suurjõgi Member is about 1 m, except the Toolse - Vihula area where it exceeds 5 m (Fig. 26).
Türisalu Formation
Up to the 1970s, the Türisalu Formation (Müürisepp 1958, 1960a, b) was considered as the upper part of the Pakerort Stage. The studies of the distribution of graptolites (Kaljo & Kivimägi 1970, 1976) allowed to divide the formation between the Pakerort and Varangu stages. The older part is characterized by the occurrence of graptolites of the Rhabdinopora flabelliformis Zone and the younger part by the graptolites of the Kiaerograptus Zone. West of the Tallinn - Rapla line (Fig. 28) only the older part of the formation is represented. It comprises the Rhabdinopora f. flabelliformis (in the lower part) and R. flabelliformis multithecata graptolite subzones (Kaljo & Kivimägi 1970, 1976), and the upper part of the C. angulatus conodont Zone. The most complete stratigraphical sequences of the Türisalu Formation occur between Tallinn and Tapa where the lower part of the formation belongs to the Pakerort and the upper part to the Varangu Stage. All the sections of the Türisalu Formation east of this area are of Varangu Age.
The Pakerort Stage is represented mainly by dark-brown horizontal laminated graptolite argillite. The lamination is caused by the different content of organic matter (intercalation of darker and lighter laminae) or by different grain-size (Heinsalu 1990a, Kivimägi & Loog 1972). In some cases, wavy or cross-bedded structures or thin (a few cm) interbeds of light, often pyritized, quartzose siltstone occur in the lower part of the formation.
There is no lithological markers for identification of the boundary between the Pakerort and Varangu stages in the limits of the Türisalu Formation and the maps of the distribution of Tremadoc rocks have been compiled by formations (Figs. 27, 28, 29). The thickness of the Türisalu Formation is up to 7 m (Fig. 28).
H. Heinsalu & V. Viira
The later Tremadoc rocks, which belong to the Varangu Stage (Männil 1990, = Ceratopyge Stage, Männil 1966, Viira et al. 1970) and have a thickness of 4-5 m extend, as a relatively narrow (20-50 km) belt in northern Estonia (Fig. 29). In the argillites, the lower boundary of the stage can be established by the appearance of graptolites of the Kiaerograptus Zone and conodonts of the Paltodus deltifer pristinus Subzone. The appearance of adelograptids marks the lower boundary of the Varangu Stage in the lithologically quite uniform Türisalu Formation. The upper part of the formation differs from the lower part, which belongs to the Pakerort Stage, by the occurrence of interbeds of very fine-grained quartzose sands from some mm up to 4-5 cm in thickness. Frequently, these interbeds abound in pyrite concretions.The Toolse area, where the Toolse Member was defined, has been studied in particular detail (Kivimägi & Loog 1972, Heinsalu 1980).
Varangu Formation
The Varangu Formation, the youngest part of the Tremadoc, is widely distributed in northwestern Estonia (Fig. 29). It is at its thickest (ca 3 m) between Haljala and Kunda in northeastern Estonia where the Varangu Formation can be subdivided into three lithologically different parts. The lower and upper parts are predominantly clayey, consisting mostly of compact claystone which comprises glauconite and pyrite, scattered or concentrated in small lenses. The middle part is rich in glauconite and very fine-grained quartz, sometimes prevailing over pelitic material. The sand is hardly pyritized. A similar three-part sequence of the Varangu Formation occurs also on the Pakri Cape in northwestern Estonia, but its thickness there is only 0.3-0.4 m.
In most of western Estonia, the Varangu Formation is characterized by the greenish-grey compact silty clay or sandy deposits with glauconite grains. In some sequences the clays of the Varangu Formation are dark in colour which makes them similar to the Dictyonema shale.
T. Meidla
The Hunneberg Stage was introduced by Tjernvik (1956) as the Hunneberg Group in Sweden, based mainly on trilobite faunas. During several decades, the stage has been recorded as the lower substage of the Latorp Stage in Estonia (after Jaanusson 1960a, Männil 1966, Männil & Meidla 1994 etc.). Following Jaanusson (1982), Mägi (1984) and Hints et al. (1994) considered this unit in the rank of stage.
During the last years, Sweden has served as the key area for biostratigraphical research of the Ontika Subseries comprising the stages from Hunneberg to Kunda. Detailed studies of the earliest post-Tremadoc sequences by Lindström (1954), Tjernvik (1956), Jaanusson (1963) and several other researchers have been supplemented by recent studies of sequences and distribution of graptolites (Lindholm 1991 a.o.) and conodonts (Löfgren 1993a, b, 1994, 1996). In the East Baltic region, the stratigraphy of the Ontika Subseries has been studied by Lamansky (1905), Öpik (1930b) and Männil (1963a, b, 1966). In Estonia modern biostratigraphy of this interval bases mainly on conodonts studied by Viira (1966, Mägi & Viira 1976, Mägi et al. 1989).
In northern Estonia, the Hunneberg Stage is represented by poorly lithified glauconitiferous terrigenous sediments: glauconitic siltstones of the Klooga Member (thickness up to 2.9 m) and glauconite silt and sand of the Joa Member (up to 1.2 m) which together form the lower, main part of the Leetse Formation (Figs. 30, 31). The content of glauconite is increasing upwards. The glauconitic siltstones of the Klooga Member are dominated by quartz with a supplement of glauconite (Mägi 1970), while the silt- and sandstones of the Joa Member consist mainly of glauconite (50–70%) and quartz (about 10–20%; Mägi 1970, 1984, 1990). The lower boundary of the Leetse Formation and the Hunneberg Stage represents a well defined lithological marker level with the glauconitic sandstones and siltstones overlying conformably, sometimes with a discontinuity surface, the dark-brown argillites of the Türisalu Formation or light-grey clays of the Varangu Formation.
The maximum thickness of the stage reaches 4 m in northwestern Estonia, but usually it is less than 2 m. As the Hunneberg and Billingen stages have not been differentiated in most sections, the thickness map (Fig. 30) shows only total thicknesses for both stages. In western Estonia, mainly on the islands of the West-Estonian Archipelago, the absence of the Hunneberg Stage has been documented from several sections.
In Sweden, the Hunneberg Stage corresponds to the Megistaspis armata and M. planilimbata trilobite zones. The base of the stage is close to that of the Paroistodus proteus conodont zone (Fig. 32, Löfgren 1993a). In the East Baltic, the lower M. armata Zone has been established only with confidence in Latvia (Männil 1966, Ulst et al. 1982). In northern Estonia, the lowermost part of the Leetse Formation has generally been assigned to the Paroistodus proteus Zone (Männik & Viira 1990). In the Mäekalda section, the thin Klooga Member at the base of the Leetse Formation is referred to the Paltodus deltifer Zone (Fig. 31, Mägi 1984).
Among macrofossils, a distinctive assemblage of lingulate brachiopods (Goryansky 1969) has been recorded. It is characterized by Thysanotos siluricus (Eichwald) and Leptembolon lingulaeformis (Mickwitz) constituting the Leptembolon-Thysanotos assemblage, widely distributed in eastern and central Europe (Popov & Holmer 1994).
The fossil evidence from most of central and southern Estonia, is too fragmentary yet for the limitation of the Hunneberg and the overlying Billingen stages. The Hunneberg age of glauconitic sandstones (up to 0.5 m) has been established in the Karula core (Männil 1966), but in most cases the detailed stratigraphy of the undivided Hunneberg–Billingen strata is unclear.
In Latvia, and in some sections in southern Estonia, close to the Estonian - Latvian border, the mudstones of the Zebre Formation, reaching a thickness of 46 m in Latvia, have been considered as equivalents of the Varangu, Hunneberg and Billingen stages (Ulst et al. 1982, see also Fig. 30). The middle part of this formation (Zirni Member) is of Hunneberg age, as it yields a zonal trilobite Megistaspis planilimbata (Angelin) (Fig. 32) and several graptolites, including Tetragraptus phyllograptoides Strandmark, together with a zonal conodont Paroistodus proteus (Lindström) (ibid.). The occurrence of the peripheral parts of the Zebre Formation can be assumed in southern Estonia.
T. Meidla
The Billingen Stage (Tjernvik 1956, Jaanusson 1982), understood here in the sense of the upper Billingen Substage of the Latorp Stage (Jaanusson 1960a, Männil 1966, 1990, Männil & Meidla 1994), consists of two distinctive parts in northern Estonia. The lower one is represented by the glauconitic calcareous sandstones and limestones of the upper part of the Leetse Formation (Mäeküla Member, equal to BIβ by Lamansky (1905), whereas the glauconitic limestones of the lowermost part of the Toila Formation (Päite Member, equal to BIγ by Jaanusson (1951), form the upper half of the Billingen Stage (Table 7, Fig. 31). In some publications, including the detailed lithological study of the Volkhov and Kunda stages by Orviku (1960a), the Päite Member is interpreted as the lowermost unit of the Volkhov Stage. This interpretation is also in use in the Leningrad Region of Russia.
The original concept of the Billingen Stage (Tjernvik 1956) was based on the evidence from the trilobite faunas, but its lower boundary can best be recognized by the distribution of conodonts. In Sweden, it is situated fairly close to the boundary of the Prioniodus elegans Zone and this level is recognizable also in several sections of northern Estonia where it nearly coincides with the lower boundary of the Mäeküla Member in the upper part of the Leetse Formation (Fig. 31).
The Mäeküla Member consists of glauconitic sandstones which are replaced upwards by calcareous sandstones and glauconitic limestones. The lower boundary of this member is lithologically fairly distinct in the klint area, and is marked by the change from poorly cemented silts and sands to well cemented sandstones. For the purposes of correlation, the occurrence of the conodonts Prioniodus elegans Pander and Oepikodus evae (Lindström) is most important (Figs. 31, 32). The thickness of the Mäeküla Member varies from 0 to 0.5 m. Sandy material with the grain-size over 0.1 mm forms up to 80% of the rock, whereas glauconite grains make up some 80% of this fraction. The quartz content varies from 10 to 40% (Mägi 1984, 1990). The highest content of glauconite has been recorded in central northern Estonia. The macrofauna of the Mäeküla Member has not been described monographically but, according to the available evidence, its main, upper part is comparable with the Megistaspides dalecarlicus Zone of Sweden (Fig. 32).
The Päite Member is represented by limestones or dolomites which dominate in the easternmost sequences, with a low content of mainly fine-grained glauconite. The greatest thickness of the member is 1.13 m, and it decreases in the northwest direction. In the Leningrad Region, the presence of several distinctive lithological marker horizons within the equivalents of the Päite Member (roughly equal to the informal Red Dikari Member) has been demonstrated by Dronov et al. 1996, part of those can be supposedly distinguished in northeastern Estonia. On the islands of Väike-Pakri (Photo 18) and Osmussaar, the member is sandy and may contain a layer of calcareous, glauconitic sandstone at the base (Orviku 1960a). The Päite Member is roughly equivalent to the Megistaspis (Paramegistaspis) estonica Zone of Sweden (Fig. 32).
In central Estonia, the presence of the Billingen Stage needs further approval. Glauconitic sandstones occuring locally in a restricted thickness in this area, like in the Äiamaa and Võhma cores (Rõõmusoks 1960, 1983), may belong to this stage but their precise age is not yet clear. Further to the south (at Tartu), the reddish-brown, occasionally glauconitic dolomites may be tentatively assigned to the Billingen Stage. The distribution of the Zebre Formation in Latvia (Ulst et al. 1982) suggests that it extends as far as the southernmost Estonia. In Latvia, the red or mottled clays of the topmost Zebre Formation (Zante Member) contain Megistaspis (Paramegistaspis) estonica (Tjernvik) and a zonal conodont Oepikodus evae (Lindström) (Gailite & Ulst 1975, Ulst et al. 1982), characteristic of the Billingen Stage in several sections of northern Estonia (Mägi 1990, Einasto et al. 1996, see Fig. 31).
Fossil evidence from southern Estonia is too fragmentary to enable the limitation of the Hunneberg and Billingen stages over the study area. Figure 30 shows only their sum thickness. The absence of the Billingen Stage in several sections of the West-Estonian Archipelago (Saaremaa, Hiiumaa) should be mentioned.
The Mäeküla Member contains the oldest Ordovician representatives of articulate brachiopods of the genera Plectella, Panderina, Prantlina and Angusticardinia (Rubel 1961) bryozoans, ostracodes (H. Aru, pers. comm) and trilobites. Frequent occurrence of trilobites (“Megistaspis”) has been recorded in some levels (Orviku 1960a). In the Leningrad Region, the Mäeküla Member yielded the material for original definition of Conodonta by Pander (1830). The yet poorly studied fossil record of the Päite Member contains conodonts, brachiopods, trilobites and ostracodes.
T. Meidla
The Volkhov Stage, corresponding roughly to the “Glauconit kalk” by Schmidt (1879, 1881), forms a lithologically distinctive unit in the sections of the North-Estonian Klint and nearby river valleys (Photo 19). The term “Volkhov” was introduced by Raymond (1916) as the “Walchow Formation” in a broader meaning (Table 1) corresponding to the lower part of the Ontika glauconitic limestones in northern Estonia. Lamansky (1905) was the first to introduce three substages in the Leningrad Region. In Estonia these are in ascending order the Saka, Vääna and Langevoja substages (Männil & Meidla 1994), conceptually largely based on trilobite zonation. The two first units are not accepted in Sweden.
In northern Estonia, the Volkhov Stage is represented by the main part of the Toila Formation, locally succeeded by the lower part of the Sillaoru Formation (Pada Member, Fig. 33) which has tentatively been assigned to the Volkhov Stage (Orviku 1960a, a.o.). In central Estonia, the Toila Formation is assumed to compose the whole Volkhov sequence. Southward the formation grades into the Kriukai Formation (Table 7, Fig. 34).
The Toila Formation is a complicated stratum made up of various, partly dolomitized glauconitic limestones resting on sandstones of the Leetse Formation. The lower boundary of the Volkhov Stage and the Saka Member is marked by a smooth discontinuity surface with “amphore-like borings” (Orviku 1960b) in the uppermost bed of the Päite Member (Orviku 1960a), known as Püstakkiht (Orviku 1961). In northern Estonia, the main part of the Toila Formation, corresponding to the Volkhov Stage, is subdivided into five members (Orviku 1960a) which are partly lateral equivalents. Only the lower, Saka Member consisting of dolomitized glauconitic limestone (up to 1.2 m) forms the base of the Volkhov Stage all over northern Estonia (Fig. 33). It is overlain by two laterally equivalent units: the Telinõmme Member in the west (interbedded greenish-grey limestones and marls, up to 2 m) and the Künnapõhja Member in the east (mottled dolomitic limestone, up to 1 m) (Fig. 33). East of Tallinn, the upper part of the formation is represented by the Kalvi Member (grey argillaceous glauconitic limestones) with a thickness of up to 1.7 m. West of Tallinn, it is composed of the Lahepera Member (glauconitic limestones, partly sandy or conglomeratic, up to 0.5 m) which is assumed to represent the youngest part of the formation, as it locally overlies the Kalvi Member.
According to the trilobite evidence, the lower, Saka Member comprising trilobites Megistaspis “elongata” (Schmidt) and Megistaspis “lata” (Törnquist), and a zonal conodont Baltoniodus navis (Lindström) (as referred to by Männil 1966 and Männil & Meidla 1994), represents the lower, Saka Substage in Estonian succession. The Telinõmme and Künnapõhja members comprising Paroistodus originalis (Sergeeva) may correspond to the middle, Vääna Substage, while the Kalvi and Lahepera members probably correspond to the upper, Langevoja Substage (Fig. 32). Ostracodes (Fig. 35) and brachiopodes (Rubel 1961) are common in the Toila Formation. Gastropods, cephalopods and cystoids have also been recorded.
The up-to-3.5-m-thick Toila Formation is poorly developed in northwestern Estonia (Fig. 34). In northeastern Estonia where the formation is at its thickest, the rocks have undergone extensive dolomitization.
The Sillaoru Formation, the “untere Linsenschicht” by Schmidt (1897), consists of two distinct members of oolitic limestones with noticeably different ages. The lower, Pada Member (up to 0.5 m of limestone with small ferriferous ooids and occasional glauconite grains) comprises Metaptychopyge truncata (Nieszkowski) (Resheniya… 1978), Ptychopyge angustifrons (Angelin) (Mägi 1990) and is apparently of Late Volkhov (Early Kunda?) Age. Its age relationship with the Lahepera Member of the Toila Formation remains still open due to the different distribution areas of these members (Fig. 31).
In central Estonia, the sequence of the Toila Formation comprises the same members with similar thicknesses as in northeastern Estonia (Saka, Künnapõhja and Kalvi). In southern Estonia, the formation grades into the Kriukai Formation (Table 7, Fig. 34) which consists mainly of reddish-brown marls, with limestone and mudstone intercalations. The thickness of these rocks does not exceed 20 m in Estonia, but in Latvia it is much greater, reaching 32.5 m. The age of the formation in southern Estonia has not been established biostratigraphically. However, in Latvia it displays a rich fauna of trilobites, ostracodes, conodonts, more rarely articulate brachiopods and lingulates (Megistaspis “limbata” (Boeck), Ptychopyge angustifrons Angelin, Tallinnellina primaria (Öpik), Rigidella mitis (Öpik) (Gailite & Ulst 1975, Ulst et al. 1982, Männil & Meidla 1994), suggesting the Volkhov Age (Sarv 1959, Meidla & Sarv 1990, Männil & Meidla 1994).
T. Meidla
The Kunda Stage (Kunda Formation by Raymond 1916) is represented by oolitic, glauconitic (Lamansky 1905) and sandy limestones corresponding to the emended Vaginatum Limestone by Schmidt (1897). A three-part subdivision of the strata, based on the trilobite zonation, was introduced already by Lamansky (1905) in the Leningrad Region. He also assumed the absence of the lower unit - the Asaphus expansus Zone in northern Estonia which was afterwards confirmed by several authors (Raymond 1916, Orviku 1960a, Männil 1966, a.o.). Orviku (1958b) proposed to name the Lamansky’s subdivisions in ascending order the Hunderum, Valaste and Aluoja substages, and presented a detailed lithostratigraphical description of the corresponding interval in northern Estonia (Orviku 1958b, 1960a, b; Fig. 33).
In northern Estonia, the Kunda Stage comprises the Valaste (corresponding to the Asaphus “raniceps” Zone) and the Aluoja (zones of Megistaspis obtusicauda and Megistaspis gigas) substages (Fig. 32, Table 7). In most of northern Estonia, the lower boundary of the Valaste Substage is drawn at the base of the oolitic limestone of the Sillaoru Formation, or locally within the unit, being marked by a discontinuity surface on the boundary between the Pada and Voka members, and the disappearance of the glauconite grains, characteristic of the underlying strata (Fig. 33).
In northwestern Estonia, the Kunda Stage is represented by the Pakri Formation, eastwards replaced by the upper part of the Sillaoru Formation and the Loobu Formation (Fig. 33). In northeastern Estonia, the Napa Formation forms the topmost part of the Kunda Stage and grades into the Rokiškis Formation in central Estonia. In southern Estonia, the entire Kundan sequence, including the Hunderum Substage, is represented by the Šakyna and Baldone formations (Fig. 36).
The Pakri Formation (Öpik 1927), up to 4.5 m of yellowish-grey sandy limestones and calcareous sandstones, sometimes with conglomerate beds, occurs in northwestern Estonia in the area west of Tallinn. In the westernmost part of its distribution area, the main, lower part of the Pakri Formation consists of up-to-4m-thick nodular kerogenous calcareous sandstone of the Suurupi Member, overlain by the thin (0.5 m) sandy limestone of the Osmussaar Member. In the surroundings of Tallinn, the formation is represented by limestones with quartz and glauconite grains, locally with a basal conglomerate (Kallaste and Jägala members, up to 0.8 m). On the Island of Osmussaar and, to a lesser extent, in the neighbouring mainland areas, a system of sedimentary dikes cuts through the Pakri Formation and the underlying Volkhov-Billingen strata (Puura & Tuuling 1988). The time of the formation of the dikes is dated as middle-late Kunda.
In most of northern Estonia (except the distribution area of the Pakri Formation) and in central Estonia, the basal part of the Kunda Stage consists of the oolitic limestone of the Sillaoru Formation (Resheniya… 1978, Männil & Rõõmusoks 1984). The main, Valaste time part of the formation (Voka Member, up to 0.6 m) consists of clayey limestones with abundant ferriferous ooids, developed around skeletal particles or glauconite grains (Mägi 1984). Among the skeletal particles, fragments of trilobites and ostracodes dominate (50-70%, Mägi 1984). The Voka Member generally serves as a good marker level in the North-Estonian sequence, although in restricted areas of northern and northeastern Estonia it overlies the thin oolitic Pada Member which differs from the main part of the formation by the presence of glauconite grains and has been included in the Volkhov Stage by Orviku (1960a) and subsequent authors. In northern and northeastern Estonia, the Loobu Formation constitutes the main part of the Kunda Stage. Detailed study of the formation (Orviku 1958b, 1960a) has revealed its two-part subdivision; both the lower and the upper parts consist of two laterally equivalent units. In central northern Estonia, east of Tallinn, the formation is represented by clayey limestone of the Nõmmeveski Member (up to 2 m) and glauconitic limestone of the Ubari Member (up to 2 m, Fig. 33). In northeastern Estonia, the lower part of the formation consists of glauconitic limestone of the Utria Member (up to 3 m), overlain by clayey limestones of the Valgejõgi Member (up to 4.7 m, Männil 1987). In the outcrop area, large nautiloids Cyclendoceras vaginatum (Schlotheim), Estonioceras ariense (Schmidt), Para-cyclendoceras cancellatum Eichwald etc., (Rõõmusoks 1960) are characteristic of most of the Loobu Formation. In northeastern Estonia, the rocks have undergone extensive dolomitization resulting in a mottled red colour and cavernous structure. The Loobu Formation reaches its maximum thickness (7 m) in the central part of northern Estonia (Fig. 36), in central Estonia it is less than 3 m thick (Resheniya… 1987). In that area the formation consists of grey, partly clayey glauconitic limestones, overlying the oolitic limestones and marls of the Sillaoru Formation (0.5 m).
The Napa Formation, an oolitic marl and limestone body (up to 4 m), is supposed to replace the upper part of the Loobu Formation in northeastern and central Estonia (Fig. 33).
The relation of the formations and members forming the Kunda Stage in northern Estonia is well demonstrated by Orviku (1960a). The correlation is largely based on the trilobite evidence. Asaphus “raniceps” Dalman has been identified in the Suurupi Member of the Pakri Formation and in the lower part of the Loobu Formation (Nõmmeveski Member). The Osmussaar Member comprises Pseudoasaphus globifrons (Eichwald), which is known from the upper part of the Loobu Formation (Ubari and Valgejõgi members). The Napa Formation is characterized by Megistaspis gigas Angelin (Resheniya… 1978, Mägi 1990). In terms of conodont zonation, the Valaste and Aluoja substages roughly correspond to the Eoplacognathus variabilis Zone (Fig. 32). The shelly fauna is represented by brachiopods, ostracodes, gastropods and cephalopods (Öpik 1927, Sarv 1959, Rubel 1961, Mägi 1990).
In central Estonia, the Napa Formation grades into the Rokiškis Formation (Fig. 36), which is represented by red mottled oolithic limestone (up to 15 m). The fauna of this unit is poorly known in Estonia. Based on Panderodus cf. sulcatus (Fåraeus) and Pinnatulites procera (Kummerow) recorded by Männil (Resheniya… 1987), the Kunda-Aseri age has been suggested. In southern Estonia, the sequence of the Loobu and Rokiškis formations grades into the Šakyna and Baldone formations, represented by grey glauconitic limestone and clayey red limestone, respectively. Palaeontologically, these units are poorly characterized in Estonia and the age relationship to the northern and central Estonian sequences is obscure. The fauna of the Šakyna Formation in Latvia contains trilobites, more rarely brachiopods and graptolites, the Baldone Formation is more fossiliferous (Gailite & Ulst 1975, Ulst et al. 1982). In this area the succession of the above-named formations comprises the entire Kunda Stage and stratigraphically the section of southern Estonia is the completest in this interval.
The thickness of the Kunda Stage demonstrates an obvious decreasing trend towards northwestern Estonia. In most of northern and central Estonia, it does not exceed 10 m, but in southeastern Estonia may locally reach 20 m (Fig. 36).
L. Hints
The term Aseri Stage was used first by Bekker (1922, 1923) for the Schmidt´s (1897) Upper Oolitic Limestone (Obere Linsenschicht). In nowadays understanding the Aseri Stage bases in a great deal on the studies carried out by Orviku (Jaansoon-Orviku 1927, Orviku 1929, 1930a, 1940), Rõõmusoks (1960, 1970) and Männil (1966).
In northern and central Estonia, the Aseri Stage is 0.1 - 5 m thick (Fig. 37) and consists of bioclastic limestones with unevenly distributed ooids, predominantly brown ferriferous (goethitic) ooids (Orviku 1940, 1960b). In places, the ooids are frequent in the lower and upper parts of the stage, but in the dolomitic limestones of northeastern Estonia they occur only in the upper part (Fig. 38). White phosphatic ooids are distributed mainly in the westernmost sequences. These, early Middle Ordovician oolitic limestones have been treated as the Kandle Formation (sensu stricto; Männil & Rõõmusoks 1984). Afterwards, Männil (1990, Männil & Meidla 1994) proposed the name Aseri Formation for the oolitic limestones of Aseri Age. Here preference is given to the term Kandle Formation, because in many cases the upper boundary of the Aseri Stage is difficult to determine; it may coincide with the upper boundary of the oolitic limestones or fall into the upper part of it.
The Kandle Formation is subdivided into the Malla (Männil & Rõõmusoks 1984) and Ojaküla (Orviku in Aaloe et al. 1958) members. The lower, Malla Member (Asaphus and Echinosphaerites limestones by Jaansoon-Orviku 1927) is lithologically the most variable part of the Kandle Formation and differs from the predominantly oolitic limestones of the Ojaküla Member (Cephalopod Limestone) in the occurrence of glauconite, e.g. in the surroundings of Jägala, or in the absence of ooids in some places or parts of the sequence. The thickness of the Malla Member decreases from 2.5 m in the eastern to 0.30 m in the central part of the klint area. West of Jägala, the Malla Member is missing and the Aseri Stage is represented by the 10–20-cm-thick sandy oolitic (mainly with phosphatic ooids) limestones of the upper, Ojaküla Member.
The Kandle Formation extends to central Estonia (Fig. 38) with the dominantly grey-coloured limestones of northern sections turning southwards brownish-grey or yellowish-grey. In southern Estonia, the stage is represented by the up-to-9m-thick reddish-brown limestones of the Segerstad Formation (Männil 1966, Männil & Meidla 1994). In the transitional area between the Kandle and Segerstad formations, reddish-brown and mottled limestones with occasional goethitic ooids (Männil 1990, Männil & Meidla 1994) have been distinguished. They belong presumably to the upper part of the Rokiškis Formation (Laškov et al. 1984). In practice, identification of the latter unit in sections seems in a great deal subjective, and its distribution area is difficult to determine.
The lower boundary of the Aseri Stage in recent use was defined by Orviku (Jaansoon-Orviku 1927, Orviku 1929). In contrast to Bekker (1922), he excluded from that stage the lowermost part of the oolitic limestones which comprises several early Ordovician (Oelandian) taxa (Ahtiella baltica Öpik, Antigonambonites sp., Megistaspis sp., Rõõmusoks 1970 p. 30). The boundary is marked by essential changes in the faunal composition, especially in trilobites and cephalopods (Rõõmusoks 1970, table 3, see also Jaanusson 1960a, 1963). Notable is the disappearance of the trilobite genus Megistaspis and appearance of Asaphus (Neoasaphus), represented at least by six species in northern Estonia (Rõõmusoks 1970; table 3). Asaphus platyurus, a characteristic species in the Segerstad Limestone in Sweden and Latvia (Jaanusson 1960a, Männil 1963b, 1966) occurs also in southern Estonia (Karula core, Männil 1966, fig. 12). Of new faunal elements, Echinosphaerites as a quite easily notable fossil is also worth of mentioning (Jaansoon-Orviku 1927, p. 15, 16; Orviku 1929, p. 9-11).
The data published on the distribution of ostracodes in the Aseri Stage in Estonia is scanty (Sarv 1959, Männil 1966). The ostracode Pinnatulites procera Zone of the Kunda Stage is replaced by the Piretella tridactyla Zone in the Aseri Stage (Meidla & Sarv 1990). In several core sections (Männil 1966, figs. 12-14), Euprimites effusus Jaanusson appears close to the lower boundary of the Aseri Stage.
The chitinozoans are known only in the grey-coloured rocks of the Kandle Formation which comprises the Cyatochitina regnelly and C. striata zones (Table 7). The boundary between these zones coincides with the boundary between the Malla and Ojaküla members.
The Aseri Stage corresponds to the lower part of the Didymograptus murchisoni graptolite zone and roughly to the Eoplacognathus suecicus conodont zone (Männil 1990, Männik & Viira 1990, Einasto et al. 1996).
L. Hints
In northern Estonia, the fairly uniform Early Viru sequence of comparatively thick-bedded, hard bioclastic limestones abounding in discontinuity surfaces (Saadre 1992, 1993), was first distinguished as a separate unit - the Building Limestone (Baukalkstein), by Orviku (Jaansoon-Orviku 1927). Subsequently, this unit, determined mainly by the lithological criteria, was termed (Orviku 1940) the Lasnamägi Stage after the sections in the Lasnamägi quarry in the northeastern part of Tallinn. The Lasnamägi Stage is well-exposed also in some other sections, including Suhkrumägi (Photo 20) and Mäekalda (see Einasto et al. 1996, fig. A16) in the vicinity of the type section. In general lines, Orviku’s interpretation of the Lasnamägi Stage kept valid until the 1970s (Jaanusson 1945, Rõõmusoks 1960, 1970, Männil 1963a). In 1966, Männil (1966) stated that the Building Limestone comprises two distinct successive faunal associations of which the upper one with several characteristic trilobites, such as Xenasaphus devexus devexus (Eichwald), Asaphus (Neoasaphus) lepidus Törnquist, and graptolites including Gymnograptus linnarssoni (Moberg), is closely related to the fauna of the overlying argillaceous limestones of the Uhaku Stage. The beds containing the “upper” fauna were included (Männil 1966, 1976, Resheniya… 1978) to the Uhaku Stage, while the term Lasnamägi Stage was restricted to the lower half of the Väo Formation (Photo 21) in recent use (Männil & Rõõmusoks 1984), corresponding roughly to the Kallaste Substage by Rõõmusoks (1970), and to the lower part of the former Building Limestone.
The 4–10-m-thick Väo Formation (Fig. 39) is subdivided into three units; in ascending order these are the Rebala (the relatively argillaceous part, thickness up to 3 m), Pae (dolomites, up to 1.5 m) and Kostivere ( hard limestones, up to 6 m) members. Besides, in the stratotype area where the formation has a detailed bed-by-bed stratification, each layer has a name of its own given by quarry-workers (Mägi 1990, Einasto et al. 1996). Männil (1976) included to the Lasnamägi Stage the Rebala and Pae members and the lower part of the Kostivere Member, up to the discontinuity surface, above which Gymnograptus linnarssoni appears.
In northern Estonia, the lower, Lasnamägi part of the Väo Formation is up to 4.5 m and in the stratotype area at Lasnamägi up to 4 m thick (Männil 1976, Männil & Saadre 1987, Mägi 1990). Still in many sections the exact level of the upper boundary of the Lasnamägi Stage is not established and in Fig. 39 the total thickness of the Väo Formation is given. In southern Estonia, the Lasnamägi Stage is represented by red (lower part) to grey (upper part) bedded, mostly micritic limestones of the Stirna Formation (Ulst & Gailite 1976), equivalent to the Seby and Folkeslunda limestones of Öland Island and mainland Sweden (Männil & Meidla 1994). The Stirna Formation is up to 15 m thick (Fig. 39) which corresponds to the maximum thickness of the formation in Estonia and northwestern Latvia (Ulst et al. 1982, fig. 45). In the transitional area between the Väo and Stirnas formations in central Estonia, the oolitic lithofacies is developed (Põlma 1982, fig. 7).
In northern Estonia, the lower boundary of the Lasnamägi Stage falls into the upper part of the oolitic limestones, predominantly with goethitic ooids, of the Kandle Formation. It is marked by the discontinuity surface above which there appear brachiopods (Equirostra, Noetlingia), trilobites (Illaenus schroeteri (Schlotheim), Illaenus schmidti Nieszkowski, cephalopods (Lituites sp.) and others (Jaanusson 1945, Rõõmusoks 1970). The level of the appearance of phosphatic ooids in the top of oolitic limestones is used as the lower boundary of the Lasnamägi Stage, if the boundary discontinuity surface is absent or the palaeontological data are insufficient.
The lists of fossils published earlier for the Lasnamägi Stage (Rõõmusoks 1970, table 4) can be used with consideration that only the data from the Kallaste Substage by Rõõmusoks (1970) characterize the Lasnamägi Stage in recent meaning. In northern Estonia, the macrofauna is quantitatively dominated by sedentary forms, particularly articulate brachiopods and bryozoans (Jaanusson 1984). Cephalopods occur mostly in lower quantities, except northeastern Estonia where the lowermost beds of the stage abound in orthocones and where lituitids are also fairly common. The same groups of fossils seem to be relatively abundant in core sections as well (Hints & Põlma 1981).
Important information on the range of North Atlantic conodont zones (see Bergström, 1971 for the reference), graptolites and chitinozoans of the Lasnamägi Stage in northern Estonia is provided by Männil (Männil 1976, fig. 2; 1986, fig. 2.1.1). According to him, the stage is comparable to the Eoplacognathus foliaceus Subzone and the main lower part of E. reclinatus Subzone. Although the chitinozoans are not very dignostic for the distinction of the Lasnamägi Stage, the Cyathochitina sebyensis Zone is a good marker for the Aseri-Lasnamägi boundary beds (Table 7).
L. Hints
The Uhaku Stage comprises, in the revised and amended form (Jaanusson 1960a, Männil 1966, 1976, 1990), the Caryocystites Zone (Jaansoon-Orviku 1927, = Uhaku Stage by Orviku 1940) and the upper part of the Building Limestone (Väo Formation). These two parts of the Uhaku Stage are considered also as substages (Männil 1976, 1990).
The thickness of the Uhaku Stage varies from 5-10 m in western to about 20-25 m in eastern Estonia (Fig. 39). In northern Estonia, the hard bioclastic limestones of the Väo Formation, forming the lower part of the Uhaku Stage, are of a rather stable thickness (4-5 m). The upper part of the Uhaku Stage, made up of relatively thin-bedded argillaceous limestones of the Kõrgekallas Formation, is subdivided (Table 7) into the Koljala, Pärtlioru and Erra members (Männil & Rõõmusoks 1984). The thickness of the formation decreases from about 18 m in northeastern to 1-2 m in northwestern Estonia (Figs. 39, 40). The lower boundary of the Kõrgekallas Formation and the Koljala Member, formed of argillaceous limestones with marly intercalations, supposedly coincides with the lower boundary of the Conochitina tuberculata Zone (Männil & Bauert 1986, p. 17; Table 7). In the Pärtlioru and Erra members, the argillaceous intercalations in the bioclastic limestones are partly kerogeneous. In the Oil Shale Basin in northeastern Estonia (Puura 1986), thin kukersite beds (up to 2 cm) occur, or they form together with limestones and marls distinct intervals (up to 1.6 m) between relatively pure limestones (Männil & Bauert 1986). These are the oldest kukersite beds in the Middle Ordovician sequence in northern Estonia. In central Estonia, the kukersite beds appear in the Kukruse Stage (cf. Männil 1966, 1986).
In southern Estonia and also in Latvia, the Uhaku Stage is represented mainly by micritic limestones with intercalation of bioclastic limestones and marls of the Taurupe Formation (= Furudal Formation in Männil 1966) with a thickness of 6 - 19 m. Only in a few sequences in western Latvia, the Taurupe Formation is over 20 m thick (Ulst et al. 1982, fig. 46). The limestones of the transitional belt between northern and southern Estonia are characterised by an interfingering pattern which resembles that of northern Öland in Sweden (Männil 1966). Both goethitic and phosphatic ooids occur in some places in the basal part of the Uhaku Stage testifying to continuous shift of the oolitic lithofacies in time (Põlma 1982, fig. 7, Pärnu and Ikla cores, Fig. 40).
In northern Estonia, the lower boundary of the Uhaku Stage falls into the lithologically rather uniform Väo Formation. In practice, a prominent discontinuity surface is used as a boundary marker above which several new taxa appear, some of them widespread and frequent throughout the Baltic Basin. Mass occurrence of the trilobite Xenasaphus d. devexus (Eichwald) is recorded from the Island of Osmussaar in the west as far as Ingria (L. Popov and R. Einasto, pers. comm., Alichova 1960, 1969) in the east. Of the graptolites of the Hustedograptus teretiusculus Zone, Gymnograptus linnarssoni (Moberg) is identified from the Oslo Region up to the Moscow Syneclise (Männil 1976). According to Männil (1986), the lower part of the Uhaku Stage corresponds to the Eoplacognathus robustus and E. lindstroemi subzones of the Pygodus serra Zone (Table 7). The upper part of the stage corresponds to the Pygodus anserinus Zone. The latter zonal species appears close to the base of the Kõrgekallas Formation. Nevertheless, on the basis of the distribution of conodonts, the boundary between the Lasnamägi and Uhaku stages is unclear, at least on the subzones level (Table 7). The chitinozoan Conochitina clavaherculi Subzone comprises the most part of the Väo Formation, including the strata with the first finds of G. linnarssoni (Männil 1986, fig. 2.1.1).
The Uhaku Stage comprises a varied sedentary benthic fauna, particularly articulate brachiopods, bryozoans and cystoids. Since there is no generally acknowledged interpretation of the Uhaku Stage, the lists of fossils presented by different researchers comprise taxa from different stratigraphical intervals (Rõõmusoks 1960, 1970; Männil 1963a, 1966).
Macrofossils are poorly known in the subsurface area of the Uhaku Stage (Fig. 40). The Taurupe Formation, which is distributed in southern Estonia, includes many elements, such as Nileus and Upplandiops (=Estoniops sp. n. in Männil 1966, fig. 12) among trilobites, and both Alwynella? and Christiania among articulated brachiopods, which are widely distributed in the Furudal limestones in Sweden (Jaanusson 1960a, 1963, Jaanusson & Ramsköld 1993).
L. Hints
The Kukruse Stage (Kuckerssche Schicht by Schmidt 1879, 1881) as a stratigraphical unit comprises the commercially exploited oil shale (kukersite) seams (Chapter X) and the richest and most diverse faunal assemblage in the Ordovician of Estonia represented by more than 330 species and subspecies (Rõõmusoks 1970, table 10).
The stratigraphy of the Kukruse Stage has been dealt with in several papers (Rõõmusoks 1957, Männil 1984, Männil & Bauert 1984, 1986, Bauert 1993, Saadre & Suuroja 1993b). The bed-by-bed stratification of the kukersite complex with special sets of indices for the individual kukersite seams form the base for the correlation of the sequences within the kukersite basin (Männil 1984, fig. 2; Bauert & Puura 1990).
The thickness of the Kukruse Stage (Fig. 41) ranges from about 3 m in western to more than 20 m in eastern Estonia (Saadre & Suuroja 1993a). The stage consists of three lithologically distinct formations. The argillaceous bioclastic limestones with intercalations of kukersite (oil shale) and marls of the Viivikonna Formation (Männil & Rõõmusoks 1984) are distributed northeast of the line Osmussaar Island (southwestern Estonia) - Mehikoorma (south coast of Lake Peipsi) (Fig. 41). Based on the frequency of kukersite seams or the content of the kerogenous component, the Viivikonna Formation is subdivided into the Kiviõli, Peetri and Maidla members (Fig. 42). The Kiviõli (lower) and Peetri (upper) members differ from the Maidla (middle) Member by the occurrence of 10—14-cm-thick kukersite seams, while the middle part of the formation consists of kerogenous and variously argillaceous limestones (Männil et al. 1986, Bauert 1993, figs 3, 4). Due to the facies shift of the kukersite beds (Männil et al. 1986), the boundaries of the Viivikonna Formation are diachronous. As a result, the upper part of the Viivikonna Formation (Peetri Member) is missing in northeastern Estonia, but it is exposed in the vicinity of Tallinn (Fig. 42, Nõlvak & Hints 1996) and is well-known by core sections south of the outcrop area (Männil 1984, Männil & Bauert 1984, Männil & Saadre 1987).
Westwards, the Viivikonna Formation grades into the bioclastic limestones of the Pihla Formation with a thickness of about 3 - 6 m (Saadre & Suuroja 1993b) and southwards into the limestones with dark pyritized skeletal detritus and nodular intercalations of argillaceous marls of the Dreimani Formation (Fig. 41, Springis 1974). The thickness of the latter varies from 7 to 14 m and only in southeastern Estonia it is about 20 m, which is nearly the same as in eastern Latvia (Ulst et al. 1982, fig. 47).
For the lower boundary of the Kukruse Stage, Bekker (1923, 1924b) proposed the base of the lowermost commercially important kukersite seam “A” at the base of the Viivikonna Formation. The renovation of faunal association begins with the appearance of new bryozoans in seam “A”. Somewhat higher, in seam “C” several new species, including the brachiopods Bilobia musca (Öpik), Sowerbyella (S.) liliifera (Öpik), Estonomena estonensis (Bekker), and the trilobites Asaphus (Neoasaphus) nieszkowskii Schmidt, Estoniops exilis (Eichwald), Paraceraurus aculeatus (Eichwald) appear (Rõõmusoks 1970, table 9). In western and southern Estonia, the base of the Pihla or Dreimani Formation is used as the lower boundary of the Kukruse Stage. This level is marked by the appearance of indicator ostracodes Baltonotella kuckersiana (Bonnema), Conchoprimitia leperditioides Thorslund, Euprimites locknensis Thorslund and others, several of which are common with the lower part of the Dalby Limestone in Sweden (Männil 1966, Jaanusson 1976). At the same time, several early Viru taxa, such as Chasmops odini odini Eichwald, Sowerbyella (Viruella) uhakuana (Rõõmusoks), Platystrophia biforata (Schlotheim), Dianulites fastigiatus (Eichwald) and others, disappear close to the lower boundary of the Kukruse Stage (Rõõmusoks 1970, p.156, 157). The graptolite Orthograptus uplandicus whose range zone corresponds to the Kukruse Stage (Männil 1984) and the chitinozoa Cyathochitina savalaensis appear roughly on the lower boundary of the Kukruse Stage. In all likelihood, also the boundary between the North Atlantic conodont anserinus and tvaerensis zones (Männil & Bauert 1986) falls into the lower part of the Kukruse Stage.
The diverse assemblage of Kukruse macrofossils is represented first of all by bryozoans (more than 60 species), brachiopods (about 90 species) and trilobites (about 50 species: Rõõmusoks 1970, table 10) which form about two thirds of the species identified. The most abundant and diverse association occurs in the Kiviõli Member in the lower part of the stage. Still some species, such as Hesperorthis inostrantzefi inostrantzefi (Wysogorski), Echinosphaerites aurantium suprum Hecker, are notable due to their mass occurrence in the upper part of the stage (Rõõmusoks 1970, p. 169). The character of the distribution of some brachiopods and trilobites, such as Estlandia marginata magna Öpik, Otarion planifrons (Eichwald), Pharostoma nieszkowskii (Schmidt) and others, shows a facies shift from the lower part of the Kukruse Stage (Kiviõli Member) in northeastern to the upper part (Peetri Member) in northwestern Estonia.
In the core sections, macrofossils are of secondary importance due to their scarcity, especially in western Estonia (Fig. 40). Still, the occurrence of some species should be noticed. In some northernmost core sections, the brachiopod Kullervo panderi (Öpik) marks the lowermost part of the Kukruse Stage (Rõõmusoks 1970). In the outcrops, this species appears presumably in the kukersite seam “G”, which lies 1-4 m above the lower boundary of the stage. In the southern periphery of the Viivikonna Formation and in the Dreimani Formation, Asaphus (Neoasaphus) ludibundus Törnquist and Bilobia musca (Öpik) appear in the Kukruse Stage and in some areas Echinosphaerites becomes frequent.
L. Hints
Jaanusson (1995) proposed the term Haljala Stage for the unit which comprises the Idavere and Jõhvi chronostratigraphical subdivisions, previously regarded as separate stages. These two subdivisions, now classified as the Idavere and Jõhvi substages (Table 7), comprise most of K‑bentonite beds which lie below the thickest bed (“d” by Jürgenson 1958a) established in eastern Baltic. The substages are difficult to differentiate in southern Estonia, in areas where K-bentonite beds are uncertain or absent. Also the faunal distinction between the substages is rather inconsiderable (Põlma et al. 1988, figs. 9-11). In Estonia, the thickness of the Haljala Stage varies mostly from 10 to 20 m (Fig. 43).
The Idavere Substage comprises the regularly bedded hard bioclastic limestones of the lower, Tatruse Formation and argillaceous limestones with intercalations of marls and some thin K-bentonites of the upper, Vasavere Formation (emended by Männil & Meidla 1994). This substage has the most reduced sequence in northern Estonia and in some places in the surroundings of Tallinn it is entirely missing (Jaanusson 1945). The Tatruse Formation (Põlma et al. 1988) corresponds roughly to Schmidt’s (1881) original concept of “Itfersche Schicht” and the fauna recorded by him and his contemporaries from the “Itfer” belongs only to this formation. The Vasavere Formation contains usually two, but in the west sometimes up to 18 K-bentonite beds, which belong to the Grefsen complex of bentonites (Bergström et al. 1995). In the areas where only beds “a” and “b” (Jürgenson 1958a) are recognizable, the upper bed is regarded as the top of both the Idavere Substage (Männil 1963a) and the Vasavere Formation. In areas farther south where bentonite beds of the Grefsen complex disappear or in the west where they are numerous, a distinction between the Idavere and Jõhvi substages is difficult.
The Jõhvi Substage, which is at its thickest (more than 10 m, Fig. 44) in northwestern Estonia, comprises argillaceous bedded to nodular limestones with argillaceous intercalations in the middle part (Männil & Rõõmusoks 1984, Põlma et al. 1988). These limestones form the lower part of the Kahula Formation (Männil & Meidla 1994). A fairly persistent K‑bentonite bed (bed “c” by Jürgenson 1958a, “Sinsen K-bentonites” by Bergström et al. 1995) occurs close to the boundary between the middle and upper parts of the Jõhvi Substage.
In southern Estonia, the Haljala Stage, 8 - 18 m in thickness, is represented by argillaceous limestones with thin K-bentonite beds and in places with phosphatic ooids of the Adze Formation (Ulst et al. 1970).
In the outcrop area, the lower boundary of the Haljala Stage and the Tatruse Formation is formed by a conspicuous discontinuity surface – a hardground which in places is penetrated by cavities, some 5 cm or even more in diameter at the surface and extending sometimes some 40 cm downwards (Põlma et al. 1988). The limestones above the basal discontinuity (Kisuvere Member) comprise up to 16% of quartz sand. The most detailed stratification of the lowermost beds of the Haljala Stage is based on chitinozoans. The oldest part of the stage, the Armoricochitina granulifera and Angochitina curvata zones (Männil 1986, fig. 5.1.1, Nõlvak & Grahn 1993) occurs in the Laeva area in eastern central Estonia (Fig. 45) where the stage has the maximum thickness (ca. 25 m, Fig. 44, core No. 285). These two zones do not occur in the northenmost sequences, where the Lagenochitina dalbyensis Zone forms the basal part of the substage.
The gap on the boundary between the Kukruse and Haljala stages is rather well expressed by differences between faunas, especially of brachiopods and trilobites in northern Estonia. Both these groups are represented in the Idavere Substage with about 40 species, a few of which occur also in the underlying Kukruse Stage (Rõõmusoks 1970, table 12). The occurrence of several Kukruse bryozoans in the upper part of the Idavere Substage (in the Vasavere Formation) is seemingly of facies origin. The lower boundary of the Haljala Stage is also marked by rather sharp changes in the composition of ostracodes (Põlma et al. 1988, figs. 7, 9-11), though some typical Idavere - Jõhvi species, e.g. Braderupia asymmetrica (Neckaja) appear in the top of the Kukruse Stage. The frequent occurrence of Leiosphaeridia above it, is a rather good marker for the lower boundary of the Haljala Stage. The base of the Haljala Stage is close to both the graptolite Diplograptus multidens Zone and the conodont Baltoniodus gerdae Subzone (Männil 1990, Jaanusson 1995).
The changes in the faunal composition on the transition between the Idavere and Jõhvi substages are continuous which is clearly revealed in core sections, especially by ostracodes (Põlma et al. 1988). Several new macrofossil taxa, such as Toxochasmops maximus (Schmidt), Clinambon anomalus (Schlotheim), presumably appear somewhat higher (1.5 - 2 m, Männil 1963a, b) of the boundary K-bentonite bed between the Idavere and Jõhvi substages.
In the core sections located far away from the outcrop area, the alga Mastopora concava (Eichwald), spicules of Pyritonema subulare (Roemer) and also some brachiopods (Bilobia) occur (Fig. 40), but there is no characteristic species among macrofossils for determination of the lower boundary of the Haljala Stage.
L. Hints & T. Meidla
In most of northern Estonia, the Keila Stage (Kegelsche Schicht, Schmidt 1881) comprises the argillaceous bioclastic limestones, with intercalations or occasionally thicker (up to 4 m) intervals of relatively pure limestones of the Kahula Formation (Table 7). Only in a restricted area in northwestern Estonia, the upper part of this formation is replaced by the Vasalemma Formation where the greatest thickness of the Keila Stage (more than 30 m) has been recorded (Fig. 46).
Initially, due to the unclear relationship between the fossilifereous argillaceous limestones of the Keila Stage and the Schmidt‘s “Wassalem’sche Schicht”, the term Keila-Vasalemma Stage was introduced by Bekker (1922, see also Öpik 1930b). Later, Jaanusson (1945) and Männil (1958c, 1963b, 1966) subdivided the Keila Stage into several members and defined the lower boundary of the stage on the level of the thickest K-bentonite (bed “d” by Jürgenson 1958a, see also Jaanusson & Martna 1948, Vingisaar 1972). The composition of the Kahula Formation and the distribution of members overlying the boundary K-bentonite is shown in Figure 47.
The lowermost part of the Keila Stage (Kurtna Member) is represented by argillaceous limestones. The Kurtna Member is overlain by relatively pure limestones, in places with argillaceous intercalations of the Pääsküla Member. This unit, although differently understood by stratigraphers (Nõlvak 1996), can be identified in the core sections of northwestern Estonia as a complex of biomicritic limestones, up to ca. 7 m in thickness (Põlma et al. 1988, Fig. 47). It may be replaced by intercalating argillaceous bioclastic and biomicritic limestones with a thickness of up to 20 m (Ainsaar 1991), seemingly corresponding to a longer time interval than the Pääsküla Member in the sense of Põlma et al. (1988).
The younger part of the Keila Stage comprises the Saue and Lehtmetsa members, the fossiliferous argillaceous limestones and detrital marls with thin layers of argillaceous limestones, respectively. Contemporaneously, the formation of carbonate buildups (interpreted as reefs, Raymond 1916, or bioherms, Männil 1960, or mud mounds, Põlma & Hints 1984) has been developed in northwestern Estonia. They belong to the Vasalemma Formation, nearly constituting the upper half of the Keila Stage in the surroundings of Keila - Vasalemma, whereas a distinct eastward shift of the corresponding facies is recorded during late Keila time (Fig. 48). The Vasalemma Formation, in thickness up to 15 m, consists of several principal lithotypes (Männil 1960, Põlma 1967, Hints 1996). The most characteristic type of rock is the bedded grainstone (cystoid limestone), which is intercalated with clayey limestones in the lower part of the formation. Cystoid limestone consists mainly of skeletal sand particles aggregated with pure calcite cement (content of terrigenous material less than 3%). The grainstones contain irregular buildups, measuring up to 10 m vertically and up to 300 m horizontally and consisting of pure limestones with a low content of skeletal sand (less than 10%) and terrigenous material (up to 6%), occasionally with inclusions of fossiliferous marls. The buildups mostly lack the reef-like framework and are considered as carbonate mounds. Still, in some “mounds” the edrioasteroid Cyathocystis rhizophora Schmidt is frequent and may form frame-like structures. The lower and middle parts of the Vasalemma Formation contain a number of species common with the Kahula Formation - Estlandia pyron silicificata Öpik, Clinambon anomalus (Schlotheim), Horderleyella? kegelensis (Alichova), Sowerbyella (S.) cf. forumi Rõõmusoks a.o., which indicate the Keila Age of corresponding rocks.
In southern Estonia, the Keila Stage comprises the upper part of the bioclastic limestones of the Adze Formation and clayey limestones and marls, which contain some species common with the Blidene Formation in Latvia (Ulst et al. 1982). The lowermost part of the Mossen Formation may also be of Keila Age (Table 7, Meidla 1996). Still, the correlation of the Lukštai and Blidene formations with a unit of siltstone and silty limestone identified in southern Estonia (Ainsaar 1995) needs to be adjusted. Due to this uncertainty, the identification of the Keila Stage is complicated in the transition between the distribution areas of the Kahula and Adze formations.
The total thickness of the Kahula Formation may exceed 30 m, and in northwestern Estonia its main part corresponds to the Keila Stage. In general, the thickness of the Keila Stage part of the formation (mostly 10-15 m) decreases in the southeast direction. In the same direction, the formation becomes lithologically more uniform and argillaceous. In southern Estonia, the thickness of the equivalents of the Keila Stage presumably does not exceed 10(?) m.
A rich and diverse fauna of bryozoans, brachiopods, trilobites, echinoderms and other sedentary and vagile groups (see Rõõmusoks 1970) is distributed in the Kahula Formation. In the upper part of the formation, corresponding to the Keila Stage, several macrofossil taxa are common with the Haljala Stage, but a specific component in this particular association comprises last representatives of several brachiopod genera (Clinambon, Cyrtonotella), trilobites (Asaphus (Neoasaphus) nieszkowskii Schmidt and Toxochasmops maximus (Schmidt)), crinoids (Ristanacrinus marinus Öpik and different baltocrinids) or species characteristic of the Keila Stage only Keilamena occidens (Männil), Longvillia asmusi (Verneuil), Horderleyella? kegelensis (Alikhova). The data on macrofauna come mostly from northern Estonia. In southern Estonia, the Keila Stage is characterised by a brachiopod - trilobite association, which comprises several taxa (Skenidioides, “Sampo”, Eoplectodonta), appearing on a higher stratigraphical level in northern Estonia or being related to the Scandinavian faunas.
The Keila Stage presumably corresponds to the uppermost part of the Diplograptus multidens and the lowermost part of the Dicranograptus clingani graptolite zones (Männil 1990). The lower boundary of the stage, the level of the K-bentonite bed “d” corresponds to the lower boundary of the chitinozoa Angochitina multiplex Subzone (Table 7) and is close to the Northern Atlantic conodont superbus Zone (Männik & Viira 1990).
L. Hints & T. Meidla
In northern Estonia, the Oandu Stage comprises rocks of two different lithofacies forming the Vasalemma and Hirmuse formations. The Vasalemma Formation, distributed in northwestern Estonia, consists of bedded fine- to coarse-grained bioclastic limestones with irregular bodies of aphanitic massive limestones (carbonate buildups). These rocks were identified first as the Hemicosmites Limestone (Eichwald 1854a) or Wasalemm’sche Schicht (Schmidt 1881). Vasalemma, as the name of a chronostratigraphic unit was applied also to the rocks of another lithofacies – the argillaceous limestones and marls, named Oandu beds by Öpik (1933) and the Hirmuse Formation by Männil and Rõõmusoks (1984), which are exposed on the banks of the Oandu River in northeastern Estonia (Rõõmusoks 1953, Aaloe et al. 1958). Later studies (Männil 1958c, 1960) showed that the lower and middle parts of the Vasalemma Formation are of Keila Age (Fig. 48) and the name Oandu was proposed for chro-nostratigraphic unit of post-Keila Age. The Oandu Age of the uppermost Vasalemma Formation is presumed by the appearance of the corals Lyopora tulaensis (Sokolov), Eofletcheria orvikui Sokolov and the brachiopods Rhynchotrema? parva Oraspõld, Rostricellula nobilis (Oraspõld), Dactylogonia luhai (Sokolskaya) (Männil 1960, Rõõmusoks 1970), or it is supposed by the disappearance of Leiosphaeridia and brachiopods of the Keila Stage (Rummu core, Põlma et al. 1988). In northern Estonia, the Oandu Stage is restricted in thickness (1-4 m, Fig. 49); only in the limits of the Vasalemma Formation it is up to 6 m thick.
In the stratotype area in northeastern Estonia, the lower boundary of the Oandu Stage and the Hirmuse Formation, is known only by the core sections where it is marked by a sharp discontinuity surface with up-to-35-cm-deep pockets, on the upper boundary of the Kahula Formation (= Kahula Group, Männil & Meidla 1994). Below this level, a great number of Middle Ordovician species and even genera common with several older stages, including the brachiopods Cyrtonotella, Estlandia, trilobites Asaphus (Neoasaphus) nieszkowskii Schmidt, Pseudobasilicus, ostracodes Tetrada (Tetrada) harpa (Krause), Polyceratella spinosa Sarv (Fig. 50) disappear. Notable is the disappearance or sharp decrease in the frequency of the acritarch Leiosphaeridia which is abundant in the Keila Stage (Fig. 51). This fossil seemingly can be used for the preliminary establishing of the above-mentioned boundary in core sections, especially when the uppermost part of the Kahula Formation is more argillaceous and possibly belongs to the Lehtmetsa Member of late Keila Age (Fig. 47, Põlma et al. 1988, fig. 32). A new complex of fossils with the ostracodes Bolbina rakverensis (Sarv), Klimphores minimus (Sarv), Disulcina perita perita (Sarv), brachiopods Howellites wesenbergensis (Alichova), Equirostrata wesenbergensis (Teichert) appears near the lower boundary of the Oandu Stage (Fig. 50, see also Põlma et al. 1988). In some cases, these species are found even below the boundary discontinuity surface, seemingly they occur in the deep pockets filled with deposits of Oandu Age. Due to the essential changes in the faunal composition (Männil et al. 1966, Hints et al. 1989), several authors have suggested to use the lower boundary of the Oandu Stage as the regional subseries or series boundary (Jaanusson 1945, Rõõmusoks 1956).
The Hirmuse Formation thins out within a rather short distance in the southern direction, and in many places in central Estonia the Oandu Stage is represented only by the Tõrremägi Member of the Rägavere Formation with a thickness less than one metre (Fig. 49). This area separates the rich and diverse fauna of bryozoans, brachiopods, echinoderms and trilobites of the Hirmuse Formation in northern Estonia (Põlma et al. 1988) from the relatively rich brachiopod and trilobite fauna in the marls and argillaceous limestone in southern Estonia, corresponding presumably to the Lukštai Formation. Beside some species common for northern and southern Estonia (Howellites wesenbergensis Alichova, Rhactorthis kaagverensis Hints a.o.), several brachiopods (Reushella magna Hints, Laticrura sp. Skenidioides sp., Leptellina? sp.) have been identified only in the latter region and are also known in the Lukštai Formation in Lithuania or in the Moldå Limestone in Sweden (Jaanusson 1982, fig. 7).
Identification of the Oandu Stage is most complicated in southeastern Estonia where the black shales and the overlying marls of the Mossen Formation are distributed (Karula core in Fig. 51). The shales, encountered in various sections have been included to the Keila (Meidla 1996), Oandu (Hints 1975) or Rakvere (Männil 1966) Stage. The overlying marls of the Priekule Member are correlated with the uppermost Oandu and/or the Rakvere Stage. In some sections, the marls below the black shale (about 4 m in the Karula core, Fig. 51) comprise brachiopods known in the other sections (Otepää, Laeva) mainly in the beds presumably of Oandu Age. In favour of this age testifies also the disappearance of Leiosphaeridia at a depth of about 3 m below the shales. At the same time, the ostracode record allows to suppose Keila Age of the lowermost part of the Mossen Formation (Meidla 1996). The contradiction in interpreting the age by macrofossils and ostracodes may be caused by insufficient data available, or it may indicate the patchy distribution of deposits in the Keila - Oandu boundary interval in southeastern Estonia.
Still, in most of southern Estonia, the Oandu Stage can be identified most realiably on the basis of ostracodes. The lower boundary of the Oandu Stage is marked by the appearance of Sigmoopsis granulata Sarv, Bolbina rakverensis Sarv, Pelecybolbina illativis Neckaja and Klimphores minimus (Sarv), and a general rapid faunal change which occurred throughout the Estonian part of the palaeobasin (Meidla 1996).
L. Hints & T. Meidla
The “Wesenbergsche Schicht” by Schmidt (1881) corresponds roughly to the Rakvere Stage in nowadays understanding (Männil 1958b, 1963a, Kõrvel 1962, Põlma et al. 1988). In northern Estonia, the Rakvere Stage forms the lowermost, relatively thick part of Late Viru and Harju pure micritic (aphanitic) limestones which intercalate with more or less argillaceous varieties. The cycles of different lithotypes generally constitute distinct lithostratigraphical units (Põlma 1982, Hints et al. 1989), whereas the clayey parts of the cycles are characterized by the appearance of abundant new taxa .
The Rakvere Stage consists of the Piilse and Tudu members (Kõrvel 1962) which form the main part of the Rägavere Formation. The stage is at its thickest (28 m) in western Estonia (Fig. 52) and it thins notably in the southeastern direction. The lower, Piilse Member with a thickness of up to 27 m (Rõõmusoks 1983) consists of pure, in places dolomitized micritic limestones with a low content of terrigenous material (3 - 9%) and skeletal sand (less than 5%, Kõrvel 1962, Põlma et al. 1988). The member is characterised by a distinct pyritic pattern, following abundant burrows in the former sediments. The upper, Tudu Member is up to 10 m thick and differs from the Piilse Member in the higher content of skeletal sand (commonly 15 %, Põlma et al. 1988), in the occurrence of thin, up to 3 cm thick kukersite layers and rare and weakly developed pyritic patterns.
Southwards the limestones of the Rägavere Formation become more argillaceous and in southern Estonia they are supposedly replaced by the carbonate marls of the Priekule Member in the upper part of the Mossen Formation (Männil & Meidla 1994, Meidla 1996). On the basis of chitinozoan distribution (Nõlvak & Grahn 1993) it is supposed that in some places (Ruhnu and Ohesaare cores) the Rakvere Stage is missing.
The data on the distribution and composition of macrofossils in the Rakvere Stage, particularly in the Tudu Member is scanty due to relatively few outcrops. The Rakvere Stage with the relatively sparse macrofauna of bryozoans, brachiopods and trilobites is characterized by frequent and diverse association of calcareous algae (Cyclocrinites, Rhabdoporella etc., Kõrts et al. 1990). From the Rakvere Stage up to the end of Ordovician, calcareous algae and their fragments dominate in the composition of skeletal particles (Põlma 1972, 1982) where they may account even for 97.6%.
Unlike macrofossils, the ostracode record of the Rakvere Stage is rich, comprising more than 80 species (Meidla 1996). Several distinct associations have been recorded in the lower part of the stage corresponding approximately to the Piilse Member of the Rägavere Formation (Meidla 1996, fig. 47). Valuable is the ostracode record of the upper part of the stage which contains only sparse macrofauna. This interval, nearly equal to the Tudu Member, corresponds to the Daleiella admiranda Subzone (subzone of Daleiella sp. n. in Meidla & Sarv 1990, Table 10), a range zone prominent in the sections of northern and central Estonia. The appearance of several long-ranging taxa, such as Steusloffina cuneata (Steusloff), Medianella blidenensis (Gailite), Pullvillites laevis (Abushik & Sarv ), etc., has been recorded within this interval (Fig. 50).
The lower boundary of the Rakvere Stage is lithologically more or less distinct in northern Estonia where it coincides with the pyritized discontinuity surface on the top of the Tõrremägi Member in the lower part of the Rägavere Formation. The appearance of several new brachiopods, including Microtrypa estonica Rõõmusoks, Platystrophia lutkevichi satura Oraspõld, P. quadriplicata Alichova, Sowerbyella (Sowerbyella) raegaverensis Rõõmusoks, Vellamo wesenbergensis (Pahlen) and trilobites Chasmops wesenbergensis (Schmidt), Encrinuroides seebachi (Schmidt), Pharostoma pediloba (Roemer) and others, above this boundary shows the renovation of faunal associations. However, macrofossils are very scarce, particularly in core sections, and cannot be used for the purposes of detailed biostratigraphy. The same applies to ostracodes, because most of the species characteristic of the Late Ordovician ostracode fauna appear below this boundary, in the Tõrremägi Member of the Oandu Stage which represents a facies similar to the Rakvere Stage. Also the zonal chitinozoa Fungochitina fungiformis, characteristic of the Rakvere and Nabala stages, appears in the Tõrremägi Member. The suggestion to include the Tõrremägi Member to the Rakvere Stage (Meidla 1996) follows partly the earlier wider interpretation of that stage, according to which the Oandu beds by Öpik were included to the Rakvere Stage (Jaanusson 1945, Alichova 1960).
The Rakvere Stage corresponds roughly to the lower part of the graptolite Pleurogratus linearis Zone and the chitinozoan Cyathochitina angusta Subzone of the Fungochitina fungiformis Zone (Nõlvak & Grahn 1993, Table 7).
L. Hints & T. Meidla
The Nabala Stage was distinguished by Männil (1958b) as the lower part of the Schmidt’s (1858) “Lyckholm’sche Schicht”. He subdivided the stage into the lower, Paekna and the upper, Saunja substages (see also Öpik & Laasi 1937, Jaanusson 1994). Nowadays they are used as lithostratigraphical units (formations) which represent the Nabala Stage with a total thickness of 10 to 35 m in northern and partly in central Estonia (Fig. 53). The Paekna Formation is up to 16 m thick and comprises predominantly argillaceous bioclastic limestones intercalating with micritic limestones. The thickness of micritic interlayers is usually about 0.1 m, but occasionally it may reach 1-3 m. The up-to-28-m-thick micritic limestones of the Saunja Formation are lithologically uniform and occur all over Estonia. Their thickness decreases towards the south until it is only 0.3 m (Meidla 1996). South of the Muhu - Mustvee line, the lower Paekna Formation is replaced by the Mõntu Formation. This 3–7-m-thick complex consists of argillaceous bioclastic limestones with rare thin (5–30 cm) layers of micritic limestones containing glauconite (Oraspõld 1995).
On the transition from the Rakvere to the Nabala Stage the shelly fauna undergoes notable renovation. Of about 150 species and subspecies occurring in the Nabala Stage, only one third is common with the Rakvere Stage (Männil et al. 1966). Several new species of brachiopods, including Bekkeromena semipartita (Roemer), Ilmarinia sinuata (Pahlen), Laticrura rostrata Hints, Sulevorthis lyckholmiensis (Wysogorski), Pseudolingula quadrata (Eichwald), appear in the Nabala Stage. Some of these species are missing in the Saunja Formation, but appear again in the overlying Vormsi Stage. The micritic limestones of the Saunja Formation contain a notably abundant and diverse fauna of molluscs (about 30 gastropod and more than 10 cephalopod species).
The lower boundary of the stage is exposed only in the Paekna quarry (Nõlvak & Meidla 1990). It is marked by a series of uneven discontinuity surfaces, above which there appears a new association of chitinozoans, including the zonal Armoricochitina reticulifera (Grahn). The latter can be used as the most reliable fossil for the identification of the lower boundary of the Nabala Stage in the core sections in central and southern Estonia.
The composition of ostracodes changes remarkably on the Rakvere - Nabala transition. Several new taxa, including Disulcina perita explicata Sarv, Tetrada neckajae Meidla, Oepikella luminosa Sarv a.o., appear in the lower part of the Nabala Stage, but mainly somewhat higher of the lower boundary of the Paekna Formation (Fig. 50). For this reason the lower boundary of the stage is marked better in the ostracode record by the disappearance of the species Disulcina perita perita (Sarv) and Daleiella admiranda Meidla (a zonal species) in the uppermost part of the Rakvere Stage (Meidla 1996). Brachiopods are found mostly in the lower Paekna Formation and among them new faunal elements Pseudolingula quadrata (Eichwald) and Sulevorthis lyckholmiensis (Wysogorski) appear close to the lower boundary of the Nabala Stage (Fig. 54).
The Nabala Stage corresponds to the middle part of the Pleurogratus linearis graptolite Zone (Table 7) and the upper part of the North Atlantic superbus conodont Zone. In the ostracode record, the summary differences between the Paekna and Saunja formations are not significant, but the very uneven distribution of ostracodes in the Saunja Formation should be mentioned (Meidla 1996).
L. Hints & T. Meidla
The Vormsi Stage (Jaanusson 1944b, = middle part of the Lyckholm’sche Schicht, Schmidt 1858) consists of a facies succession of bioclastics limestones (Kõrgessaare Formation, up to 21 m) in northern Estonia, argillaceous limestones with glauconite (Tudulinna Formation, up to 17.1 m) in central Estonia and black shales (Fjäcka Formation, up to 4.5 m) in southern Estonia. The thickness of the stage decreases from 10 - 20 m in northern Estonia to 0.3 m in southern Estonia (Ohesaare core, Fig. 55). In the transitional area between the Kõrgessaare and Tudulinna formations, interfingering of these units can be followed (Oraspõld 1982a, figs. 3, 4).
The association of the diverse shelly fauna of corals, bryozoans, brachiopods, molluscs and trilobites includes some 200 species (Rõõmusoks 1967) in northern Estonia in the Kõrgesssaare Formation. Southwards this fauna is replaced by a specific and less diverse association.
The Tudulinna Formation is characterized by an association of brachiopods, comprising species of the genera Dicoelosia, Christiania, Skenidioides, Leptellina?, and a facies dependent ostracode association prevailed by Uhakiella curta Sidaraviciene, Medianella blidenensis (Gailite) and Rectella nais Neckaja (Meidla 1996). The Fjäcka Formation comprises a brachiopod association typical of shally facies, consisting mainly of inarticulated small brachipods Paterula, Hisingerella a.o., and of a few articulate brachiopods, such as “Sericoidea” and Onniella (Fig. 56). In general, the association is similar to that in the Mossen Formation.
The lower boundary of the Vormsi Stage coincides with a lithologically sharp boundary in most of Estonia. Above that boundary the frequency of macrofossils and ostracodes increases notably. In the Kõrgessaare Formation, several new species appear, including the corals Proheliolites dubius (Schmidt), Kenophyllum siluricum (Dybowski), Streptelasma (Streptelasma) distinctum Wilson, the brachiopods Eoplectodonta schmidti (Lindström), Equirostra gigas Schmidt, Sampo hiiuensis Öpik, Glyptorthis plana Oraspõld, Triplesia insularis (Eichwald), the first Dicoelosia (Wright 1968) and the trilobites Encrinurus moe Männil a.o. Ostracodes are dominated by the species common with the older strata (Fig. 57). The earliest conodonts of the ordovicicus Zone occur also in the Vormsi Stage (Männik1992b).The zonal species Amorphognathus ordovicicus has been recorded from the basal part of the Kõrgessaare Formation and also from the Tudulinna Formation.
In spite of distinct lower and upper boundaries of the Vormsi Stage, the detailed correlation of the formations belonging to this stage is not yet very clear. The distribution of zonal chitinozoans allows to suppose that the oldest part of the Vormsi Stage is missing in central and southern Estonia (Nõlvak & Grahn 1994). In these areas the Vormsi Stage corresponds to the Tanuchitina bergstroemi Zone which forms the upper part of the stage in northern Estonia, overlying the Fungochitina fungiformis Zone (Table 7) of the lower part of the stage in this area.
The distribution of ostracodes (Meidla 1996) does not support the correlation schemes based on chitinozoans.
The topmost part of the Vormsi Stage correponds to the chitinozoa Acanthochitina barbata Subzone (Table 7). The level of the disappearance of the index species marks well the traditional upper boundary of the Vormsi Stage.
L. Hints & T. Meidla
The Pirgu Stage (Jaanusson 1944b) is a lithologically variable (Table 7) and thick (up to 66 m, Fig. 58) stratigraphical unit (=upper part of the “Lyckholm’sche Schicht”, Schmidt 1858). The notable changes in the thickness of the stage, sometimes within a short distance, are due to various reasons, e.g. the development of mud mounds, denudation during late Pirgu and/or Porkuni time, intensive tectonical movements and changes in sea-level.
In northern Estonia, within the stage two succesive rock units of grey-coloured limestones are distinguished - the lower, Moe and the upper, Adila Formation (Rõõmusoks 1960) which correspond approximately to the former Nyby (Jaanusson 1944b, Männil 1966) and Piirsalu (Jaanusson 1945) substages.
The Moe Formation, up to 40 m in thickness, consists of micritic and bioclastic nodular or bedded limestones with argillaceous intercalations. The lower part of the formation contains abundant calcareous alga Palaeoporella? (=Dasyporella in Männil 1966) and in some places (Hoitberg on Vormsi Island, Võhma core in central Estonia) typical carbonate mounds are developed, quite similar to the Boda mounds in the Siljan district of Sweden.
The Adila Formation comprises predominantly bioclastic limestones with a thickness of 10-15 m. Numerous discontinuity surfaces and cyclically alternating pure and argillaceous limestones are characteristic to the upper part of the formation. In this topmost part of the formation, the pentamerid brachiopod Holorhynchus has been recorded in the Island of Hiiumaa (Hints 1993). The boundary between the Moe and Adila formations coincides with the boundary between the rugata and bergstroemi chitinozoan zones (Table 7).
Within a rather large area in central Estonia, the Pirgu Stage is characterized by the interfingering of different rock units (Oraspõld 1975a, Table 7), whose correlation with the northernmost and southernmost sequences is complicated. In this transitional area, the lowermost part of the stage consists of argillaceous bioclastic limestones with glauconite (Tootsi Member), overlying the upper part of the Vormsi Stage, corresponding to the chitinozoan barbata Zone. The Tootsi Member contains a distinct association of shelly fauna (Fig. 59). This unit is succeeded upwards by the grey-coloured, sometimes red mottled marls and highly argillaceous limestones of the main part of the Halliku Formation whose relationship with the Moe Formation is not yet very clear (Männil & Meidla 1994). The diverse and abundant ostracode fauna of the Halliku Formation comprises taxa (Fig. 60) common with the upper part of the Moe Formation (Meidla 1996). Most uncertain is the age of the Kabala Member on the transition between the Pirgu and Porkuni stages in central and westernmost Estonia. According to the ostracode record, the formation contains two different associations. The older association comprises the Pirgu species Brevibolbina pontificans Schallreuter, Bullaeferum tapaensis (Sarv), whereas the species Apatochilina falacata Sarv and Gryphiswaldensia plicata Schallreuter from the supposedly younger part are common with the fossiliferous part of the Ärina Formation of the Porkuni Stage.
In southern Estonia, in the limits of the central confacies belt, the Pirgu Stage is represented by red-coloured or mottled argillaceous limestones and mudstones of the Jonstorp and Jelgava formations, including the Kuili Member (Table 7).
According to the traditional understanding, the lower boundary of the Pirgu Stage coincides with the lower boundary of the Moe Formation in northern Estonia. This level is underlain by the chitinozoan Acanthochitina barbata Zone. The zone has also been established under the red-coloured limestones of the Jonstorp Formation (Ruhnu core). On this basis, the lower boundary of the latter formation has been equalized with the lower boundary of the Pirgu Stage in southern Estonia. In terms of graptolite zonation, the stage boundary corresponds to the base of the D. complanatus Zone (Männil 1990). The presence of Climacograptus supernus Elles et Wood in the upper part of the Pirgu Stage suggests its correlation with the Dicellograptus anceps Zone (Männil 1976, 1990).
The lower boundary of the Pirgu Stage is poorly reverberated in the distribution of most shelly fossil groups. As a rule, the new faunal elements appear 1 - 2 m above (or even higher) of the distinct lithological changes (Fig. 59). In northern Estonia, the biostratigraphical boundary is best expressed by the appearance of the concentrations of the alga Palaeoporella?. In central and southern Estonia, the faunal renovation is best revealed in the ostracode record (Meidla 1996, see also Fig. 60).
The Pirgu Stage comprises three different assemblages of macrofauna, related to different facies zones of the palaeobasin. In northern Estonia, in the grey-coloured Moe and Adila formations a rich assemblage of large articulated brachiopods Plaesiomys solaris (Buch), Equirostra gigas Schmidt, Triplesia insularis (Eichwald), Luhaia vardi Rõõmusoks, corals Sarcinula, Catenipora, Palaeofavosites, and stromatoporoids, together with different molluscs, is distributed (Fig. 59). In the ostracode composition nonpalaeocopes are dominating: the associations of Steusloffina cuneata-Medianella blidenensis and S. cuneata-Olbianella fabacea (Meidla 1996, Fig. 60) occur. In central Estonia, in the Halliku Formation the most common representatives of the shelly fauna seem to be brachiopods and rugose corals (Fig. 59). In the red-coloured deposits of southern Estonia, only a few macrofossils, mainly brachiopods and trilobites, have been recorded, while trilobite and echinoderm fragments dominate in the skeletal sand (Männil et al. 1968).
L. Hints & T. Meidla
The Porkuni Stage (Borkholm’sche Schicht by Schmidt 1881) represents the topmost Ordovician stage (Raymond 1916, Bekker 1922). Up to the 1960s, the stage was included to the Silurian System (Öpik 1930b, 1934, Aaloe et al. 1958, Alichova 1960, Rõõmusoks 1960). In Estonia, the Porkuni Stage is represented by variable deposits of shallow-water facies (Männil 1966; Oraspõld 1975b, 1982b; Rõõmusoks 1983), with a thickness of about 10 m in northern and up to 18 m in southern Estonia (Fig. 61). In northern Estonia, the stage is supposedly represented by its older part only, because afterwards, during late Porkuni time, this area turned into dry land as a result of the glacioeustatic sea-level lowering (Oraspõld 1975b).
In northern Estonia, the Porkuni Stage is represented by the Ärina Formation comprising a succession of dolomites (Röa Member), stromatoporoid-tabulate reefs (with surrounding facies) and oolitic or sandy limestones (Kamariku Member) in the top. The assignment of the Röa Member (0.5-5.5 m of dolomites) has been problematic over the years. Some researchers have assigned it to the Pirgu Stage (Rõõmusoks 1991), others to the Porkuni Stage (Rosenstein 1943, Jaanusson 1956, Resheniya… 1987). The unit, usually poor in fossils, yields some species common with the Pirgu Stage (Rõõmusoks 1989). In many sections the lower boundary of the member is lithologically sharp, except the areas where the topmost part of the Adila Formation is dolomitized. The upper boundary is transitional. Here, the assignment of the Röa Member to the Porkuni Stage is conventional.
Small reef bodies, recorded in the middle part of the Ärina Formation (2-3 m high, up to 20 m wide, traditionally treated as the Tõrevere Member), yield the tabulate corals of the Palaeofavosites rugosus community (Klaamann 1986) and the stromatoporoids Clathrodictyon mammillatum (Schmidt), Ecclimadictyon porkuni (Riabinin) a.o. The reefs are surrounded by skeletal limestones (Vohilaid Member, up to 3.7 m) and kerogenous limestones (Siuge Member, up to 2.6 m; see Oraspõld 1975b, Rõõmusoks 1983), apparently representing the pre-reef and inter-reef facies, respectively. In the western part of mainland Estonia, the Vohilaid Member, which is often represented by pure skeletal sand in sparry (?) calcite matrix, contains thin (up to 20 cm) layers of oolitic limestone, whereas the ooids make up 10-45% of the rock volume (Oraspõld 1975b).
In core sections, the common succession of the three lithotypes of the “reef complex” begins with skeletal limestones which are overlain by kerogenous and reef limestones (Fig. 62). The rocks contain a rich and diverse macrofauna of corals, brachiopods, gastropods etc., (more than 150 species and subspecies; Männil 1962, Rõõmusoks 1970). The associations characteristic of the particular lithotypes have many species in common: rugose corals Konodophyllum rhizobolon (Dybowski), Streptelasma (Streptelasma) giganteum (Kaljo), brachiopods Streptis undifera (Schmidt), Schmidtomena acuteplicata (Schmidt), trilobite Platylichas mastocephalus (Öpik) and others. Among microfossils, ostracodes are abundant (Meidla 1996), whilst conodonts are extremely rare (Männik 1992b).
South of the distribution area of the Ärina Formation, distinction of the Porkuni Stage is complicated. In some sequences (Ohesaare core) the Porkuni Stage is obviously missing. In many sections in central Estonia (Are and Kahala etc., Fig. 62), the topmost part of the Ordovician sequence is represented by 1- 2m-thick dolomites, which may correspond to some part of the Ärina Formation (?Röa Member). The distribution area of these dolomites coincides roughly with the area of the pre-Silurian (early Silurian?) channeling where the erosion reached the pre-Porkuni rocks (Perens 1995).
In southern Estonia, the Porkuni Stage is represented by the peripheral parts of the Kuldiga and Saldus formations (Ulst & Gailite 1982). The Kuldiga Formation of bioclastic limestones and marls, overlain by the silty and sandy limestones of the Saldus Formation, comprises the cosmopolitan Hirnantia fauna (Rong & Harper 1988). Hirnantia sagittifera (M’Coy), Dalmanella testudinaria (Dalman), Plectothyrella crassicosta (Dalman), typical elements of that fauna, have been identified in the core sections of Ruhnu, Ikla and Taagepera. These species appear in the lower part of the Kuldiga Formation roughly on the level where the ostracodes common with the Ärina Formation and a zonal chitinozoa Spinachitina taugourdeaoui (Eisenack) disappear. The new ostracodes appearing in the Kuldiga Formation seem to have an extraordinarily wide geographical distribution and probably form a part of the Hirnantia fauna sensu lato (Meidla 1996).
The youngest Ordovician deposits corresponding to the Glyptograptus persculptus graptolite Zone are identified only on the western coast of the East Baltic (Ulst 1992). There is no certain evidence on the occurrence of shallow-water deposits of the persculptus Zone in Estonia, although they may be present as the unfossiliferous topmost Ordovician or even in strata assigned to the lowermost Silurian (Kaljo & Hints 1996).
Concluding the data on the terminal Ordovician in Estonia, it should be mentioned that the Porkuni Stage, in the presents limits, comprises rocks of different age. The oldest part of the stage is present in the stratotype area in northern Estonia, while the most complete sequences presumably occur in southern Estonia. The appearance level of the Hirnantia fauna, which may be correlated with the lower boundary of the Hirnantia Stage in Scandinavia, lies seemingly in the lower part of the Porkuni Stage in the East Baltic.
H. Nestor
The first stratigraphical classification of the Silurian rocks in Estonia was worked out by Schmidt (1858, 1881, 1892). Bekker (1922, 1925) and Luha (1930, 1933, 1946) established the present nomenclature of the Silurian regional chronostratigraphical units - regional stages. Lithostratigraphical divisions have been adequately defined in the monograph “The Silurian of Estonia” (Kaljo 1970c). Further amendments to the stratigraphical nomenclature and correlation with the sequences of the adjacent areas have been published in the unified regional stratigraphical charts of the Baltic Republics (Resheniya… 1978) and of the East-European Platform (Resheniya… 1987). The latest version of the Silurian stratigraphical chart, approved by the Stratigraphical Commission of Estonia, was published by H. Nestor (1995a) and is followed in the present publication (Table 8 ).
The Silurian sequence in Estonia consists of ten regional stages grouped directly into the series of the global chronostratigraphical standard. In most cases the boundaries of the regional stages and series have been considered more or less congruent, based on the graptolite or conodont datings (Kaljo 1962, Viira 1982). An exception is the Wenlock/Ludlow boundary which is only conventionally fitted with the junction of the Rootsiküla and Paadla stages. The lower limit of the Silurian System coincides with the boundary between the Porkuni and Juuru stages. It is proved by the presence of the Hirnantian trilobites and brachiopods in the Kuldiga and Saldus formations of the Porkuni Stage and records of Stricklandia lens prima Williams from the lowermost beds of the Varbola Formation of the Juuru Stage (Kaljo et al. 1988b).
Based on the sharply expressed lateral, facies changes of the Silurian rocks, the Mid-Estonian and South-Estonian confacies belts have been distinguished (Kaljo 1977). The Mid-Estonian Confacies Belt is dominated by various lime- and dolostones, rich in shelly fauna. The belt covers the islands of the West-Estonian Archipelago and the western and central parts of mainland Estonia (Table 8). In the latter area, the Silurian sequence is less complete; its upper part has undergone severe dolomitization. The South-Estonian Confacies Belt consists mostly of marl- and mudstones with a more unilateral deeper-water shelly fauna, graptolites and planktonic microfossils (chitinozoans). Within the confacies belts separate sets of lithostratigraphical units have been established.
Many parts of the Silurian sequence have a clearly expressed cyclical nature, especially in the more shallow-water Mid-Estonian Confacies Belt. In such cases a cyclostratigraphical unit, the so-called beds consisting of alternating types of rocks with a certain trend of succession, has been distinguished and treated as a subdivision of formation. In some cases formations can be subdivided into members.
The Juuru strata were established by Schmidt (1858) as the Bed (“Jördensche Schicht”), later transferred to the rank of Stage (“Stufe”) (Schmidt 1892). Nestor and Kala (1968) determined the present stratigraphical extent of the stage and worked out its classification. With the Juuru Stage they united the Tamsalu Formation, earlier treated as an independent stage, and the lowermost beds of the Raikküla Stage (now the Karinu Member). The former Hilliste Member of the Juuru Stage was recently expanded and raised into formation rank partly corresponding to the Raikküla Stage (Männik 1992b, Nestor 1995a).
The Juuru drill core in the interval of 0.4-16.2 m has been selected as a neostratotype for the Juuru Stage (Nestor 1993). The Juuru Stage spreads on the islands of Hiiumaa and Saaremaa and in the western, central and southern parts of mainland Estonia. The outcrop extends as a west-eastwards widening belt (4 to 25 km) from midsouthern Hiiumaa as far as the eastern slope of the Pandivere Upland. The main localities are ancient coastal cliffs at Kallasto and Pullapää, quarries at Hilleste, Kirimäe, Karinu, Tamsalu and Rakke (Kamariku) and a well in the ancient Varbola stronghold (Fig. 63). The full thickness of the stage varies from 20.1 m in the Asuküla borehole to 63.7 m in the Viljandi borehole (Fig. 63).
The stage is dominated by biomicritic limestones (packstones, wackestones) rhythmically intercalating with thin layers of marl- and mudstones (argillites, clays) and containing interlayers of sparitic limestones (grain- or rudstones). The proportion of marlstones increases southwards and the number of sparitic interlayers towards the north-west and upwards in the sequence.
The lower boundary of the stage coincides with the base of a thin band of micro- to cryptocrystalline limestone of the Koigi Member or, if the latter is absent, with the base of the marl- or mudstones of the Varbola and Õhne formations overlying various sparitic limestones of the Porkuni Stage, including bioclastic and oolitic grainstones, lithoclastic rudstones of shallow-water origin. Above the boundary, the brachiopod Stricklandia lens, the chitinozoans Ancyrochitina laevaensis and Spinachitina fragilis or the conodont Ozarkodina ex gr. oldhamensis appear.
The Juuru Stage contains a rather rich benthic shelly fauna, whereas planktonic fossils are rare. The most characteristic species are ( abbreviations in brackets: vr - Varbola Formation, tm - Tamsalu Formation, õh - Õhne Formation, pt. - part) Clathrodictyon boreale Riabinin (vr, tm), Paleofavosites paulus Sokolov (vr, tm, õh), Stricklandia lens prima Williams (vr, lower pt.), S. lens lens Williams (vr, upper pt.), Zygospiraella duboisi (Verneuil) (vr), Borealis borealis (Eichwald) (tm), Acernaspis estonica Männil (õh), Calymene ansensis Männil (vr, tm), Aitilia senecta Sarv (vr), Steusloffia eris Neckaja (vr, tm, õh), Ozarcodina ex gr. oldhamensis (Rexroad) (vr, tm, õh), Distomodus kentuckyensis (Branson et Mehl) (vr, tm, õh), Ancyrochitina laevaensis Nestor (õh, basal pt.), Spinachitina fragilis Nestor õh, basal pt.), Conochitina postrobusta Nestor (õh), Dimorphograptus confertus Nicholson) (õh, top), Pribylograptus incommodus (Toernquist) (õh, top). Records of S. lens prima, A. laevaensis and S. fragilis from the basal part of the stage suggest that the base of the Juuru Stage lies on the level of the Parakidograptus acuminatus Zone (Cocks 1971, Nestor V. 1994). Graptolites D. confertus and P. incommodus from the top of the stage confirm that the upper boundary of the stage roughly coincides with the boundary between the Orthograptus vesiculosus and Coronograptus cyphus zones (Kaljo & Vingisaar 1969).
In the Mid-Estonian Confacies Belt, the Juuru Stage is divided into the Varbola (below) and Tamsalu (above) formations. In the South-Estonian Confacies Belt, the Õhne Formation corresponds to both of them (Figs. 64, 65).
The Varbola Formation is represented by nodular biomicritic limestones (skeletal to coquinoid pack- and wackestones) with thin intercalations of marlstone. The formation contains tempestitic interlayers of skeletal grainstones, often with intraclasts, the number of which increases upwards in the sequence and northwestwards in the space. Brachiopods of the Stricklandia Community are characteristic to the formation. The thickness of the formation varies from 8.8 m in the Pusku borehole to 24.6 m in the Käru borehole. The 0.1—3.5-m-thick Koigi Member of micritic (aphanitic) limestones is developed at the base of the Varbola Formation.
The Tamsalu Formation consists of various, prevailingly sparitic limestones (skeletal and pelletal grainstones, coquinoid or lithoclastic rud- and floatstones). The thickness of the formation varies from 8.8 m in the Pusku 2 borehole to 18.5 m in the Rumba borehole. The formation is subdivided into the Tammiku (below) and Karinu (above) members.
The Tammiku Member is typically represented by a bank of coquinoid limestone consisting of shells and debris of the brachiopod Borealis borealis. The thickness of the bank reaches 13.5 m on the Pandivere Upland. In the same area, the Karinu Member consists of skeletal and pelletal grainstones and bio- or lithoclastic rudstones. South- and westwards the latter are replaced by fine-grained grain- and packstones with numerous hardgrounds.
The Hilliste Formation consists of a highly variable assemblage of rock types in which the most characteristic are crinoidal limestones (grainstones) with coral-stromatoporoid bioherms. The formation also contains fine-grained pelletal and skeletal grain- and packstones and micritic limestones. The formation corresponds to the upper part of the Tamsalu Formation (Karinu Member) and to the lower part of the Raikküla Stage (Nestor 1995a). It occurs on Hiiumaa Island and in the vicinity of Haapsalu - Rohuküla and Rapla - Käru in mainland Estonia.
The Õhne Formation is represented by marlstones, mudstones and micritic limestones. It corresponds to the whole stratigraphical extent of the Juuru Stage in southern Estonia. The rather poor fauna corresponds to the brachiopod Clorinda Community. The maximum thickness (63.7 m) has been fixed in Viljandi 91 borehole. The thin, up-to-2.7-m-thick Puikule Member of marlstones and the overlying, up-to-8-m-thick Ruja Member of micritic limestones occur in the basal part of the Õhne Formation along the southern and eastern margins of the area of distribution of the formation.
The Raikküla beds were originally defined (Schmidt 1858) as the “Intermediate zone” (Zwischenzone) between the strata with Pentamerus borealis and P. oblongus. In 1881, Schmidt introduced the geographical name - Raikküllsche Schicht. Kaljo and Vingisaar (1969) presented the currently used subdivision of the stage for southern Estonia. Perens (1992) and H. Nestor (1995a) modernized the classification for the outcrop area. The Mõhküla beds, earlier attributed to the Adavere Stage, were replaced into the Raikküla Stage as they are separated from the rest of the Adavere Stage by a structural disconformity (Nestor 1995a). However, since the stratigraphical level of the Mõhküla beds was changed only recently, it is not yet reflected in the limit between the outcrops of the Raikküla and Adavere stages on the printed geological maps.
The Raikküla-Paka scarp and Raikküla drill core in the interval of 0.5 to 35.0 m have been defined as the composite stratotype of the stage (Nestor 1993). The Raikküla Stage is distributed on the islands of the West-Estonian Archipelago and in the western, central and southern parts of mainland Estonia. The outcrop extends as a latitudinal, eastward widening belt (6 to 45 km) from southern Hiiumaa as far as the southeastern slope of the Pandivere Upland near Palamuse. The main localities are active or abandoned quarries at Pusku, Orgita, Keava, Mündi, Kalana and Rôstla, and ancient coastal scarps (inland cliffs) at Pakamägi and Raikküla-Paka. The thickness of the stage varies from 16.3 m in the Murika borehole to 176.3 m in the Ikla borehole (Fig. 66) and decreases abruptly in the northwest direction due to the end-Raikküla denudation of the upper layers of rocks.
The Raikküla Stage consists of a variety of carbonate rocks. The most characteristic are micritic (micro- and cryptocrystalline) limestones cyclically interbedding with marl- or mudstones in the south and with different bioclastic limestones (wacke-, pack- and grainstones) in the north. In the northernmost sections of central Estonia, the shallowing-up sedimentary cycles may end with argillaceous primary dolomites. In the southernmost sections, the marl- and mudstones contain graptolites on certain levels. In central Estonia, in the Paide - Pärnu belt of faults and eastwards, the Raikküla rocks are strongly dolomitized.
The lower boundary of the stage coincides with the base of a band of marl- or mudstones overlain by thick deposit of monotonous micritic limestones of the Järva-Jaani beds in the north and Slitere Member in the south. In the area of distribution of the most shallow-water sequences of the Hilliste and Raikküla formations the boundary is less definite. Above the boundary, sparse graptolites of the Pristiograptus cyphus Biozone and chitinozoans of the Conochitina electa Biozone (C. electa, C. maennili, etc.) appear.
Fossils are of uneven distribution in the rocks of the Raikküla Stage. The widespread micritic limestones and different dolostones (from pure dolomite to dolomitic marl) contain occasional macrofossils. The most characteristic species are (abbreviations in brackets: rk - Raikküla Formation, nr - Nurmekund Formation, sr - Saarde Formation, u.pt. - upper part, m.pt. - middle part, l.pt. - lower part): Clathrodictyon clivosum Nestor (rk, u. pt.), Parastriatopora celebrata Klaamann (rk, u. pt.), Borealis pumilus (Eichwald) (nr), Borealis borealis osloensis Mjork (nr), Meifodia ovalis Williams (sr), Hermannina hisingeri (Schmidt) (rk, nr), Bythrocyproidea sarvi Neckaja (nr), Icriognathus cornutus Männik (rk, l. pt.), Kockelella manitoulinensis (Pollack, Rexroad et Mehl), (rk, nr), Conochitina electa Nestor (rk, nr, sr, l. pt.), C. iklaensis Nestor (nr, sr), Spinachitina maennili Nestor (sr), Coronograptus cyphus (Lapworth) (sr, l. pt.), C. gregarius (Lapworth) (nr, sr, m. pt.), Demirastrites triangulatus (Harkness) (sr, m. pt.), D. convolutus (Hisinger) (sr, u. pt.). The presence of zonal graptolites shows that the Raikküla Stage spans from the C. cyphus Biozone to the D. convolutus Biozone.
The Raikküla Stage consists of the Raikküla, Nurmekund and Saarde formations, laterally replacing each other from north to south (Figs. 64, 65). The upper part of the Hilliste Formation is of Raikküla Age (Table 8).
The Raikküla Formation is distributed in central and western Estonia, in the Lääne, Rapla and Järva counties. It is represented by two shallowing-up sedimentation cycles star-ting with biomicritic or micritic limestones, succeeded by ske-letal grainstones, pelletal or coral-stromatoporoid limestones, and ending with argillaceous lagoonal dolostones. These cycles are treated as the lower and upper subformations (Nestor 1995a). The thickness of the formation varies from 30 m in the Kiideva borehole to 56 m in the Käru borehole. The upper layers of the formation have undergone considerable denudation and in the westernmost sections of mainland Estonia the upper subformation thins totally out.
The Nurmekund Formation south and east of the Raikküla Formation consists of five sedimentary cycles which begin with a relatively thin layer of marlstone or argillaceous limestone. The main, middle part of the cycle is represented by wavy-bedded micritic limestone, the upper part by bioclastic limestones containing numerous discontinuity surfaces. In central Estonia, the formation is strongly dolomitized, particularly its upper half. The first, third and fifth cycles from below are thicker and more complete, the second and fourth being thinner and less typical. In ascending order, the cycles are termed the Järva-Jaani, Vändra, Jõgeva, Imavere and Mõhküla beds (Table 8). Westwards the upper beds gradually thin out and on Saaremaa Island only the Järva-Jaani and, partly, the Vändra beds are present. The thickness of the formation ranges from 16 m in the Murika borehole to 73+ m in the Võhma borehole.
The Saarde Formation is distributed in southwestern Estonia. It consists of cyclically alternating deposits of horizontally-bedded micritic lime- and marlstones or mudstones and is subdivided into six members. The lowermost, rather thin mudstone member, comprised mostly of argillites, has no name and was earlier included in the Õhne Formation of the Juuru Stage. In ascending order, the Slītere, Kolka, Ikla, Lemme and Staicele members follow. The shaly mudstone interlayers in the Ikla Member abound in graptolites of the Demirastrites triangulatus Zone. In other members graptolites are less frequent. In its full thickness (176.3 m) the Saarde Formation occurs only in the Ikla borehole.
The Adavere Stage as a stratigraphical unit was established by Schmidt (1858) as the uppermost unit (zone 6) of the group of smooth pentamerids (“Gruppe der glatten Pentameren”). Afterwards it was termed the Esthonus-Schicht (Schmidt 1881), Addifer Formation (Twenhofel 1916), Adavere Stage (Bekker 1922). Kaljo (1962) fitted the upper boundary of the stage with the Llandovery and Wenlock boundary and included in it the marlstones of the present Velise Formation. Recently, Perens (1992) and Nestor (1995a) excluded the Mõhküla beds and replaced them into the Raikküla Stage.
The Päri quarry in western Estonia has been selected as the neostratotype of the stage (Nestor 1993, 1995a) and the Kirikuküla core at the depth of 50.3 m may be treated as the boundary stratotype of the stage. The Adavere Stage is distributed in the southernmost part of Hiiumaa Island, on Saaremaa and Muhu islands and in the southwestern part of mainland Estonia as far as the Viljandi fault. The outcrop extends as a 10—15-km-wide belt from the southernmost Hiiumaa Island and the Soela Strait over Matsalu Bay up to the vicinity of Türi - Vändra being denudated eastwards the Paide-Pärnu belt of disturbances. The main localities are the Saastna coast, Päri quarry, river banks at Päärdu, Jädivere, Velise, Valgu, Vändra and ditches at Lätiküla and Valgu (Fig. 67). The thickness of the stage increases westwards — from 10.7 m in the Ristiküla borehole to 56.3 m at Nässumaa.
The Adavere Stage is represented by thin-bedded to nodular biomicritic limestones (wackestones to packstones) with Pentamerus oblongus (below) and marl- to mudstones (above). The former unit is treated as the Rumba Formation and the latter as the Velise Formation. The clay content increases westwards. The lower boundary of the stage coincides with the strongly pyritized erosion surface at the base of the nodular biomicritic limestones of the Rumba Formation, transgressively overlying different strata of the Raikküla Stage. The Adavere Stage contains rather rich shelly fauna of Pentamerus (below) and Clorinda (above) communities. Microfossils (chitinozoans, ostracodes, conodonts) are more frequent in the mud- and marlstones of the Velise Formation, almost devoid of corals and stromatoporoids. The most characteristic species are as follows (abbreviations in brackets: rm - Rumba Formation, vl - Velise Formation): Clathrodictyon variolare (Rosen) (rm), Mesofavosites obliquus Sokolov (rm), Angopora hisingeri (Jones) (vl), Palaeocyclus porpita (Linnaeus) (vl), Prodarwinia speciosa (Dybowski) (rm), Pentamerus oblongus (Sowerby) (rm), Stricklandia laevis (Sowerby) (rm), Dicoelosia baltica Musteikis et Puura (vl), Encrinurus (Nucleurus) rumbaensis Rosenstein (rm), Calymene frontosa Lindström (vl), Beirichia valguensis Sarv (rm), Longiscella caudalis (Jones) (vl), Conochitina emmastensis Nestor (rm), Eisenackitina dolioliformis Umnova (rm, vl), Angochitina longicollis Eisenack (vl), Pterospathodus celloni Walliser (vl), P. amorphognathoides Walliser (vl), Spirograptus turriculatus (Barrande) (vl), Monograptus discus Törnquist (vl), Monoclimacis griestoniensis (Nicol) (vl).
The presence of the index species of graptolites (S. turriculatus, M. griestoniensis) and conodonts (P. celloni, P. amorphognathoides) in the upper half of the Adavere Stage demonstrates that most probably the stage corresponds to the Monograptus sedgwickii to Monoclimacis crenulata biozones.
In Estonia, the Adavere Stage consists of the Rumba (below) and Velise (above) formations. The Rumba Formation spreads on the islands of the West-Estonian Archipelago and in the southwestern part of mainland Estonia. It is represented by horizontally-bedded to nodular biomicritic limestones (wackestones, packstones) with clayey partings and scattered shells or tempestitic accumulations of the brachiopod Pentamerus oblongus. The formation consists of twelve low-grade sedimentary cycles beginning with argillaceous rocks (marlstones, argillaceous limestones) and ending with a layer of pure, hard limestone (Einasto et al. 1972). Westwards the clay content of the rocks increases and on Saaremaa Island marlstones are prevailing in the sequence of the Rumba Formation. A characteristic yellowish-green tuffaceous (metabentonite) interlayer (8 to 18 cm) occurs at the level of the base of the upper third of the sequence.
The thickness of the Rumba Formation is mostly 15 to 19 m and it decreases at the western and eastern margins of the distribution area. Local hiatuses occur in the Ohesaare and Are sections.
The Velise Formation overlies the Rumba Formation and consists of different marlstones and mudstones up to plastic clays. The mostly greenish- to bluish-grey rocks are south- and eastwards replaced by red-coloured (purple) varieties. In the southwesternmost sections (Ohesaare, Ruhnu) graptolites are present in the dark-grey interlayers of argillite. Thin (0.5 to 5.0 cm) metabentonite interlayers are characteristic to the formation. The thickness of the formation is greatest in northwestern Saaremaa, reaching 37 - 38 m in the Viki and Eikla sections. In the southeast direction, it decreases until thinning out in the Ristiküla section, eastern Pärnumaa.
The Jaani Stage was defined by Luha (1933) as a marlstone unit corresponding to the lower part of the “Untere Oeselsche Gruppe (Stufe)” by Schmidt (1858, 1892). Kaljo (1962) separated the lower part of the marlstones (now the Velise Formation), corresponding to the uppermost Llandovery, and joined it with the Adavere Stage. V. Nestor (1984) determined the position of the upper boundary in the subsurface area. Aaloe (1960, 1961) subdivided the Jaani Stage into the Mustjala, Ninase and Paramaja members. Later, Aaloe & Kaljo (1962) distinguished the Tõlla Member for the South-Estonian subsurface area.
A historical stratotype of the Jaani Stage is the sea shore with the Paramaja Cliff in the vicinity of the Jaani Church (Resheniya… 1987, Nestor 1993). The Ohesaare drill core at the depth of 345.8 m may be treated as the boundary stratotype of the stage. The Jaani Stage spreads on Saaremaa and Muhu islands and in the southwestern part of mainland Estonia (Pärnumaa and southern Läänemaa). The outcrop runs along the northern coast of Saaremaa and Muhu islands and the southern coast of Matsalu Bay towards the Vändra Borough (Fig. 68). The main localities are the cliffs at Undva, Suuriku, Ninase, Panga, Liiva, Paramaja, Kautliku, Püssina and Uisu, the quarries at Koguva and Anelema (lower part), and the river bank at Jädivere. The thickness of the stage increases westwards and varies from 24.2 m in the Lihula borehole to 70 m in the Kaugatuma borehole.
The stage consists mainly of various marl- and mudstones. Limestones (skeletal wacke-, pack-, grain- and boundstones) are of minor importance and occur only in the upper half of the stage in the northwestern part of Saaremaa Island (Ninase Member). The lower boundary of the stage has been made congruous with the Llandovery/Wenlock boundary (Kaljo 1962), established by the appearance of the graptolite Cyrtograptus murchisoni in the Ohesaare drill core above the depth level 345.8 m and in other sections by chitinozoans of the Margachitina margaritana Zone (Nestor V. 1994). Lithologically, it usually coincides with a certain increase in the carbonate content of rocks.
The Jaani Stage contains rather rich shelly and planktonic faunas with characteristic deeper-water elements (graptolites, chitinozoans, trilobites). The most typical species are as follows (abbreviations: M - Mustjala Member, P - Paramaja Member, N - Ninase Member, T - Tõlla Member, u.pt. - upper part, l.pt. - lower part): Stromatopora impexa Nestor (M, u. pt.), Halysites senior Klaamann (M), Thecia podolica (P), Neocystiphyllum keyserlingi (Dybowski) (P), Leptaena rhomboidalis (Wahlenberg) (M,P), Eocoelia angelini (Lindström) (N), Pseudobollia krekenawaiensis Neckaja (T,P), Craspedobolbina (C.) mucronulata Martinsson (N,P), Beirichia (B.) suurikuensis Sarv (N,P), Calymene orthomarginata Schrank (T,P), Encrinurus punctatus (Wahlenberg) (P), Conochitina cf. mamilla Laufeld (T,N,P), Calpichitina acollaris (Eisenack) (P), Pterospathodus amorphognathoides Walliser (T,M, l.pt.), Kockelella ranuliformis (Walliser) (M,N,P), Cyrtograptus murchisoni Carruthers (T), Monograptus riccartonensis Lapworth (T), M. flexilis Elles (P). The presence of the index species of graptolites proves that the Jaani Stage spans from the C. murchisoni Biozone to the M. flexilis Biozone.
In Estonia, the Jaani Stage is mainly represented by the Jaani Formation. Only in the southernmost sections, the lower part of the stage has been treated as the Tõlla Member of the Riga Formation (Figs. 69, 70).
The Jaani Formation consists of marlstones and, to a lesser extent, of bioclastic and biohermal limestones. The lower part is formed by the Mustjala Member comprising argillaceous marlstones (Figs. 69, 70), which are often dolomitized, particularly in eastern sections. In the middle of the sequence of the Jaani Stage, the carbonate content increases abruptly and, respectively, the upper half of the Jaani Formation is represented by calcareous marlstones or argillaceous limestones of the Paramaja Member in the eastern part of Saaremaa Island, on Muhu Island and in mainland Estonia. In the northwest direction the Paramaja Member is laterally replaced by bioclastic limestones (wackestones to grainstones) of the Ninase Member containing also bioherms. In many sections, a tongue of the Paramaja marlstones overlaps the Ninase limestones.
In Estonia (Tõlla, Ikla, Ruhnu, Ohesaare drill sections) the Riga Formation is represented only by its lower part - the Tõlla Member, which is characterized by graptolite-bearing grey mudstones. Northwards it is replaced by the greenish-grey marlstones of the Mustjala Member and upwards with the marlstones of the Paramaja Member, both belonging to the Jaani Formation (Figs. 69, 70).
The present unit was established by Luha (1933) provisionally as the Muhu-Kurevere Stage, later as the Jaagarahu Stage (Luha 1946). It corresponds roughly to the upper half of the “Untere Oeselsche Gruppe (Schicht)” by Schmidt (1858, 1881). The subdivision of the stage has been recurrently changed (Bekker 1925, Luha 1930, Aaloe 1970, Aaloe et al. 1958, 1976, etc.). Recently, some additional units, including the Jamaja and Riksu formations, were introduced (Resheniya… 1987, Nestor 1995a) and the Muhu dolomites by Luha (1930) were re-established as a formation (Nestor 1995a). V. Nestor (1984) determined the scope of the stage in subsurface area.
The historical stratotype of the stage is an abandoned quarry at Jaagarahu supplemented with the Jaagarahu drill core in the interval of 0.3 to 21.4 m (Aaloe 1970). The Jaagarahu Stage spreads on Saaremaa and Muhu islands and in the southwestern part of mainland Estonia (Pärnumaa and southern Läänemaa). The outcrop extends as a 10—30-m-wide belt from Vilsandi and Vaika islands through northern Saaremaa and Muhu as far as Eidapere and Tori at Tallinn - Pärnu railway (Fig. 71). The main localities are the quarries at Jaagarahu, Tagavere, Koguva and Anelema, and recent and ancient coastal cliffs at Vilsandi, Abula, Panga (Photo 22), Pulli (Oiu), Üügu, Püssina, Kesselaid, Salevere and Kirbla (Photo 23). The thickness of the stage is variable and increases southwestwards from 32.3 m in the Viki core to 145.0 m in the Ohesaare core.
In the western part of Saaremaa, the Jaagarahu Stage is dominated by comparatively pure limestones, while dolomites are prevailing in the eastern part of Saaremaa, on Muhu Island and in mainland Estonia. Reefs (bioherms and mounds) are widespread in the Jaagarahu Stage, especially in its lower part (Vilsandi beds and Kesselaid Member). In the South-Estonian Confacies Belt, the lower part of the Jaagarahu Stage is represented by the marlstones of the Jamaja Formation, and the upper part by the nodular biomicritic limestones of the Sõrve Formation. Temporal analogues of the latter formation are absent in northern sequences due to the long stratigraphical hiatus (Nestor & Nestor 1991). In the Mid-Estonian Confacies Belt, the lower boundary of the stage has been drawn by an abrupt increase in the carbonate content of the rocks coinciding with the base of the Jaagarahu and Muhu formations. In the more argillaceous sequences of the South-Estonian Confacies Belt and transition area, the lower boundary is determined by the appearance of chitinozoans of the Linochitina cingulata Biozone (Nestor V. 1994) at the base of the Jamaja and Riksu formations.
The Jaagarahu Stage contains a wide spectrum of fossils from lagoon-related eurypterids and thelodonts to deep-water communities of chitinozoans, ostracodes and trilobites. Severe dolomitization has destroyed skeletal remains over a vast area in the eastern part of the stage (e.g. in the Muhu Formation). The most characteristic fossils are as follows (abbreviations: jg - Jaagarahu Formation, jm - Jamaja Formation, srv - Sõrve Formation, rks - Riksu Formation, mh - Muhu Formation, V - Vilsandi beds, M - Maasi beds, u.pt. - upper part): Vikingia tenuis (Nestor) (jgV), Ecclimadictyon astrolaxum Nestor (jgM), Favosites mirandus Sokolov (jgV), Thecia confluens (Eichwald) (jgM), Coenites juniperinus Eichwald (jgV,M), Acervularia ananas (L.) (jgV), Kodonophyllum truncatum (L.) (jgM), Dolerorthis rustica (Sowerby) (jm), Howellella cuneata Rubel (jg,mh,srv), Encrinurus balticus Männil (jm), Warburgella estonica Männil (jgM), Craspedobolbina insulicola Martinsson (jm), Leptobolbina quadricuspidata Martinsson (srv), Conochitina lagena Eisenack (jm, rks), C. pachycephala Eisenack (jm, srv), C. cribrosa Nestor (srv), Ozarkodina sagitta rhenana Walliser (jgV), Kockelella walliseri Helfrich (jm), K. amsdeni Barrick et Klapper (srv), Monograptus flemingii Salter (jm, srv), Gothograptus nassa (Holm) (srv, u.pt.), Logania taiti (Stetson) (srv). Taking into account the few findings of zonal species of graptolites in the Ohesaare core, it seems that the Jaagarahu Stage probably spans from the M. flexilis Biozone (partly) to the Gothograptus nassa Biozone (Table 8). The main, lower part of the Jaagarahu Stage consists of the Jaagarahu, Muhu, Riksu and Jamaja formations laterally replacing one another (Figs. 69, 70). The upper part of the stage is represented by the Sõrve Formation which is distributed only in the South-Estonian Confacies Belt; in the northern sequences a stratigraphical cap corresponds to it.
The Jaagarahu Formation occurs in northwestern Saaremaa and consists of very variable, prevailingly sparitic limestones of shallow-water origin. Coral-stromatoporoid limestones, including reefs or bioherms, and fine-grained skeletal and pelletal grainstones are the most widespread rocks. In some places they are dolomitized. The formation contains some bands of lagoonal argillaceous dolostones, the so-called eurypterus and pattern dolomites which divide the formation into three subunits: the Vilsandi, Maasi and Tagavere beds. The Vilsandi beds comprise an abundance of large bioherms. The Maasi beds contain biomicritic interlayers of deeper-water genesis. The Tagavere beds are capped by the thickest (5-8 m) deposit of lagoonal dolomites treated sometimes as the Selgase Member. The thickness of the Jaagarahu Formation varies from 32 to 46 m (Fig. 71 ).
The Muhu Formation is distributed in northeastern Saaremaa, on Muhu Island and on mainland Estonia north of the Pärnu latitude. It consists mostly of rather monotonous flaggy dolomites containing numerous large massive reef-mounds (Fig. 70, Photo 23) in its lower part (Kesselaid Member). Almost everywhere the thickness (20 to 40 m) of the formation is uncomplete due to the post-Silurian denudation.
The Riksu Formation bounds the Jaagarahu and Muhu formations from the south, and spreads along the southern coasts of Saaremaa, the Tõstamaa Peninsula and around Pärnu Bay. It is mostly represented by nodular biomicritic to micritic limestones containing layers of argillaceous limestones and marlstones causing a cyclical nature of the sequence. In the eastern part of the distribution area, the rocks are usually dolomitized. In places (Pärnu, Kihnu, Seliste), the Riksu Formation is underlain by a tongue of the Jamaja Formation and, in most places, it is overlain by the tongue of the Jaagarahu or Muhu formations. The thickness of the Riksu Formation varies from 34.5 m in the Nässumaa borehole to 50.8 m in the Kaugatuma borehole.
The Jamaja Formation forms the lower part of the Jaagarahu Stage in the South-Estonian Confacies Belt. The formation is represented by different marl- and mudstones. The thickness of the formation reaches 95.2 m in the Ohesaare section.
The Sõrve Formation overlies the Jamaja Formation in the southernmost sections of Estonia (Ohesaare, Ruhnu, Ikla). It is represented by biomicritic to micritic nodular limestones (pack- and wackestones) similar to the Riksu Formation but lying stratigraphically higher in the sequence and corresponding to the hiatus in northern sequences. The thickness of the formation reaches 49.8 m in the Ohesaare boring.
The Rootsiküla strata were established by Bekker (1925) as the Rootsiküla-Kaarma Substage of the Saaremaa Stage (= “Obere Oeselsche Gruppe” by Schmidt 1858, 1881). Later Luha (1933) raised the unit into the stage rank and introduced the name Kaarma. Einasto (1970) motivated the use of the name Rootsiküla. He also defined the boundaries of the stage and subdivided it into beds. Viita quarry, the historical stratotype, has been destroyed. The Kipi drill core in the interval of 25.6 to 53.6 m has been treated as the hypostratotype of the stage (Einasto 1970, Nestor 1993). The Rootsiküla Stage spreads in middle and southern Saaremaa, in the western part of the Tõstamaa Peninsula and on Kihnu and Ruhnu islands. The outcrop forms a 4–10-km-wide belt running through the central part of Saaremaa from Atla to the Kübassaare Peninsula. On mainland, it reaches the Seliste Village on the Tõstamaa Peninsula (Fig. 72). The main localities are the coastal cliffs at Elda, Soeginina, Anikaitse and Kübassaare, the Vesiku Rivulet and an abandoned quarry at Pamma. In Estonia, the full thickness of the stage varies from 20 to 40 m and increases rapidly southwards, towards the Kurzeme Peninsula (Fig.72).
The Rootsiküla Stage consists of various skeletal, pelletal, lithoclastic, coquinoid and micritic limestones cyclically interbedding with argillaceous sedimentary dolomites (the so-called Eurypterus and pattern dolomites). Limestones form the lower and dolomites the upper part of the shallowing-up sedimentary cycles. Microbial-algal structures (oncolites, stromatolites) are frequent. The limestones are often dolomitized and in the eastern part of the distribution area the whole sequence consists completely of dolomites. The lower boundary of the stage has been determined at the base of a stratum of skeletal pack- or grainstones forming the lowermost part of the Viita beds which disconformably overlie the first thick stratum of lagoonal dolomites (Selgase Member of the Jaagarahu Formation) in the Middle-Estonian Confacies Belt and unnamed members of skeletal grainstones at the top of the Sõrve and Riksu formations in the South-Estonian Confacies Belt. In the Ohesaare boring these basal-Rootsiküla nodular packstones contain the Beirichia subornata ostracode fauna characteristic to the Mulde Marls on Gotland and correlatable with the Gothograptus nassa graptolite zone.
The Rootsiküla Stage contains sparsely distributed and specific fossil biota. Eurypterids, thelodonts, leperditian ostracodes, specific gastropods, bivalves, oncolites, stromatolites are common indicating the shallow, near-shore environments. The most typical species are (abbreviations: rt - Rootsiküla Formation, Vt - Viita beds, K - Kuusnõmme beds, Vs - Vesiku beds, S - Soeginina beds): Araneosustroma stelliparratum (Nestor) (rtK), Parastriatopora commutabilis Klaamann (rtK,S), Howellella cuniculi Rubel (rtVt), Straparollus (S.) helicites (Sowerby) (rtVs), Murchisonia (Hormotoma) compressa Lindström (rt), Hermannina phaseola (Hisinger) (rt), Bingeria vesikuensis Sarv (rtVt), Beirichia subornata Martinsson (rtVt), Balteurypterus remipes tetragonophtalmus (Fischer) (rtVt,Vs), Ctenognathodus murchisoni (Pander) (rt), Ozarkodina bohemica bohemica (Walliser) (rtVt), Logania martinssoni Gross (rt), Tremataspis schmidti (Rohon) (rtVt,K,Vs). The presence of Ozarkodina bohemica bohemica enables to date the Rootsiküla Stage as top Wenlock - basal Ludlow.
In Estonia, the Rootsiküla Stage is represented by the Rootsiküla and Sakla formations laterally replacing each other (Fig. 70). The Rootsiküla Formation is distributed on Saaremaa Island, except its easternmost part. The formation is represented by cyclically alternating limestones (often secondarily dolomitized) and argillaceous sedimentary dolostones. Limestones are prevailingly skeletal and pelletal grainstones containing in places oolites, oncolites and intraclasts. Biomicritic and micritic limestones also occur on some levels. Argillaceous dolomites and dolomitic marlstones form the upper part of the sedimentary cycles. They are laminated Eurypterus dolomites or massive bioturbated pattern dolomites. Four cycles have been distinguished in the Rootsiküla Formation (Einasto 1970) defined as beds. The lowermost, Viita beds begin with comparatively normal-marine biomicritic limestones and end with typical Eurypterus and pattern dolomites. The Kuusnõmme beds form a thin uncomplete cycle with coral-stromatoporoid or oncolitic limestone in the lower and pattern dolomite in the upper part. The Vesiku beds begin with sparitic limestones, pelletal, oolitic, coquinoid or lithoclastic grain- or floatstones and their dolomitized counterparts and end with the thickest band of Eurypterus and pattern dolomites. The Soeginina beds form an untypical cycle with totally dolomitized porous grainstones in the lower, thick stromatolite band in middle and pattern dolomites in the upper part. The Sakla Formation is developed in southeastern Saaremaa, on Kihnu Island, and on the Tõstamaa Peninsula in mainland Estonia. It is represented by comparatively monotonous, thick-bedded, bioturbated dolomites with numerous pyrite patterns and undefinite cyclicity.
The present unit was treated by Schmidt (1892) as the Ilionia Schichten (Beds) of the “Obere Oeselsche Gruppe”. Bekker (1925) introduced the geographical name Paadla and considered these beds as a substage of the Saaremaa Stage. Aaloe (1963b) included equivalents of the present-day Himmiste beds earlier correlated with the Kaarma Stage. Aaloe et al. (1976) redefined the upper boundary of the stage excluding the Tahula beds and gave the current stratification of the stage (Table 8).
The historical stratotype of the stage - Paadla quarry, has been destroyed. The Kuressaare-GI (Kingissepa) drill core in the interval of 19.8 to 43.4 m has been chosen as the neostratotype of the stage (Nestor 1993). The rocks of the Paadla Stage occur in middle and southern Saaremaa, on Kihnu and Ruhnu islands and in the western part of the Tõstamaa Peninsula in mainland Estonia. The outcrop forms a 12–20-km-wide belt passing through midsouthern Saaremaa from the vicinity of Karala to Kõiguste and extending eastwards as far as the Tõstamaa Settlement on mainland (Fig. 73B). The main localities are cliffs at Roopa and Katri, quarries at Lümanda, Himmiste-Kuigu, Kogula, Kaarma and Uduvere and the walls of the Kaali meteorite crater. The full thickness of the stage varies from 2.8 m in the Kihnu borehole to 28.4 m in the Kaugatuma borehole, increasing westwards (Table 9).
The Paadla Stage consists of various bioclastic, pelletal and argillaceous limestones containing coral-stromatoporoid bioherms and biostromes in the west and different primary and secondary dolomites in the east. The base of the stage coincides with the top of the Soeginina pattern dolomites of the Rootsiküla Stage, overlain by argillaceous limestones and dolomites with Didymothyris didyma and Ilionia prisca.
In Estonia, the Paadla Stage contains a rather specific shallow-water fauna of corals, stromatoporoids, agnathans, brachiopods and molluscs in the north-west, more diverse shelly fauna in the south-west, and almost barren dolomites in the east. The most typical species are as follows (abbreviations: pd - Paadla Formation, tr - Torgu Formation, kh - Kihnu Formation, S - Sauvere beds, H - Himmiste beds, U - Uduvere beds, m.pt. - middle part, u.pt. - upper part): Conochitina latifrons Eisenack (tr), Angochitina elongata Eisenack (tr), Parallelostroma typicum (Rosen) (pd), Lophiostroma schmidtii (Nicholson) (pd), Thecia swindereniana (Goldfuss) (pd), Laceripora cribrosa Eichwald (pdU), Phaulactis cyathophylloides Ryder (pd), Didymothyris didyma (Dalman) (pd, tr), Howellella elegans Muir-Wood (pd, tr), Ilionia prisca Hisinger (pd, tr), Cardiola interrupta Sowerby (pd, tr), Megalomphala taenia (Lindström) (pdU), Hemsiella hemsiensis Martinsson (pdU, tr u.pt.), Neobeirichia nutans (Kiesow) (tr), Hammariella pulchrivelata Martinsson (pdU, tr), Amygdalella paadlaensis Sarv (pd, tr), Balizoma obtusus (Angelin) (pdU), Ozarkodina crispa (Walliser) (pdU, tr u.pt.), O. cf. snajdri (Walliser) (tr m.pt.), Tremataspis mammillata Pander (pdH, kh), Phlebolepis elegans Pander (pd, tr, kh), Andreolepis hedei Gross (pdU, tr u.pt.). The presence of chitinozoan species Conochitina latifrons and C. lauensis enables to correlate the strata of the Paadla Stage in Estonia with scanicus to tauragensis (leintwardinensis) graptolite zones of middle Ludlow in Latvia which suggests that the basal Ludlow beds are probably absent (Nestor & Nestor 1991).
In Estonia, the Paadla Stage is represented by the Paadla, Torgu and Kihnu formations (Aaloe et al. 1976) laterally replacing one another (Fig. 73). The Paadla Formation occurs in the southern part of Saaremaa, except the Sõrve Peninsula. It is dominated by argillaceous biomicritic to sparitic limestones and dolomites with bands of marlstones, coral-stromatoporoid biostromes, pelletal and coquinoid (Didymothyris) limestones. The formation is subdivided into the Sauvere, Himmiste and Uduvere beds (Klaamann 1970a). The Sauvere beds are represented by nodular argillaceous bioturbated biomicritic limestones (pack- and wackestones), containing small bioherms in the west and being gradually replaced by argillaceous dolomites towards the east. The Himmiste beds are mainly represented by micro- to cryptolaminated argillaceous dolomites with the remains of eurypterids and agnathans. They also contain bands of pelletal-skeletal grainstones at the base and (less often) at the top. The well-known Kaarma building dolomite is tentatively attributed to these beds now (Einasto in Kaljo & Nestor 1990, p. 173). The Uduvere beds are represented by variable rocks of shallow-water genesis: skeletal-, pelletal- lithoclastic-, oncolitic grainstones, packstones and rudstones, interbedded with bands of marlstones, coral-stromatoporoid biostromes, etc. Rocks are partly or totally dolomitized, particularly east of the Kuressaare Town.
The Torgu Formation spreads on the Sõrve Peninsula and Ruhnu Island. It mainly consists of nodular argillaceous biomicritic limestones with rather rich shelly fauna, but corals and stromatoporoids are rare.
The Kihnu Formation is distributed on the Tõstamaa Peninsula and Kihnu Island. It is represented by monotonous platy dolomites (below) and argillaceous dolomites (above) of reduced thicknesses containing agnathans of Paadla and Kuressaare ages, respectively (Einasto et al. 1977), and consequently spanning from the Paadla to Kuressaare Stage.
The Kuressaare Stage was separated from the Kaugatuma Stage by Klaamann (Klaamann 1970a). Later on, the Tahula beds were added to the stage from among the Paadla Stage (Aaloe et al. 1976). The stratotype section is the Kuressaare-GI (Kingissepa) drill core in the interval of 1.5 to 19.8 m (Aaloe et al. 1976). The Kuressaare Stage spreads in the southernmost Saaremaa, on Ruhnu and Kihnu islands and in the southwestern part of the Tõstamaa Peninsula. The outcrop forms a 2-to-10km-wide belt along the southern coast of Saaremaa Island (Fig. 73). The rocks of the stage crop out in temporary excavations and ditches in the Town of Kuressaare and its surroundings. The full thickness of the stage varies from 5.4 m in the Ruhnu to 27.4 m in the Ohesaare borehole (Tab. 9).
The Kuressaare Stage consists of different marlstones (below) and nodular argillaceous biomicritic limestones (above), both containing interlayers of skeletal, lithoclastic and coquinoid grain-, float- and rudstones. The base of the stage coincides with a sharp increase in the clay component and appearance of the elements of a new microfossil assemblage: Pterochitina perivelata, Ozarkodina remscheidensis aff. scanica, Calcibeirichia altonodosa, Thelodus sculptilis.
The Kuressaare Stage contains a rich assemblage of shelly fossils, especially ostracodes.The brachiopod Atrypoidea prunum is extremely numerous and forms coquina banks in the upper, Kudjape beds of the stage. The most typical species are as follows (abbreviations: T - Tahula beds, K - Kudjape beds): Pterochitina perivelata (Eisenack) (T, K), Conochitina granosa Laufeld (T, K), “Parallelopora” ornata Mori (K), “Paleofavosites” moribundus Sokolov (K), Entelophyllum articulatum (Wahlenberg) (K), Tryplasma loveni (M.Edw. et Haime) (K), Atrypoidea prunum (Dalman) (T, K), Calcaribeirichia altonodosa Sarv (T, K), Plicibeirichia numerosa Sarv (K), Retisaculus sulcatus Gailite (K), Limbinariella malornata Sarv (K), Calymene flabellata Männil (K), Pulcherproetus kuressaarensis (Männil) (K), Ozarkodina remscheidensis aff. scanica (Jeppsson) (T, K), O. snajdri parasnajdri Viira et Aldridge (T, K), Thelodus sculptilis Gross (T, K). The Kuressaare Stage has been indirectly correlated with the upper part of the Ludlow Series.
In Estonia, the Kuressaare Stage is represented by the Kuressaare Formation which is subdivided into Tahula beds (below) and Kudjape beds (above) (Aaloe et al. 1976). The Tahula beds mainly consist of argillaceous or dolomitic marlstones with bands of various bio- and lithoclastic limestones. The content of the calcareous component increases northeastwards.
The Kudjape beds are represented by nodular argillaceous biomicritic limestones containing coquinoid interlayers with Atrypoidea prunum and numerous colonial rugose corals.
Twenhofel (1916) introduced the name Kaugatoma in the sense of the upper subdivision (Zone) of his Oesel Formation (=“Upper Oeselsche Gruppe” by Schmidt 1858, 1881). The present-day limits and stratification of the stage were proposed by Klaamann (1970a) who separated the Kuressaare Stage as an independent unit. The historical stratotype of the stage is the Kaugatuma Cliff supplemented by Kaugatuma-GI drill core in the interval of 0.6 to 37.2 m (Resheniya… 1987, Nestor 1993). The rocks of the Kaugatuma Stage are distributed on the southern peninsulas of Saaremaa Island, and also on Ruhnu and Abruka islands. They crop out in the northern part of the Sõrve Peninsula and on the Roomassaare, Muratsi, Vätta and Leina peninsulas (Fig. 73). The main localities are the cliffs at Kaugatuma and Lõu and the abandoned quarries at Muratsi, Väike-Rootsi and Äigu. The full thickness of the stage varies from 41.6 m in the Ruhnu borehole to 85.7 m in the Sõrve-514 borehole (Table 9).
The Kaugatuma Stage is represented by interbedded marlstones and bioclastic to coquinoid limestones displaying certain cyclicity. In the lower part of the cycle, marlstones are dominating; in the upper part the limestone interlayers become more frequent and the cycle ends with a thick (2-4 m) deposit of crinoidal limestones. In the upper cycles and southwards, the role of limestone layers decreases. The lower boundary of the stage coincides with a notable increase in the clay component. Higher in the sequence, there appear species of ostracodes Amygdalella nasuta, Sleia equestris, Frostiella groenvalliana, Neobeirichia buchiana; chitinozoans Ancyrochitina fragilis; conodonts Ozarkodina remscheidensis eosteinhornensis etc.
The Kaugatuma Stage contains a rich shelly fauna, particularly ostracodes. The guide fossils of the stage are as follows (abbreviations: Ä - Äigu beds, L - Lõo beds): Ancyrochitina fragilis Eisenack (Ä, L), Fungochitina pistilliformis (Eisenack) (L), Densastroma astroites (Rosen) (Ä), Actinostromella vaiverensis Nestor (Ä), Parallelostroma tuberculatum (Yavorsky) (Ä), Favosites pseudoforbesi muratsiensis Sokolov (Ä), Syringopora blanda Klaamann (Ä, L), Cystiphyllum cylindricum Lonsdale (Ä), Atrypoidea prunum (Dalman) (Ä), Stegerchynchus pseudobidentatus (Rybnikova) (Ä, L), Acaste dayiana Richter et Richter (Ä), Pulcherproetus nieszkowskii (Männil) (Ä), Amygdalella nasuta Martinsson (Ä, L), Sleia equestris Martinsson (Ä), Frostiella groenvalliana Martinsson (Ä), Nodibeirichia tuberculata (Klöden) (L), Crotalocrinites rugosus (Miller) (Ä, L), Ozarkodina remscheidensis eosteinhornensis (Walliser) (Ä), O. remscheidensis remscheidensis (Walliser) (L), O. confluens nasutus (Viira) (L), Thelodus admirabilis Märss (Ä), Nostolepis gracilis Gross (Ä, L). The presence of Ozarkodina remscheidensis eosteinhornensis in the lower part of the Kaugatuma Stage shows that its base roughly corresponds to the Ludlow/Pridoli boundary.
In Estonia, the Kaugatuma Stage is represened by the Kaugatuma Formation which is subdivided into the Äigu (below) and Lõo beds (above) (Fig. 73). The Äigu beds consist of two regressive sedimentary cycles, sometimes regarded as the Lower and Upper Äigu beds (Nestor 1995a). In the lower part of these cycles, marlstone layers are prevalent; in their upper part limestones dominate. Among the latter, coarse-grained crinoidal limestones are the most typical rocks, but interlayers of coquinoid or bio-lithoclastic limestones are also quite common, among these bands with Atrypoidea prunum. Both cycles are capped by a thick deposit of crinoidal limestones. The Äigu beds roughly correspond to the ostracode Frostiella groenvalliana Biozone.
The Lõo beds also consist of two sedimentary cycles of the same type but marlstones prevail throughout the whole sequence. Bioclastic to coquinoid limestones occur as thin intercalations. A thicker band of crinoidal limestones occurs at the top of the lower cycle considered sometimes as the Lower Lõo beds. The Upper Lõo beds lack crinoidal limestones at the top. The Lõo beds roughly correspond to the Nodibeirichia tuberculata Biozone.
The Ohesaare strata were originally established by Bekker (1925) as a substage of the Saaremaa Stage and were raised into the stage rank by Luha (1933). Klaamann (1970a) defined the lower boundary of the stage. Aaloe et al. (1976) distinguished the Kaavi Member. The Ohesaare Cliff is the historical stratotype of the stage. Ohesaare-2 drill core at the depth of 4.10 m has been selected as the boundary stratotype of the stage.
The Ohesaare Stage crops out in the southern part of the Sõrve Peninsula and spreads also on Ruhnu Island under the Devonian cover. The only exposures of the stage are the Ohesaare and Loode cliffs (Fig. 73). In Estonia, the upper limit of the stage is erosional and the stage does not reach its full thickness anywhere. The thickest section (33.7+ m) has been recorded in Kaavi-568 boring.
In Estonia, the Ohesaare Stage is represented by the Ohesaare Formation which mostly consists of argillaceous-dolomitic marlstones or calcareous mudstones with thin intercalations of partly to totally dolomitized bio- to lithoclastic limestones. At the base of the stage there is a rather thick (4-5 m) deposit of various thin-bedded bioclastic to micritic limestones with thin intercalations of marlstone. In the upper part of the sequence the argillaceous-dolomitic marlstones are of red colour and contain silt and sand admixture. This part of the sequence is regarded as the Kaavi Member. The lower boundary of the stage coincides with the junction between the marlstones of the Lõo beds and the platy bioclastic limestones in the basal part of the Ohesaare Stage. Above this level there appear some new elements among ostracodes (Juviella piltenensis, Nodibeirichia protuberans), chitinozoans (Urochitina sp. sp., Eisenackitina lagenicula) and vertebrates (Poracanthodes punctatus, Goniporus alatus), etc.
The Ohesaare Stage contains a rather rich shelly fauna and a diverse association of agnathans and fish remains. Most of the palaeontological records come from the Ohesaare locality and characterize the lowermost part of the stage. From the Kaavi Member (K) only vertebrate fossils have been identified up to now. The species characteristic of the whole stage include Eisenackitina lagenicula (Eisenack), Urochitina cf. simplex Eisenack, U. verrucosa Eisenack, Favosites forbesi ohesaarensis Klaamann, F. vectorius Klaamann, Fistulipora tenuilamellata (Bassler), Eridotrypa parvulipora Ulrich et Bassler, Shaleria dzwinogrodensis (Kozlowski), Collarothyris collaris (Rubel), Grammysia obliqua (McCoy), Tentaculites scalaris (Schlotheim), Lonchidium inaequale Eichwald, Calymene conspicua Schmidt, Eophacops serotinus Männil, Juviella piltenensis Gailite, Nodibeirichia protuberans (Boll), Klodenia leptosoma Martinsson, Orcofabella testata (Gailite), Ozarkodina confluens nasutus (Viira), Poracanthodes punctatus Brozen, Tylodus deltoides Rohon, Goniporus alatus (Gross) (K), Nostolepis alta Märss (K).
The presence of the conodont species Ozarkodina remscheidensis remscheidensis allows to correlate the Ohesaare Stage with the upper Pridoli.
A. Kleesment & E. Mark-Kurik
The first data about the Devonian of Estonia and fossils date from the first half of the 19th century (Engelhardt & Ulprecht 1830, Kutorga 1835, 1837). A great contribution was made by Asmuss (1856), professor of Tartu University, with his valuable collection of fish fossils from the Aruküla caves (Photo 13). Grewingk (1861, 1879) was the first to describe and correlate the Devonian strata with neighbouring areas. Eichwald (1854a) presented the first reasonably complete description of the cross bedding of the Devonian outcropping strata.
Systematic research into the Devonian was started in the first half of the 20th century. During the 1920s-1940s, the outcropping Devonian strata were described and correlated by Bekker (1924a), Orviku (1930c, 1932, 1935b, 1946, 1948), Gross (1930, 1931, 1933, 1934, 1940b, 1942), Obruchev (1931, 1933) and Bölau (1943, 1944). In general outline, the classification and correlation dating from that period are valid todate (Sorokin 1981).
In the second half of the century, the lists of the fish fossils related to Devonian strata have been essentially improved (Obruchev & Mark-Kurik 1965) and new stages differentiated (Mark 1958, 1964). Age correlation of the Devonian strata in Estonia has been revised and adjusted to the internationally acknowledged scale (Mark-Kurik 1991a, 1993c, Valiukevičius 1988, 1994).
In connection with a medium-scale geological mapping thousands of boreholes were made which enabled research into buried Devonian strata (Kajak et al. 1970, Kajak & Kajak 1983, 1986, Kõrvel et al. 1970, Väärsi et al. 1971, 1981, Vanamb et al. 1975, Kala et al. 1981a, Polivko et al. 1981, Arvisto et al. 1987). In geological mapping, the stratigraphical schemes worked out by the Commission of Stratigraphy of the Baltic and East-European Platform were taken into account (Resheniya… 1978, Rzhonsnitskaja & Kulikova 1990).
Mineralogical studies initiated by H. Viiding and carried on by A. Kleesment imparted much valuable information for subdividing and correlating the strata (Viiding 1962, 1964, 1965, 1976b, Kleesment 1976, 1977, 1984). The zonal scheme was worked out by J. Valiukevičius (1988, 1994) and Valiukevičius et al. (1985) on the basis of acanthodian scales.
The present chapter is based on the whole bulk of data available on the subject. The sources include the publications and original material of the authors of the chapter, descriptions of sections stored in the Estonian Geological Fund, the results of grain-size and mineralogical analyses, based mainly on fraction 0.1–0.05 mm. Use has also been made of H. Viiding’s unpublished results. The recent stratigraphical scheme of the Devonian in Estonia (Table 10) is based on complex studies and was accepted in 1995 by the Devonian Working Group of the Estonian Stratigraphic Commission.
A. Kleesment & E. Mark-Kurik
The Lower Devonian with a total thickness of up to 51.5 m is spread in restricted areas in southern Estonia. It is represented by three stratigraphical units of different age, separated from each other by stratigraphical gaps (Table 10, Figs. 74, 79, 80).
Liepinš was the first to acknowledge the Tilžė beds as an independent stratigraphic unit - the Lower Stoniškiai Formation (Liepinš 1955). Paškevičius (1959) determined the present stratigraphic extent of the stage and gave it the present name. Faunistically, it was determined by Karatajute-Talimaa (1962). The stratotype of the Tilžė Stage is in the interval of 1104.5–1212 m of the drill core Stoniškiai-1 in southwestern Lithuania.
The Tilžė Stage is spread in southeastern Estonia and covered with rocks of the Rēzekne Stage. It has been determined only in the Laanemetsa and Värska drill cores, but is assumed to be present also in the Väimela drill core. The sediments are absent in the Mõniste High. The total thickness of the stage varies from 2.1 to 17.7 m (Fig. 74).
The Tilžė Stage lies with a stratigraphical unconformity on the rocks of different stages of Ordovician age (Figs. 75, 76, 77). It has yielded several thelodonts of stratigraphical value: Turinia pagei (Powrie), Turinia sp., Nikolivia gutta Kar.-Tal. and N. elongata Kar.-Tal. The other fossil fishes (psammosteid heterostracans, acanthodians) are identified only on the group level (Sorokin 1981).
In the Baltic region, the Tilžė Stage is represented by the Tilžė Formation which in Estonia is composed of grey and purplish-grey horizontally bedded silt- and sandstones with interlayers of grey clay and yellowish-grey dolomite. Silt- and sandstones are predominantly strongly cemented with dolomitic or gypsum (Värska) matrix. Siltstone is often mottled, conglomeritic sandstone occurs in the basal part.
The rocks of the Tilžė Formation are quartzose or feldspatic arenites with the content of quartz 60–85%. Micaceous arenites (content of micas up to 60%) occur in some places.
The heavy fraction is dominated by the group of transparent allothigenic minerals (50–70% ). Garnet with a considerable supplement of zircon is prevailing. Tourmaline and apatite are also important (Fig. 78).
The stage was established by Liepinš (1960, 1964). The Ķemeri drill core in the interval of 461 to 547.65 m has been selected as a neostratotype for the Ķemeri Stage (Sorokin 1981). The probable Ķemeri rocks in Estonia have not revealed any fossils which makes the correlation of sections rather difficult.
In Estonia the presumable Ķemeri Stage occurs in a limited area in the southwestern part of the Republic and is identified only in Ikla, Ipiku, Tõlla and Abja drill cores where its thickness varies from 5.9 to 8.4 m (Fig. 79). It lies with a stratigraphical discordance on the Silurian rocks and is covered with deposits of the Rēzekne Stage (Figs. 76, 77).
In the Baltic Region, the Ķemeri Stage is represented by the Ķemeri Formation (Table 10). In Estonia the formation consists of light-grey and pinkish poorly sorted horisontally thin-bedded sandstone and dolomite cemented with conglomeratic sandstone in the basal part. It includes interlayers of grey clay and dolomitic marl, seldom bluish-grey stabby siltstone.
Mineralogically, the rocks of the Ķemeri Formation are predominantly quartzose arenites with the quartz content reaching 80-98%. In the heavy fraction, ilmenite and transparent allothigenic minerals dominate, accounting for 17–57% and 37–53%, respectively. Among the latter group zircon is clearly prevailing, but garnet and tourmaline are also important (Fig. 78).
The Rēzekne Stage was established by Lyarskaya (1974) with the stratotype in Akniste-5 drill core (interval 361.8 - 487.8 m) in southeastern Latvia. Earlier it was treated as the Kemeri (Liepinš 1952) or Viesite (Liepinš 1964) Formation. After destruction of Akniste-5 drill core, the interval of 427 - 446 m in Ludza-15 drill core in eastern Latvia was selected for the neostratotype of the Rēzekne Stage (Lyarskaya 1978). In Estonia, the corresponding strata were earlier treated as the Pärnu Stage (Mark & Paasikivi 1960) and the Viesite Formation (Kleesment 1966). As the Rēzekne Stage they were first mentioned in 1975 (Kleesment et al. 1975). The age of the stage was palaeontologically determined by Mark-Kurik (1991a).
The Rēzekne Stage is spread in southern Estonia and covered with the rocks of the Pärnu Stage (Figs. 75, 76, 77). The best examined section is the Mehikoorma drill core (interval 220.3 - 246.2 m, Kleesment et al. 1975). The total thickness of the stage varies from 0.7 m in the Asuküla and Kaagvere boreholes up to 51.5 m in the Kavastu borehole. The stage is at its thickest in eastern Estonia (Fig. 80).
The Rēzekne Stage is characterized by greenish-, purplish- and light-grey sandstone. In southeastern Estonia, the upper part of the section is represented by dolomitic marls. The stage lies with a stratigraphical unconformity on the different stages of Ordovician or Silurian age, in a few cases on the rocks of the Tilžė Stage (Värska, Laanemetsa, Väimela), in the Mõniste High it overlies the basement complex. The lower part of the section consists of sandstones with dolomitic matrix, often conglomeratic (Figs. 75, 76, 77).
The fossils known from the Rēzekne Stage are largely confined to the Rēzekne Formation. An equivalent of the latter unit, the Lemsi Formation contains a few unidentified fish remains (Sorokin 1981). Fossil fishes coming from the Rēzekne Formation include Schizosteus sp., Psammosteidae gen. indet., Cephalaspidida gen. indet., Kartalaspis belarussica Mark-Kurik in litt., Antiarchi gen. indet., Laliacanthus singularis Kar.-Tal., Diplacanthus kleesmentae Valiuk., Acanthodes ? sp. B Valiuk., Acanthodes? sp. C Valiuk., Ptychodictyon ancestralis Valiuk., Cheiracanthus gibbosus Valiuk., Markacanthus parallelus Valiuk., Ectopacanthus flabellatus Valiuk., Rhadinacanthus primaris Valiuk., Nostolepis sp., Pruemolepis wellsi Vieth-Schreiner, Crossopterygii gen. indet. The presence of otoliths is noteworthy.
Invertebrates comprise “Lingula” sp., Ostracoda, Glyptasmussia? sp., Gastropoda. Microfossils include a simple conodont and miospores: Retusotriletes simplex Naumova, Leiotriletes microrugosus (lbr.) Naumova, L. simplex Naumova [rz1], Stenozonotriletes conformis Naumova [rz1], Acanthotriletes perpusillus Naumova [rz1], A. parvispinosus Naumova [rz1], Archaeozonotriletes memorabilis V. Umnova [rz1], Emphanisporites rotatus McGregor [rz1], Dibolisporites cf. eifeliensis (Lanninger) McGregor [rz1], Diatomozonotriletes devonicus Naumova [rz1], Retusotriletes cf. priscus V. Umnova [rz2], Leiotriletes cf. insuetus V. Umnova [rz2], Granulatisporites cf. rudigranulatus Staplin [rz2], cf. Calamospora pannucea Richardson [rz2]. The above list shows that the miospore content differs in the lower [rz1] and upper [rz2] parts of the Rēzekne Formation. The lists of fossils are by Kleesment et al. 1975 and Valiukevičius 1994 (miospores identified on generic level are not indicated).
In Estonia, the Rēzekne Stage consists of two formations, laterally replacing each other. In eastern Estonia, the Rēzekne Formation expands as far as Lake Võrtsjärv, west of it the Lemsi Formation occurs (Sorokin 1981, Figs. 77, 80).
The interval of 220.3–246.2 m of the Mehikoorma drill core has been selected as the parastratotype of the Rēzekne Formation. The thickness of the formation varies commonly from 10 to 30 m (Fig. 80). The lower part of the section is represented by grey, purplish- and pinkish-grey loose sandstone, with interlayers of brownish-black silty platy clay in its basal part. On contact with the underlying Ordovician or Silurian carbonate rocks (3–5 m) the sandstones are strongly cemented with dolomitic matrix. In southeastern Estonia, the upper part of the Rēzekne Formation is represented by grey silty dolomitic marl up to 10 m in thickness, in other regions by an up-to-1–2-m-thick layer of grey siltstone or silty sandstone (Figs. 75, 76, 77). The sandstone of the Rēzekne Formation is fine- and very fine-grained.
Mineralogically, the sandstone of the Rēzekne Formation is predominantly feldspatic arenite with the quartz content of 70–85%. The sand component of the dolomitic marl contains 60–75% of quartz. The heavy fraction is dominated by allothigenic transparent minerals. In sandstone its share is commonly 45–60% and in dolomitic marl it forms 30–50% of the fraction. This group is dominated by garnet (50–70%), accompanied by zircon (15–25%, Fig. 78). The content of garnet is relatively high in dolomitic marls.
The stratotype of the Lemsi Formation is the interval of 69.8–85.8 m of the Kihnu drill core (Sorokin 1981). The thickness of the formation is commonly 11–20 m (Fig. 77). It mostly consists of light grey, yellowish and brownish, most rarely of purplish loose sandstone, which in the basal contact part with Silurian carbonate rocks is strongly cemented with dolomitic matrix. The upper 0.6–2.9 m of the section consist of greenish and purplish-grey siltstones or very fine-grained sandstone, often strongly cemented with dolomitic matrix. The sandstone of the Lemsi Formation is predominantly fine- and medimum grained.
Mineralogically, the sandstones of the Lemsi and Rēzekne formations are similar. Only in the sandstone of the Lemsi Formation the content of zircon is higher, particularly in the Kanaküla – Tõlla – Ipiku area where it dominates over the garnet.
A. Kleesment & E. Mark-Kurik
The Middle Devonian is the completest part of the Devonian section in Estonia with both the Eifelian (Pärnu, Narva, Aruküla) and Givetian (Burtnieki, Gauja, Amata) standard stages being present (Mark-Kurik 1993c). The total thickness of the Middle Devonian rocks reaches 400 m. The wide outcrop area contains excellent exposures (Figs. 81, 82, 84, 86, 87, 88).
The Pärnu strata were established as an independent stratigraphical unit by Orviku (1930c, 1932). The name “Pärnu” (“Pernu”) was first used by Obruchev (1933). Palaeontologically, it was distinguished as the Schizosteus heterolepis Zone by Gross (1942) and as a stage by Mark-Kurik (Mark 1958). The stratotype is the bank of the Pärnu River near the settlement of Tori. The exposures occur on the banks of the Pärnu and Navesti rivers in central and southwestern Estonia, including Oore dairy — the boundary outcrop with the Narva Stage (Fig. 81).
The Pärnu Stage is spread in southern Estonia. The outcrop forms two narrow wedgeform areas in the northwestern and northeastern parts of the distribution area. The total thickness of the stage ranges commonly from 15 to 47 m (Figs. 75, 76, 81). In the Võrtsjärv Depression, only the topmost part of the section, up to 8 m in thickness, is represented.
The Pärnu Stage is characterized by light-yellow fine-grained cross-bedded sandstone. In most of the distribution area it lies conformably on the Rēzekne Stage. The topmost layer of the Rēzekne Stage in the southeastern part of the area is represented by dolomitic marl which is overlain by sandstone of the Pärnu Stage. In the western part, the boundary between these stages is difficult to establish because of their similar composition. In the northern part of the of distribution area, the Pärnu Stage lies with a stratigraphical disconformity on the Silurian and Ordovician carbonate rocks.
The majority of the fossils of the Pärnu Stage are confined to the Tori Member. The Tamme Member has revealed gyrogonites (?) of charophyte algae and, probably, unidentified lamellibranchs and rare fish remains (Orviku 1930c). In the Tamme Member Valiukevičius (pers. comm.) has identified scales of acanthodians Cheiracanthus gibbosus Valiuk., Rhadinacanthus primaris Valiuk., Cheiracanthus brevicostatus Gross and Acanthodes? sp. D. Fossil fishes occurring in the Tori Member are: Schizosteus heterolepis (Preob.), Psammolepis toriensis (Mark-Kurik), Tartuosteus sp., Actinolepis tuberculata Ag., Homostius sp., Byssacanthus dilatatus (Eichw.), Archaeacanthus quadrisulcatus Kade, Diplacanthus kleesmentae Valiuk., Acanthodes sp. B? Valiuk., Porolepis sp., Glyptolepis sp., Osteolepididae, Dipnoi?. Invertebrates (lingulates) are extremely rare. Common is fossil flora including macroremains of Hostinella sp., and Psilophytites sp., and miospores: Periplecotriletes tortus Egorova, Emphanisporites rotatus McGregor, Retusotriletes raisae Tchib., R. devonicus Naumova, R. concinnus Kedo, R. incomptus Kedo, R. planituberculatus Kedo, Dibolisporites antiquus (Kedo) Arkh., Hymenozonotriletes marginodentatus Kedo, H. altus Kedo, H. ludzus Kedo, H. longus Arkh., Calyptosporites velatus (Eisenack) Richardson, C. tener (Tchib.) Obukh. var. concinnus Tchib., Camarozonotriletes apertus Kedo, Sinuosisporites sinuosus (V. Umnova) Arkh., Punctatisporites tortuosus (Tchib.) Arkh. (data from Sorokin 1981, Valiukevičius et al. 1986, modified according to Abukhovskaya (pers. comm.)).
In Estonia and adjacent areas, the Pärnu Stage is represented by the Pärnu Formation. In Estonia, the formation (Table 10) is divided into the Tori (below) and Tamme (above) members.
The Tori Member is dominantly represented by yellow, light-grey or purplish-grey loose cross-bedded sandstone. Strongly cemented sandstone with dolomitic matrix forms only a basal layer with a thickness of 0.03 to 2 m on the Silurian or Ordovician carbonate rocks. Commonly, it contains pebbles of the underlying sediments. The sandstone is dominantly fine-grained, in the bottommost part medium-grained sandstone is developed. The thickness of the Tori Member varies greatly and irregularly and is highest in the Tsiistre drill core. In some places it is absent (Taagepera drill core, Fig. 77).
The Tamme Member is represented by interbedding loose and dolomitic-cemented greenish-, pinkish- and purplish-grey sandstone containing thin interlayers of siltstone and clay. The complex is horizontally-bedded. Commonly, yellowish-grey sandy dolomite (dolostone) with a thickness of 0.5 to 1 m occurs in the topmost part of the section. Strongly cemented sandstones with dolomitic matrix contain irregular vugs with a diameter of 1 to 15 cm, fulfilled with loose sandstone. The sandstone of the Tamme Member is fine and very-fine grained. The thickness of the member varies irregularly from 2 m (Valga-324 borehole) to 30 m, which is the full thickness of the Pärnu Formation (Taagepera borehole, Fig. 77).
Mineralogically, the sandstone of the Pärnu Formation is quartzose and feldspatic arenite with the quartz content of 75–85%. The heavy fraction is dominated by transparent allothigenic minerals (about 50%), among which garnet with a considerable supplement of zircon is prevailing (Fig. 78). The sandstones of the Tori and Tamme members are quite similar; only in the Tamme Member the content of garnet is higher and that of zircon is lower.
As an independent stratigraphic unit the Narva Stage was distinguished and termed by Obruchev (1933). The left bank of the Narva River near Gorodenka, and the banks of the Gorodenka Brook and the Poruni River near the place where these watercourses flow into the Narva River in northeastern Estonia, make up the stratotype area of the stage. The outcrops are concentrated in northeastern Estonia, with the Narva and Sirgala quarry sections (Fig. 82) being most noteworthy.
The Narva Stage is spread in southern and eastern Estonia. The outcrop area extends as a 10–30-km-wide belt from Ruhnu to Halliku. Besides, there is a separate area in northeastern Estonia and a few isolated spots near the outcrop line. The total thickness ranges from 30 to 109 m, increasing from north to south (Fig. 82).
The lower boundary of the stage coincides with the base of a carbonate breccia or dolomitic marl, 0.2 to 10 m in thickness, which overlie sandy dolomite or sandstone of the Pärnu Stage (Figs. 75,76).
On the basis of palaeontological (acanthodians, inarticulates, brachiopodes, spores), lithological and mineralogical characteristics, the Narva Stage has been divided into three substages (Table 10) traceable from the stratotype area up to eastern Belarus (Valiukevičius et al. 1986, Kleesment et al. 1987).
The Narva oil shale quarry section, 10 km southeast of the Sirgala Settlement, has been selected for the stratotype of the lower, Vadja Substage. Its parastratotype is the outcrop on the left bank of the Narva River, 300 m downstream from the mouth of the Gorodenka Brook. The thickness of the substage in the stratotype area is about 16 m, in Estonia it varies from 10 to 31.9 m (Figs. 75, 76, 83).
The Luutsniku-451 drill core in the interval of 317–377.7 m has been selected as a stratotype for the Leivu (middle) Substage. In the stratotype area it is exposed at the Poruni and Narva rivers. The thickness of the substage in northeastern Estonia is about 5 m and it increases considerably in a southern direction reaching 60.7 m in the Luutsniku drill core (Figs. 75, 76, 83).
For the stratotype of the upper, Kernavė Substage the interval of the Ledai borehole in central Lithuania was proposed.
The Narva Stage is rich in fossil fishes, particularly its upper part - the Kernavė Substage [k] where the majority of the macroremains come from. In the Vadja Substage [v] and especially in the Leivu Substage [l] the fishes are more scarce. The fish fauna of the stage consists of Schizosteus striatus (Gross) [k], Pycnolepis splendens (Eichw.) [l?,k], Cephalaspidida, Actinolepis tuberculata Ag. [k], Holonema sp. A Mark-Kurik [v], Holonema sp. B Mark-Kurik [k], Homostius latus Asm. [k], Coccosteus cuspidatus Miller ex Ag. [k], Protitanichthys? sp.n. Mark-Kurik [v], Byssacanthus dilatatus (Eichw.), Asterolepis estonica Gross [k], Cheiracanthoides estonicus Valiuk. [v], Acanthodes? sp.C [v], Cheiracanthus crassus Valiuk. [v], Rhadinacanthus balticus Gross, Acanthodes? sp.B, Acanthodes? sp.D, Cheiracanthus brevicostatus Gross, C. longicostatus Gross, Ptychodictyon distinctum Valiuk. [l,k], P. rimosum Gross [l, k], Cheiracanthus sp. [l,k], Diplacanthus sp. [l,k], Acanthodes ? sp.A. [l,k], Cheiracanthus intricatus Valiuk. [k], Nostolepis kernavensis Valiuk. [k], Cheiracanthoides proprius Valiuk. [k], Markacanthus costulatus Valiuk. [k], Minioracanthus laevis Valiuk. [k], Ptychodictyon sulcatum Gross [k], Diplacanthus carinatus Gross [k], Acanthodii gen.n. Valiuk. [k], Archaeacanthus quadrisulcatus Kade, Haplacanthus marginalis Ag., Homocanthus gracilis (Eichw.), Thursius fischeri (Eichw.) [k], Osteolepididae, Glyptolepis quadrata Eichw. [k], Glyptolepis sp., Onychodus sp.[v,k], Dipterus serratus (Eichw.)[k], Dipnoi [v, l], Cheirolepis sp.sp., Orvikuina sp. [v, l], O. vardiaensis Gross [k].
Invertebrates are represented by lingulate brachiopods (Bicarinatina sakalana Rõõmusoks et Gravitis a.o.) which are especially numerous in the upper portion of the Leivu Substage. Ostracods (Lepertitia tartuensis Öpik var. geographica Hecker, Ostracoda inc. gen.) are less common and so are unidentified lamellibranchs in the Kernavė Substage. Conchostracans Pseudestheria pogrebovi Lutk., Trigonestheria triangularis Mir., Glyptoasmussia quadrata Mir., G. aff. willweratica Nov., Concherisma eifelense Raym., Ulugkemia sinuata Lutk., U. mesodevonica Mir., Asmussia membranacea Pacht, Praeleaia quadricarinata Lutk. are characteristic for the Vadja Substage and the lowermost part of the Leivu Substage. The lists of animal fossils are given after Sorokin (1981), Valiukevičius et al. (1986) and Valiukevičius (1994).
Microspore assemblage from the Narva Stage is restricted to the Vadja Substage. It includes Retusotriletes raisae Tchib., R. devonicus Naumova, R. concinnus Kedo, R. incomptus Kedo, R. planituberculatus Kedo, R. cf. brandtii Streel, R. fragosus Arkh., R. microsetosus Kedo, R. lanceolatus Kedo, R. luxispinus Kedo, R. engurensis Kedo, R. clivosiformis Kedo, Hymenozonotrietes cf. marginodentatus Kedo, H. altus Kedo, H. cf. echiniformis Kedo, H. ludzus Kedo, Camarozonotriletes apertus Kedo, Grandispora protea (Naumova) Moreau-Benoit, Phylotecotriletes triangulatus Tiw. et Schaarschmidt. The Vadja Substage has also revealed acritarchs, not yet identified (data from Valiukevičius et al. 1986, modified). Gyrognites (?) of charophyte algae (Sycidium) have been found from all members of the Narva Stage. The Kernavė Substage comprises poorly preserved plant macroremains.
In Estonia and adjacent regions, the Narva Stage is represented by the Narva Formation with a highly variable lithology. Its lower part consists mostly of dolomitic marl with interlayers of dolomite and dolomitic clay. Siltstone, very fine-grained and silty sandstone intercalating with interlayers of dolomitic marl and clay, forms the upper part of the Narva sequence. The sequence of the formation is divided into three members corresponding in volume to substages (Table 10, Fig. 83).
In the basal part of the lower, Vadja Member the breccia of dolomitic marl with unsorted irregular pebbles of dolomite, dolomitic marl and siltstone are common. In general this member is characterized by a thin-bedded complex of dolomitic marl, dark-grey to black silty clay and pale yellowish-grey dolomite which often includes crystalline dolomite, chalcedony, pyrite or sphalerite filled vugs. The detrital component of the rocks in the lower part of the unit is mineralogically relatively mature with the quartz content reaching 50–80%. In the upper part, the rocks are often micaceous containing quartz 20–50, feldspars 15–30 and micas 20–60%. The heavy mineral suite is prevailed by Fe hydroxides or pyrite, in some interlayers baryte is dominating. Nonopaque heavy minerals are dominated by garnet, followed by zircon (Fig. 78).
The middle, Leivu Member is prevailed by dolomitic marl. The section varies both lithologically and in thickness. Within the section four beds have been distinguished (Valiukevičius et al. 1986, Kleesment 1995). Two lower beds are thinning towards the north-east (Fig. 83). The lowermost bed of grey dolomitic marl contains a remarkable amount of silty-sandy particles with a diameter up to 1–2 mm. The next bed is a grey thin-bedded complex formed of intercalating dolomitic marl, dolomite and dolomitic clay. The third bed from the bottom comprises interlayers of grey silt- and sandstone. The topmost bed consists of reddish-brown, purplish-grey and grey mottled massive argillaceous dolomitic marl which serves as a good correlative level. Mineralogically, the detrital part of the rocks in the lowest bed belongs to the feldspatic arenites, in other beds – to the arcosic and micaceous arenites. The heavy fraction is dominated by Fe hydroxides, more rarely by pyrite or micas. The beds differ from one another by the nonopaque heavy mineral suite. In the two lower beds garnet is clearly dominating, while in the overlying beds it is accompanied by apatite, zircon and tourmaline. Significant is the presence of sphen (titanite) and corund in the two lower beds (Fig. 78).
The upper, Kernavė Member consists of brownish-red and grey loose and dolomite-cemented silty sandstone with intercalations of siltstone, dolomitic marl and clay. The complex is horisontal-, lenticular-, more rarely cross-bedded. Mineralogically, the rocks of the Kernavė Member belong to the arcosic arenites with the quartz content of 50–80%. In heavy fraction micas (30–60%), more rarely Fe hydroxides (up to 92%) are prevailing. The heavy nonopaque mineral suite is mostly dominated by apatite, followed by tourmaline, zircon and garnet (Fig. 78).
The Aruküla strata were distinguished as an independent stratigraphical unit by Gross (1940b, 1942) and transfered to the rank of regional stage by Mark-Kurik (Mark 1958). The stratotype is the river bank near the Tartu Jaani Cemetery (Photo 24); not far from it are the Aruküla caves (Photo 13). The other important localities are Kallaste, Viljandi, Tamme and Õisu (Fig. 84). The rich material collected there during more than a hundred years contains the majority of the fauna characteristic of this unit.
The Aruküla Stage is spread in southern and southeastern Estonia. The outcrop area forms a 17–55-km-wide belt extending from Ruhnu Island and Ikla Settlement in the west to Pala and Mehikoorma settlements in the east (Fig. 84). The total thickness of the stage in Estonia ranges from 66.0 to 97.2 m.
The Aruküla Stage consists of reddish-brown cross-bedded sandstone, interbedded with siltstone, clay, and dolomitic marlstone. It lies everywhere above the Narva Stage. The lower boundary of the Aruküla Stage coincides, in general, with the lower surface of the first significant uncemented reddish-brown sandstone layer above the dolomitic siltstone or marl of the Narva Stage (Figs. 75, 76). The topmost part of the Narva Stage often shows a greenish-grey siltstone layer; the overlying Aruküla sandstone is mostly inequigranular. In the Võrtsjärv Depression, this boundary is often difficult to establish (Mark & Tamme 1964).
The Aruküla Stage is rich in fossil fishes known from its all three subdivisions (Table 10). Characteristic are psammosteid heterostracans and several placoderms, both arthrodires and antiarchs. Fishes from the Viljandi [vl] and Kureküla [kr] beds are better known than those from the Tarvastu beds [tr]. The Aruküla fish fauna includes: Tartuosteus giganteus Gross, Pycnosteus palaeformis Preobr. [vl], Ganosteus artus Mark-Kurik, Psammolepis proia Mark-Kurik [vl, kr], cephalaspidids [vl], Actinolepis tuberculata Ag. [vl,kr], Holonema obrutshevi Mark [vl], Homostius latus Asm., Heterostius ingens Asm., Coccosteus grossi O.Obr. [vl,kr], Millerosteus? sp [tr], Byssacanthus dilatatus (Eichw.) [vl,kr], Asterolepis estonica Gross [vl,kr], Archaeacanthus quadrisulcatus Kade, Haplacanthus marginalis Ag., Homacanthus gracilis (Eichw.), Rhadinacanthus balticus Gross [vl], R. multisulcatus Valiuk., Diplacanthus sp. [vl], D. carinatus Gross, D. gravis Valiuk., Markacanthus costulatus Valiuk. [vl], Minioracanthus laevis Valiuk. [vl,kr], M. alius Valiuk., Acanthodes? sp. A, Acanthodes? sp. B, Acanthodes? sp. D, Cheiracanthus brevicostatus Gross, C. longicostatus Gross, Ptychdictyon rimosum Gross, P. sulcatum Gross, Gyroptychius pauli Vorob., Glyptolepis sp. sp., Dipterus sp. sp., Conchodus sp. [kr], Orvikuina sp.n., Cheirolepis sp., Tartuosteus? luhai Mark-Kurik [kr, tr], Pycnosteus pauli Mark [kr,tr], Nodocosta pauli Gross [kr], Thursius estonicus Vorob. [kr], Hybosteus sp. [kr], Tartuosteus maximus? Mark-Kurik [tr], Nostolepis sp.n. Valiuk. [tr], Ptychodictyon distinctum Valiuk. [tr], Porolepis? sp. [tr].
Invertebrates of Aruküla Age are mostly known from the Viljandi beds. These are: ostracods Leperditia tartuensis Öpik, Aparchitellina taehtverensis (Öpik), Drepanellina orvikui (Öpik), Pontocypris rulescens (Öpik), lingulates Bicarinatina bicarinata (Kut.), B. ugalana Rõõmusoks et Gravitis and unidentified lamellibranchs. Trace fossils are fairly numerous in the Kureküla beds. Lingulate brachiopods occur also in the Tarvastu beds.
Plant remains consist of gyrogonites? of charophyte algae (Viljandi beds) and some poorly preserved fragmental branches. The list of fossils is from Sorokin (1981) and Valiukevičius (1994, modified).
In Estonia, the Aruküla Stage is represented by the Aruküla Formation. On the basis of lithological and mineralogical evidence, three cyclic units are distinguished in the Aruküla Formation. These cycles occur in all sections of Estonia and adjacent areas and are defined as the Viljandi (lower), Kureküla (middle) and Tarvastu (upper) beds of the Aruküla Formation (Table 10). Each unit begins with relatively coarse and poorly sorted sandstones of a mature mineral composition, but ends with a clayey-silty complex (Figs. 75, 76, 85, Kleesment, 1994).
The lower, Viljandi beds are dominated by very fine sandstones, often platy or slaty-bedded. The Kureküla beds are characterized by irregularly cemented interbeds of variegated siltstones, pockets of white sandstone, lenses of conglomeratic sandstone and interlayers with large clay pebbles. The section of the Tarvastu beds contains typically conglomeratic interbeds and surfaces and crusts of Fe hydroxide.
Mineralogically, the rocks of the Aruküla Formation are predominantly quartzose and feldspatic arenite with the quartz content of 60–90%. Micaceous arenites (content of micas 20–50%) occur as thin interbeds. The heavy fraction is dominated by ilmenite (30–60%) and transparent allothigenic minerals (15–40%). Among the latter, garnet and zircon are most significant. Tourmaline and rutile are also important. On the lower boundary of the formation the content of zircon and apatite increases significantly, and staurolite appears (Fig. 78).
As an independent stratigraphical unit the Burtnieki strata was distinguished by Gross (1940b, 1942). Into the rank of regional stage it was raised by Mark-Kurik (Mark 1958). The stratotype is the bank of the Salaca River, 12 km northwest of Lake Burtnieki in northern Latvia. In Estonia, main exposures are situated at Helme and on the banks of the Ahja (Photo 25) and Võhandu rivers. The outcrops of Karksi, Härma, Koorküla and Essi are known as localities of fossil fishes (Fig. 86).
The Burtnieki Stage is spread in southeastern Estonia. The outcrop area forms a 25–50-km-wide belt stretching from Ipiku and Valga in the west to Mehikoorma and Karisilla in the east. The total thickness ranges from 60.6 to 94.5 m (Fig. 86).
The Burtnieki Stage is mainly represented by light (white, yellowish, pinkish and greyish-brown) fine-grained medium- to weakly-cemented cross-bedded sandstones with interlayers of siltstone and clay. The stage lies everywhere above the Tarvastu beds of the Aruküla Stage (Figs. 75, 76, 85). The topmost layer of the Tarvastu beds is, as a rule, represented by reddish or variegated (purplish-grey to reddish-brown) siltstones, which are overlain by white, yellowish-, brownish- or purplish-grey poorly sorted loose sandstones of the Burtnieki Stage.
The Burtnieki Stage is divided into the Salaca (below) and the Abava (above) substages (Table 10). Mark-Kurik (1993a, b) has treated the latter as an independant stage.
Fossils coming from different parts of the Burtnieki Stage, the Härma [hm] and Koorküla [kr] beds and the Abava [ab] Substage belong mainly to fishes: Tartuosteus maximus Mark-Kurik [hm], cephalaspidids [hm], Pycnosteus tuberculatus (Rohon) [hm,kr], Ganosteus stellatus Rohon, Psammosteus bergi (Obr.) [hm], Actinolepis magna Mark-Kurik [hm,ab], Tropinema haermae (Mark) [hm], Homostius latus Asm. [hm,kr], Heterostius ingens Asm. [hm,kr], Coccosteus markae O.Obr. [hm], Asterolepis sp.1 Kar.-Tal.[hm], Homacanthus gracilis (Eichw.) [hm], Nodocosta sp. [kr], Acanthodes? sp. A, Acanthodes? sp. B, Acanthodes? sp. D, Acanthodes sp. [ab], Cheiracanthus brevicostatus Gross [hm,ab], C. longicostatus Gross [hm], Cheiracanthus sp. [ab], Ptychodictyon rimosum Gross [hm], P. sulcatum Gross [hm], ? Ptychodictyon sp. [hm], Diplacanthus carinatus Gross [hm], D. gravis Valiuk. [hm], Acanthodii gen.n. Valiuk. [hm], Markacanthus alius Valiuk. [hm], Rhadinacanthus multisulcatus Valiuk. [hm], Nostolepis sp.n. Valiuk. [hm], Gyroptychius elgae Vorob. [hm], Glyptolepis? karksiensis (Vorob.) [hm], holoptychiids [hm], Psammolepis sp.sp., Byssacanthus sp.sp. [kr,ab], Hamodus lutkevitshi Obr. [kr,ab], Panderichthys? sp. [kr,ab], Psammolepis abavica Mark-Kurik [ab], Psammosteus sp.sp. [ab], Watsonosteus sp.n.? [ab], Livosteus? sp. [ab], Plourdosteus? panderi O. Obr. [ab], Asterolepis essica Lyarsk. [ab], Microbrachius cf. dicki Traq. [ab], Chondrichthyes? [ab], Laccognathus sp. [ab], Osteolepididae [kr, ab], Onychodus? sp. [ab], Dipnoi [ab], Moythomasia? sp. [ab], Cheirolepis sp. [ab] (Sorokin 1981, Valiukevičius 1994, modified).
Of other fossils, silicified wood has been found in the Härma beds and rare lingulates and various remains of the pteridophyte Pseudosporochnus estonicus Kalamees in the upper clayey part of the Abava Substage (Kalamees 1988).
In Estonia and adjacent areas, the Burtnieki Stage is represented by the Burtnieki Formation. On the basis of the lithological and mineralogical data, three cyclic units are distinguished in the Burtnieki Formation. These cycles are observable in all sections of Estonia and defined as the Härma (lower), Koorküla (middle) and Abava (upper) beds (Table 10). Each unit begins with relatively coarse-grained light, variegated (yellowish, pinkish, greyish and brownish) sandstones of a mature mineral composition and ends with clayey silt layers (Figs. 75, 76, 85, Kleesment 1995).
Lithologically and mineralogically (Fig.78), these three beds are rather similar. The sandstones are prevailed by fine-grained fraction which usually forms 50–70% of the rock. The rate of medium-grained and very fine-grained sand fractions is variable, forming 10–30 and 6–20% of the rock, respectively. The share of other fractions rarely exceeds 5%. The predominating thickness of the cross-bedded sandstone series is 20–30 cm. They are dipping to the south, southwest, and southeast. In the Härma beds, the southwest inclination is prevailing, while in the Koorküla and Abava beds the inclination directions are more variable. Siltstones are mostly medium-cemented, variegated, clays are strongly silty, grey, and reddish-brown.
Mineralogically, the rocks of the Burtnieki Formation are predominantly quartzose and feldspatic arenites with the quartz content of 70–90%. Micaceous arenites (content of micas up to 50%) occur as rare thin interbeds. The heavy fraction is dominated by ilmenite (45–65%). The share of allothigenic transparent minerals in the heavy fraction is in general 15–30%. This group is dominated by zircon (40–70%). Of other accessory minerals, tourmaline (7–20%) and staurolite (3–15%) are more important, noteworthy is the appearance of kyanite (Fig. 78). The share of tourmaline is greatest in the Abava beds where it makes up 10-30% of the group of transparent allothigenic minerals.
The Gauja Stage was formally established by Liepinš (1951), although the corresponding stratigraphical unit as a Stage already existed in the scheme of Kraus (1934) and had been distinguished by Gross (1942) as “Oredesch-Stufe”. In different periods it has been treated as a separate stage or as the lower part of the Šventoji Stage (Sorokin 1981). The stratotype of the Gauja Stage is the bank of the Gauja River between Cēsis and Sigulda in northern Latvia. In Estonia more important localities are the banks of the Piusa, Pärlijõgi and Mustjõe rivers, Tuhkvitsa Brook and sand quarries near the railway station at Piusa (Fig. 87).
The Gauja Stage is spread in a restricted area in southeastern Estonia. The outcrop area forms a 14–30-km-wide belt which extends from Valga and Luutsniku in the west to Karisilla and Petseri in the east. The total thickness of the stage in Estonia ranges from 78 to 79.8 m (Fig. 87).
The Gauja Stage consists mostly of weakly- to medium-cemented white and light- to yellowish-grey cross-bedded sandstones. It lies everywhere on the topmost clayey-silty complex of the Abava beds of the Burtnieki Stage (Figs. 75, 76, 85). On the contact level the sandstone is often rich in carbonate cement.
The lower and upper parts of the Gauja Stage differ in fossils. The lower, Sietiņi Member has yielded fossil fishes: Psammolepis venyukovi Obr., P. paradoxa Ag., P. heteraster Gross, P.alata Mark-Kurik, Plourdosteus livonicus (Eastm.), Asterolepis ornata Eichw. sensu Ag., Bothriolepis? sp., Glyptolepis baltica Gross, Laccognathus panderi Gross and Megadonichthys kurikae Vorob. in litt. In the Sietiņi Member also some large fragments or stems, or both, of silicified and ferriferous wood have been found (Sorokin 1981, modified).
In the Lode Member, only plant macroremains (Hostinella sp., Archaeopteris sp., A. fissilis Schalh.) and miospores are known. The miospore assemblage includes: Retusotriletes rugulatus Riegel, Samarisporites triangulatus Allen, S. eximius (Allen) Loboziak et Streel, Geminospora micromanifesta (Naumova) Arkh., G. lemurata Balme, emend. Playford, Ancyrospora sp. cf. A. incisa (Naumova) M. Rask. et Obukh., Dictyotriletes sp. cf. Reticulatisporites perlotus (Naumova) Obukh., Perotriletes sp. cf. Rugospora? impolita (Naumova) Tchib. (Blieck et al. 1996).
In Estonia and adjacent areas, the Gauja Stage is represented by the Gauja Formation. In the latter, two cyclic complexes can be distiguished, corresponding to the Sietiņ (lower) and Lode (upper) (Table 10) members, established by Kuršs (1992). The Sietiņi Member consists mostly of sandstones, with variegated siltstone in the topmost part. The lower part of the Lode Member is represented by light, mainly white sandstones, its upper part is dominated by siltstones and clays (Figs. 75, 85).
The sandstones of the Gauja Formation are fine-grained. The share of fine-grained particles is usually 55–65%, the content of very fine-grained particles is 22.5%, on an average. The cross-bedded series are 5–40, mostly 15-30 cm thick, predominately inclined to the southwest, south and southeast. Characteristic are brown iron-rich surfaces, pebbles of purplish-brown and grey clay, quartz and Fe hydroxide.
The siltstones, which form ca 20% of the section on an average, are usually clayey, represented by variegated, grey and brownish varieties. Clays (average 15%) are strongly silty, grey, and purplish-grey, often dolomitic.
Mineralogically, the rocks of the Gauja Formation are predominantly quartzose arenites with the quartz content of 80–94%. The heavy fraction is dominated by ilmenite, transparent allothigenic minerals make up 20–30%. In the latter group, the leading mineral is zircon, although in the Lode Member tourmaline often dominates (Fig. 78). The Sietiņi and Lode members differ notably in the composition of clay minerals. In the Sietiņi Member, the average share of hydromicas is 75% and kaolinite 25%. In the Lode Member, these values are 45 and 55%, respectively. The Lode Member is the most kaolinite-rich level in the Devonian sequence of Estonia.
The Amata Stage was formally established by Liepinš (1951), although the corresponding stratigraphical unit as Stage existed in the scheme of Kraus (1934) and had also been distinguished as the “Podsnetogor-Stufe” by Gross (1942). The stratotype of the stage is situated at the lower course of the Amata River in Latvia. In Estonia, the main outcrops are the banks of the Piusa (Loosi) and Peetri rivers (Karisöödi), and in the vicinity of Vastseliina (Fig. 88).
The Amata Stage is spread in a restricted area in southeastern Estonia. The outcrop area forms a 5–10-km-wide belt from Mõniste and Ape in the west to Petseri and Dekshino in the east. The total thickness in boreholes ranges from 12–21 m (Fig. 88), but in the outcrops on the banks of the Piusa River it reaches 30 m.
In Estonia, the Amata Stage is represented by sandy-silty sediments alternating with frequent clay interbeds. The stage lies everywhere on the grey clay of the Gauja Stage and starts with a layer of breccia-like sandstone (Figs. 75, 85). According to Kuršs (1992), in the lower Amata layers the cross-bedded series are inclined to the north which is not typical of this part of the Devonian.
The Amata Stage contains Psammolepis undulata (Ag.), Psammosteus praecursor Obr., P. maeandrinus Ag., Asterolepis radiata Rohon, Bothriolepis prima? Gross, B. cellulosa? Pand., Panderichthys rhombolepis Gross and Laccognathus panderi Gross (Sorokin 1981).
In Estonia and adjacent areas, the Amata Stage is represented by the Amata Formation. According to borehole data, the predominating rock type in the Amata Formation is the greenish- and purplish-grey siltstone which forms on average of 45% of the section. In outcrops, however, the sandstones are prevailing. The sandstones of the Amata Formation are light to yellowish-grey, more rarely reddish-brown, fine-grained, medium- to strongly-cemented, with indistinct cross-bedding, the inclination of which varies in wide limits. The sandstones often contain pebbles and lens-shaped interlayers of clay, more rarely quartz pebbles. Clay interlayers are usually purplish-grey and -brown and form 30% of the section as an average.
Mineralogically, the sandy-silty rocks of the Amata Formation are predominantly quartzose arenites with the quartz content being 80-90%. The heavy mineral suite is dominated by ilmenite, the share of transparent allothigenic minerals is relatively great, varying from 26 to 40%. Among this group zircon is predominating. It is followed by tourmaline, staurolite and rutile in almost equal amounts ( Fig. 78, Kleesment 1995). The assamblage of clay minerals is dominated by hydromicas with the average content of 95%.
K. Kajak
In southeastern Estonia, the Upper Devonian is represented by carbonate rocks, the thickness of which reaches 47 m in the Parmu borehole. The outcrop belt of rocks of the Upper Devonian Pļaviņas, Dubniki and Daugava stages has a complicated configuration. In the area where the clayey-sandy sediments of the Middle Devonian Gauja and Amata stages crop out, single outlier-islands of carbonate rocks are encountered, the largest being Saarlase and Loosi (Fig. 89).
The Pļaviņas Stage and the Pļaviņas Formation have been named after the exposures in the vicinity of the Town of Pļaviņas in Latvia (Liepinš 1951). Currently, these exposures are under the waters of the Pļaviņas reservoir and, therefore, the outcrops near Izborsk (Irboska) have been selected as the neostratotype for the Pļaviņas Stage.
In Estonia and adjacent areas, the Pļaviņas Stage has a thickness of 27–32 m. In the vicinity of Izborsk it is 37 m thick (Fig. 90). The lower boundary of the stage is lithologically clear — the clayey sandy deposits of the Amata Stage are overlain by carbonate rocks of the Pļaviņas Stage. Based on the palaeontological and lithological characteristics, the Pļaviņas Stage has been subdivided into the Snetnaya Gora, Pskov and Chudovo substages.
The Snetnaya Gora Substage has been named after a type section near the Snetnaya Gora Monastery in the vicinity of the Town of Pskov in Russia. In Estonia, the rocks of the substage crop out at the Peetri River upstream of Karisöödi, at the Pärli River near the Saarlase Mill, in the Rõuge Ööbikuorg, in the environs of Loosi and Vastseliina (Fig. 89). According to the borehole data, the thickness of the substage is ranges from 5.5 to 12 m, and increases from west to east (Fig. 90).
The substage is represented by rhythmically alternating yellowish- and greenish-grey micro- to cryptocrystalline argillaceous silty dolomite (MgO 16%, CaO 24%) and dolomitic marl (insoluble residue 30%, MgO 13%, CaO 19%), less frequently by clay. Dolomites and dolomitic marls contain silty interlayers. The complex is micro- and thin-laminated. Imprints of cubical salt crystals are found in dolomite. In northern regions, thin sand interlayers occur. At the Peetri River, the lower part of the section is composed of clay, and the upper part of dolomite.
The fossils, occasionally found in the section, are represented by the brachiopods Camarotoechia aldoga Nal., conchostracans Asmussia vulgaris Lutk. and the fishes Psammosteus meandrinus Ag., Ctenurella pskovensis (Obr.) and Bothriolepis cellulosa Pand., Grossilepis tuberculata (Gross), Moythomasia perforata (Gross).
The Pskov Substage has been named after the type section on the bank of the Velikaya River near Pskov in Russia. The exposures occur in the same area where those of the Snetnaya Gora Substage are situated. According to the borehole data, the Pskov Substage is 7–13 m thick, on the base of the exposures in the vicinity of the Izborsk Castle (Russia) it is about 17 m thick (Fig. 90).
The Pskov Substage is represented by grey, in the lower part by pale purplish limestone. The rate of dolomitization grows to the west. In the Karisöödi area, the lower part of the substage consists of dolomite (MgO 20%, insoluble residue 6%) with 3–10-cm-thick clay interlayers. In the east (Tsiistre, Hino, Vungi), the substage is mainly represented by thin-layered, often cavernous dolomite (MgO 20%), partly silty-argillaceous (insoluble residue 12–20%), in the upper part of the section it is calcareous in places (CaO 30-35%). On the east margin of the Haanja Heights (Tiirhanna, Parmu), the lower beds are represented by dolomite, the upper ones by limestone.
The Pskov Substage is rich in fossils. The brachiopods Anatrypa micans (Buch), Atrypa velikaya Nal., Ladogia meyendorfii (Vern.), Ripidiorhynchus pskovensis (Nal.) dominate. Calcareous algae have also been found.
The Chudovo Substage was differentiated on the basis of the exposures in the vicinity of the Town of Chudovo, Russia. In Estonia, the Quaternary cover is thick and the rocks of the Chudovo Substage are not exposed. The substage crops out near the Pskov - Riga highway. Based on the key fossils Ripidiorhynchus tschudovi (Nal.) and Anatrypa heckeri Nal., the age of the substage has been established in the Izborsk outcrops in Russia. The thickness of the substage in the boreholes reaches 13 m. In places (Laura, Vungi, Parmu), the lower boundary of the substage is marked by a pyritized discontinuity surface.
In the easternmost part of its distribution area (Vungi, Parmu), the Chudovo Substage is represented by micro- and cryptocrystalline limestones (CaO 44–49%, insoluble residue 5–10%). Dolomitization of rocks increases westwards and the substage consists of micro- and fine-crystalline dolomitic limestones (CaO 32–34%, MgO 16–17%, insoluble residue 3–7%) to coarse-crystalline cavernous dolomites (CaO 28–29%, MgO 20%, insoluble residue 3–7%). Dolomitic facies is spread west of Misso.
The Pskov and Chudovo substages are lithologically very similar and sometimes it is expedient to treat them together as the Izborsk Member (Table 10).
The Dubniki Stage and the Dubniki Formation have been named after the former gypsum quarry which is situated east of Izborsk Town in Russia (Bekker 1924a). In the walls of the quarry up to 12.5 m thick bed of greenish-grey clay with gypsum and dolomite interlayers was exposed. Fossils are represented by the brachiopods Comiotoechia bifera (Phill.), Ripidiorhynchus strugi (Nal.) and the ostracod Acratia benevaensis Zasp.
The Formation covers a limited area in the southeastern-most part of Estonia, but the exposures do not occur there because of the thick Quaternary cover (Fig. 89). The thickness of the stage reaches 10 m (Fig. 90) in the boreholes . The section consists of bluish-grey marl (CaO 33%, MgO 3% and insoluble residue 27%) and argillaceous dolomitic marl (CaO 7%, MgO 6%, insoluble residue 58%) with clay and dolomite interlayers.
The Daugava Stage and the Daugava Formation have derived the name from the exposures on the banks of the Daugava River, Latvia (Liepinš 1951). In Estonia, the uppermost 8.5 m of the Devonian section in the Parmu borehole belong to the Daugava Stage (Fig. 90). In the borehole, the stage is represented by argillaceous micro- and cryptocrystalline limestones (CaO 49%, MgO 2%, insoluble residue 27%).
A. Raukas & K. Kajak
Estonia belongs to the zone of glacial erosion or moderate accumulation and, therefore, the Quaternary cover is rather thin. In northern Estonia, on the outcrops of the Ordovician and Silurian carbonate rocks it is usually less than 5 metres. Occasionally, on the so-called alvars, it is even lacking (Photo 26). The Quaternary cover is at its thickest (Fig. 91) in the Haanja and Otepää heights (often more than 100 m) and in the buried valleys of southern Estonia (at Keskküla 207 m). Thick Quaternary deposits are also encountered on lee-sides of heights and elevations (e.g. the Saadjärve Drumlin Field in the “shadow” of the Pandivere Upland) and in front of escarpments oriented against the movement of glaciers which favoured the accumulation of a considerable amount of sediments. In the fore-klint area, the deposits are more than 100 m and reach 143 m in the ancient Harku Valley (in the vicinity of Tallinn). In the Gulf of Finland, the greatest thicknesses have been established in megadrumlins to the north of Tallinn, for example, on the Island of Prangli, where the Pleistocene deposits are up to 123 m thick. Here the Quaternary cover comprises several till layers (Kajak 1961, 1965a, 1995, Raukas 1978).
About 95% of the Pleistocene cover is formed of glacial and aqueoglacial deposits. Glacial sediments, 70% by volume and 47.7% by surface area, dominate (Raukas 1978). Of wide distribution are also glaciolacustrine (6.8% by area) and glaciofluvial (3.1%) deposits (Fig. 92). Five till beds, often of great thickness, are to be noticed more or less distinctly. Only in a few cases they are separated from one another by deposits containing spores and pollen of interglacial or interstadial origin which considerably aggrevates the correlation and dating of the glacial strata.
In terms of glacial stratigraphy and lithogenesis, the bedrock valleys and interlobate “insular heights” are the main objects of interest. Deep ancient river valleys, which were further eroded by glaciers and their meltwaters, vary in morphology and sediment facies infill. They may be filled with glaciofluvial deposits of the last glaciation (Pada Valley), glaciolacustrine deposits of the last glaciation (Selja Valley) or one till bed with under- or overlying glacio-aquatic deposits (Kunda Valley). There are also many valleys with a complicated structure comprising several till beds and accompanying glacio-aquatic deposits. Valleys of the first three types are characteristic of northern Estonia, whereas valleys with a complicated structure prevail in the southern part of the Republic (Tavast & Raukas 1982). The deposits of radial and marginal valleys also have certain areal differences. The deposits of radial valleys have to a large extent been reworked by glaciers and contain less older deposits. With regard to their age, the Upper Pleistocene and Holocene deposits predominate in the buried valleys of northern Estonia and Middle Pleistocene deposits in the valleys of southern Estonia (Raukas & Tavast 1987). The deposits of ancient valleys are the most suitable objects for stratigraphical studies because they contain less erratics than uplands. In the latter, the blocks of bedrock and older Quaternary sediments have been displaced not only horizontally, but occasionally also a considerable up-thrusting or folding, or both have taken place. Older beds are thus found standing more or less vertically in a position tens and even hundreds of metres above their normal stratigraphic position (Raukas & Gaigalas 1993).
The so-called “accumulative insular heights” (the name is derived from the “island-like” position in the topography) form three belts with a N-S orientation in the East-European Lowland (Aboltinš et al. 1989). Estonian insular heights- Haanja and Otepää - belong to the Latgale Zone. These heights are characterized by a hummocky topography and by a considerable thickness of Quaternary deposits (up to several hundred metres). They have a rather great altitude, distinct slopes, plateau-like forms of glaciolacustrine origin in water-divide areas, and usually a bedrock core. During their development they have undergone the following four morphogenetical stages: (1) subglacial, (2) englacial, (3) marginal accumulation, and (4) a dead ice stage (Aboltinš et al. 1989). The heights have formed between rather large ice lobes as a result of frequent redeposition of older deposits, accounting for the mosaic pattern of sediments. Representative outcrops on heights have served as the main areas of stratigraphic investigations for over a century. As a result of redeposition of interglacial and interstadial sediments, the number of supposed interglacials could be erroneously increased. This, in turn, may lead to an older age being assigned to the tills separating them and to misleading palaeogeographical conclusions (cf. e.g. discussion in Liivrand 1991). A precise correlation of Pleistocene deposits assumes great knowledge of glaciosedimentation processes and elaboration of special dating methods.
The classification of the Estonian Quaternary deposits is based on the study of genetical types of deposits resulting from the development of a certain dynamic form of accumulation, and playing qualitatively different role in the structure and history of the formation of the sedimentary cover (Raukas 1978). Genetical types can be merged into paragenetical series, groups and subgroups, and, at need, into smaller taxonomical units (facies and subfacies).
Among the Estonian Pleistocene deposits six paragenetical series occur: eluvial, organogenous, colluvial and deluvial, aqueous, glacial and subaerial.
In the paragenetical series of eluvial deposits, areal and linear (along the crevasses and zones of tectonic faults) crusts of weathering, soil horizons on the boundaries of stadial or phasial beds of different age, and the deposits and relief forms of permafrost (cryogenous eluvium) occur in some places. Cryogenous phenomena (occurrences of cryoturbation, ice wedges, bedding disturbances and structural grounds) are most widespread in the Younger Dryas sediments (Photo 27).
Deposits of organogenous paragenetical series occur as interglacial peat and gyttja in few places (Rõngu, Karuküla, etc.). Organogenous deposits did not accumulate in late-glacial times and the thin interlayers of peat are most probably redeposited.
Colluvial and deluvial deposits are relatively frequent in front of Ordovician and Silurian escarpments and in the hilly topography of southeastern Estonia where in late-glacial times solifluctional processes resulting from the melting of buried ice played a significant role.
Deposits of aqueous paragenetical series are represented by several genetical types of different lithological composition. Among those, proluvial and subsurface aquatic deposits are rare. Little is also known about the interglacial and interstadial alluvial deposits represented by intermorainic silty-sandy sediments (Valguta, Peedu) and sandy gravelly deposits in the ancient buried valleys. Late-glacial alluvial deposits of the last glaciation occur in the terraces of a great many river valleys in southern Estonia where their thickness usually ranges between 3 and 10 m.
Interglacial and interstadial lacustrine deposits are represented by gyttja (Karuküla, Rõngu) and silty-sandy sediments (Otepää, Sudiste). Late-glacial terrigenous lacustrine sediments of the last glaciation are to be found under many contemporary bogs. The boundary between the Pleistocene and Holocene lacustrine deposits is rather clear and easy to notice due to an abrupt increase of organic matter in the Holocene deposits or carbonates in the form of lacustrine lime.
Marine interglacial and interstadial deposits are distributed in the fore-klint area (Prangli Island) and on the islands (Kihnu) of the Gulf of Liivi (Riga). The deposits of the Baltic Ice Lake, occurring as bottom and coastal formations, are conventionally also regarded as marine deposits.
As mentioned above, deposits of glacial paragenetical series form a great part of the Pleistocene cover in the Republic (Fig. 2). All the deposits of the glacial paragenetic series can be divided into two paragenetic groups: glacial drift deposited by glaciers on ground (subaerial tills) and underneath ice shelves (subaqueous tills). Among the subaerial varieties the lodgement, superglacial (ablation) and frontal (margin) tills and among the subaqueous varieties iceberg and submarine ablation tills can be distinguished as genetic types (Raukas 1978). On grounds of detailed micropalaeontological, geochemical, geomorphological, structural and other studies, it has been proved that the tills in Estonia are mainly of continental subaerial genesis (Raukas 1973).
Marine microorganisms (diatoms, foraminifers, ostracodes, etc.) are very rare in Estonian tills. Microfauna and -flora has been discovered in greater amount only in a narrow fore-klint area which in the past, and most likely during the interglacial stages as well, was occupied by the sea or big glaciolacustrine basins. The borehole at Vääna-Jõesuu provides an excellent illustration of the above. However, also there the content of foraminifers and diatoms in till is remarkably lower than in intermorainic silty clays (Raukas & Liivrand 1971). The occurrence of foraminifers and diatoms predominantly in the intermorainic layer evidences of the fact that the microfauna and -flora, at least partially, had redeposited here from the bottom of the Gulf of Finland where marine conditions were of repeated existence in the past.
Geochemical and most of the lithological evidence support the theory of subaerial genesis of Estonian glacial deposits. For instance, this is indicated by the great similarity between the lithological composition of tills and underlying rocks, poorly sorted sediments, lack of new authigenous formations of marine origin, poor rounding of clasts and the increase in roundness towards the south and south-east, i.e. in the direction of the supposable movement of continental ice, but also the distribution of indicator (index) boulders, a relatively high content of clayey particles and a high degree of compaction in tills, the orientation of clasts in the direction of advance of ice masses, parallel to glacial striae and almost horizontal in position, the occurrence of glaciodislocations (Photo 28) etc. (Raukas 1973).
Lodgement tills are of the widest distribution. The formation of subaerial tills is thought to have taken place both beneath an advancing glacier (lodgement till) and as a result of bottom melting of a passive glacier during the degradation of glacial covers (basal melt-out till). In Estonia, deformation tills are quite frequent. They are formed by subsole drag underneath the moving glacier (Raukas 1978).
Due to the flatness of the topography, the content of supraglacial material in tills is quite low in Estonia. By contrast, the importance of englacial and basal debris is much greater and quite different. In most cases, the material of basal debris predominates, whereas in places, lodgement tills composed of the material of englacial debris (the so-called erratic tills) are also found. Besides, local tills consisting entirely of local sedimentary material are spread (Photo 29). The vast majority of lodgement tills belong to the intermediate group between local and erratic varieties containing local and allochthonous (far-transported) material in different ratios. According to Gaigalas (1969), it would be expedient to call them transitional lodgement tills. In transitional lodgement tills the local sedimentary material is prevailing.
Ablation tills resulting from surficial melting and gravity flowing of superglacial and englacial debris are represented by flow tills and melt-out tills. Usually they are difficult to distinguish from lodgement tills. Exceptionally, lodgement tills are immediately overlain by ablation tills.
Frontal (margin) tills are present in end moraines which fall into push and dump moraines. Among dump moraines stationary and recessional forms are distinguished. There occur also end moraines of complicated structure which bear traces of ice pressure and are overlain by glaciofluvial deposits, or vice versa, as well as interior peripheral moraines of push character developed between dead and active ice (Raukas et al. 1971). Frontal tills consist of various squeeze lodgement, deformation and flow-till facies with injections of aqueoglacial deposits and bedrock erratics.
The formation of lithological and mineral composition of tills depends on a number of factors, such as the composition and topography of the underlying bedrock, the dynamics of the movement of glacier, the location of material in the body of the glacier, the character of accumulation, the nature and intensity of weathering of material, etc.
Numerous investigations have enabled to elucidate that on its way the glacier accumulated in its body great quantities of local bedrock material (Raukas 1969, a.o.). At that, maximum content of rock particles (about 60-80% ) from a certain stratigraphic unit is usually traceable near the distal boundary of the outcrop of the unit. Already at a distance of 6-8 km from the bedrock boundary, the amount of rock particles from the corresponding unit does not exceed 20-30% of the total (Raukas 1978). The content of erratic material in a typical lodgement till does not usually exceed 5-10%, but in englacially formed till it amounts to 100%. The transport distance of clasts is greatly dependent upon the resistance of rocks. Claystones and weakly cemented sandstones have travelled no more than 15 km, fine-grained limestones 120 km, dolomites 300 km and resistant varieties of crystalline rocks 800-900 km. During transport the till becomes enriched with more resistant clasts. For example, as crystalline rocks are more durable than carbonaceous ones, the South-Estonian tills on sandstones are enriched with crystalline clasts. At the same time, the fragments of carbonaceous rocks become enriched with dolomites as the tougher ones (Raukas 1978).
The local bedrock exerts also a remarkable influence on the mineral composition of tills (Raukas 1974, a.o.). For example, territorial differences are easy to trace on the outcrops of Cambrian clays and siltstones, Ordovician and Silurian carbonaceous rocks and Devonian sandstones. However, even smaller dependences can be traced from territorial aspect and with respect to the composition of different minerals. Thus, the content of carbonaceous minerals decreases abruptly and that of quartz increases southwards from the outcrops of carbonate rocks. The content of amphiboles, pyroxenes and other minerals of heavy subfraction, typical of the outcrops of Precambrian rocks of Finland, gradually decreases in the southern and southeastern direction. Correspondingly, the quantity of weathering-resistant minerals typical of the underlying Palaeozoic rocks, such as garnet, zircon, tourmaline, rutile, etc., increases. Great variations in the proportions of these minerals occur depending upon local conditions. In favour of the above speaks the content of garnet and zircon in the tills of southern Estonia where the underlying Devonian rocks display distinct regularities with respect to those minerals (Raukas 1974).
During the various glaciations, the movements of glaciers have differed (Tavast & Raukas 1982). This enables correlation of till beds on the basis of lithological (Table 11) and mineralogical (Table 12) data.
Of lithological methods, most promising for solving the problems of the Pleistocene stratigraphy and palaeogeography seems the study of crystalline indicator (index) boulders, the content of which in deposits was only slightly influenced by the differences of the local bedrock, and has remained almost stable over vast areas (Raukas 1963b).
The paragenetical subgroup of glaciofluvial deposits is divided into englacial and periglacial genetical types with frequent transitions between them. The deposits of radial eskers and fluviokames are conventionally regarded as englacial glaciofluvial deposits. The deposits of glaciofluvial deltas (Photo 30), sandurs and marginal eskers are identified as periglacial. Glaciofluvial deposits are mostly characterized by a highly variable grain-size composition and structure, and the great variation in lithological and mineralogical composition, everywhere closely connected with the composition of the adjacent till and the bedrock (Raukas 1978).
The maximum distance of transport of pebbles in glaciofluvial streams extends to 16-20 km, being naturally controlled by a great many additional factors, e.g. the hardness of rocks, the width of the outcrop of regional stages, the bedrock topography, etc. (Raukas et al. 1971). In the course of the formation of glaciofluvial deposits, the content of resistant rocks and minerals increases on account of less stable fractions that are crushed or destroyed during their transportation by water streams. Usually the content of crystalline rocks in gravel and pebble fractions of glaciofluvial deposits is 10-15% higher than in tills, whereas the content of carbonate rocks in them is accordingly lower. The content of metamorphic (predominantly gneisses) and coarse-grained magmatic rocks (predominantly rapakivi and pegmatites) is 5-10% lower than in the initial tills (Raukas 1978).
Glaciolacustrine deposits are also divided into englacial and periglacial genetical types. Englacial deposits form plateau-like limnoglacial kames (Raukas et al. 1971) and are included in superposed kames (Kajak 1963). They are of the widest distribution in northeastern Estonia, in the vicinity of Iisaku-Illuka, while periglacial deposits are most common in the Otepää and Haanja heights.
Periglacial glaciolacustrine deposits, predominantly silts and varved clays (Photo 31), are more frequent in western Estonia, in the fore-klint area and on the Narva Lowland (Fig. 2), but also in the numerous river valleys of southern and northern Estonia (Pirrus 1968). The thickness of varved clays reaches 26-27 m. The similar mineral composition of tills and varved clays, and the high content of weathering-resistant minerals in clays points to the insignificant role of chemical changes in the transformation of the initial morainic material into glaciolacustrine clay (Pirrus & Raukas 1963).
Deposits of subaerial paragenetical series are located along ancient shorelines of the Baltic Ice Lake and the larger local ice lakes where they form coastal dunes, up to 12 m high, or in the form of a thin mantle covering the sandy beach ridges. More seldom, they are represented by hillocky plains. Fine-grained sand prevails in the deposits.
Several local and regional stratigraphical schemes have been compiled for Estonia (in 1956, 1957, 1961, 1963, 1970, 1976). These were mainly correlative parts of the schemes of the European portion of the former Soviet Union or the Baltic States and Belarus (Orviku 1956, 1960d, e, Raukas 1978). In the scheme compiled by Kajak et al. (1976) local geographical names were for the first time used to denote stratigraphic units. Over a period of more than 15 years, the scheme served as a basis for large-scale geological mapping and applied works in the Republic.
On May 6, 1993, a new official stratigraphical chart of Quaternary deposits of Estonia (Table 13) was accepted by the Estonian Stratigraphic Commission (Raukas & Kajak 1995). The scheme was approved as a correlative part of the stratigraphical chart of the Baltic States at the II Stratigraphic Conference in Vilnius (May 9-14, 1993).
In the Quaternary stratigraphy, the age of tills is of special interest as it enables correlation of lithologically similar formations over vast areas (Raukas 1978). The age of tills is generally determined by bedding conditions, by their position with respect to interglacial or interstadial deposits. Unfortunately, the latter are rather uncommon. Besides, the advancing glaciers of the succeeding glaciations crushed most of unconsolidated intermorainic organic deposits which today are often found as erratics embedded in younger sediments. The deposits of the Prangli (Eemian, Mikulinan) interglacial, both continental (Rõngu) and marine (Prangli), serve as key sediments in stratigraphic subdivision and correlation of the Pleistocene cover. The Karuküla (Holsteinian, Likhvian) deposits are more complete in the Karuküla section in southwestern Estonia. The spore and pollen spectrum of all other intermorainic sections is not clear, as these sediments often contain reworked pollen. The most important type sites are shown in Figure 93.
In the Estonian stratigraphical chart (Table 13), lithostratigraphic terms have been used as basic units. As a fundamental unit, formation is used in a meaning of glacial and interglacial episodes in the event stratigraphy. Formations are the three-dimensional sedimentary bodies which have been formed by a specific geological process in the time span of one clear geological event. Big stadial episodes in a meaning of event stratigraphy are comparable with subformations. Using close in the meaning but not synonymous chrono- stratigraphical (e.g. the Prangli Stage), climatostratigraphical (Prangli Interglacial), lithostratigraphical (Prangli Formation) and event stratigraphical (Prangli Interglacial Episode) terms has been avoid. Although interglacial sediments are differentiated on the basis of spore and pollen and other fossil evidence, and pollen assemblage zones underlie their description, for the unity of the scheme, even here lithostratigraphical terms (Prangli and Karuküla formations) were preferred.
Lower Pleistocene deposits are absent in Estonia, and the oldest Middle Pleistocene deposits identified so far in the official chart belong to the Upper Sangaste Subformation. Taking into account some similarity of pollen spectrum of sandy clayey sediments in the Otepää buried valley (Harimägi borehole No. 323 at a depth of 143.3-169.2 m) with the Turgeliai and Belovežje subformations in Lithuania and Belarus (Kajak & Liivrand 1967), Kajak (1995) includes those beds to the Middle Sangaste Subformation and underlying tills and glaciofluvial deposits (Otepää borehole No. 2 at the depth 123.7-173.7 m) to the Lower Sangaste Subformation. Liivrand (1991) includes all the mentioned sediments to the Järva Formation.
The Sangaste Formation is correlated with the Dainava Stage in the southern Baltic, the Oka Stage in the European part of Russia and with the Elsterian Stage in Western Europe.
The lowest diamicton unit in Estonia termed as the U p p e r S a n g a s t e S u b f o r m a t i o n (named after the Sangaste Parish north-east of the Town of Valga) till, is very compact, brownish, sometimes greenish in colour with indications of shearing. It rests directly upon the bedrock and is found only in the bottommost part of ancient valleys. The thickness of the till bed is small - 15 m at Puiestee, 10.7 m at Sudiste and 5.4 m at Mägiste. Borehole 177 (Puiestee) at a depth of 169.0–207.0 m was chosen for the stratotype section (Raukas et al. 1993). The clast composition is different: in southwestern Estonia crystalline rocks are clearly prevailing (up to 95%), but in southeastern Estonia their amount is only 25–60%.
The high content of Vyborg rapakivi and Suursaari quartz porphyries in southeastern Estonia and the absence or a very low content of rapakivi from southwestern Finland suggest that the deposition of till was due to the southward flowing ice. The poorly sorted diamicton is richer in clay particles than the uppermost till units. The latter are high in kaolinite (up to 30–35%) derived from the weathered bedrock. Due to the influence of Devonian sand- and siltstones, the sand and silt fractions of till are richer in quartz and contain less feldspars and carbonates than other till units (Raukas 1978).
According to bedding conditions (Kajak 1995), to the Upper Sangaste Subformation belong grey and brownish till beds and glaciofluvial deposits below organogenous bog and lacustrine deposits in the Karuküla and Kõrveküla sections, up to 23 m in thickness.
The Karuküla Formation (interglacial) is palynologically correlated with the Butenai Stage (interglacial) in the southern Baltic, the Likhvinan Stage in the European part of Russia and the Holsteinian of Western Europe.
The Karuküla type site is situated in southwestern Estonia, in the Pärnu County, about seven kilometres south of the Town of Kilingi-Nõmme (Fig. 93). It displays continental deposits and was first described by Orviku (1944). The name of the stratotype proposed by Kajak et al. (1976) is inaccurate because the section is actually situated in the Keskküla Village. Due to the rather long history of investigations and wide recognition of the site, changing of the stratotype’s name was considered unpurposeful.
The Middle Pleistocene (Likhvinan, Holsteinian) age of the section was first suggested by Danilans (1966) and Voznyachuk (1966) and later established by Liivrand (1984, Velichkevich & Liivrand 1976, 1984).
The information currently available on the Karuküla site has been derived through the study of about 70 boreholes and excavations. The interglacial deposits are probably of allochtonous bedding (Levkov & Liivrand 1988). There seem to be three large and two small erratics and two lumps of Holsteinian deposits within one stratigraphic level measuring 105 m horizontally and 3.25 m vertically (Liivrand 1991). The Karuküla section and its palaeobotanical characteristics have been described in detail in several publications (Liivrand 1972, 1984, 1990, 1991). Another well investigated site of the Karuküla Formation is at Kõrveküla near Tartu (Liivrand & Saarse 1983).
The Ugandi Formation, called after an ancient South-Estonian and North-Latvian area, where those deposits are most widespread, is correlated with the Žeiminiai Formation in Lithuania, Kurzeme in Latvia, Middle Russian in Russia and Saale in Western Europe. In some places Middle Ugandi interstadial beds have been described. Borehole 6 on Prangli Island (depth 75.5-123.0 m) and borehole 268 at Valguta (13.1-35.0 m) have been established as the unit and boundary stratotypes for northern and southern Estonia, respectively (Raukas et al. 1993).
The till of the L o w e r U g a n d i S u b f o r m a t i o n, which is correlated with the Dniepr till in Russia and the Žemaitia till in Lithuania, is reddish-brown both in northern (Prangli, Naissaar, Suurpea) and southern Estonia (Mägiste, Lanksaare, Sudiste) and up to 50 m thick (Mägiste). The till is compact and occurs in uplands mainly in sheltered position or rests in ancient valleys. The clast lithology (high content of Vyborg rapakivi in eastern Estonia) indicates that the Lower Ugandi till was deposited by the southward flowing ice (Raukas 1978). Clasts in northern Estonia are almost completely fragments of crystalline rocks, whereas in southern Estonia their composition reflects both the local provenance (up to 10% of local Devonian sand- and siltstones) and the influence of the outcropping carbonaceous rocks on the way of the moving ice (50-60% carbonaceous clasts). Of all Estonian tills, it has the highest content of sandy fraction. In clay fraction illite (50-70%) prevails, but also the content of kaolinite is rather high (20-45%).
M i d d l e U g a n d i sands, silts, loams and sandy loams often contain rebedded pollen and their stratigraphic position is not clear (Liivrand 1991).
U p p e r U g a n d i till is massive to slightly stratified, reddish-brown in the fore-klint area (Prangli, Juminda) and grey in northern (Sõrve, Saadjärv) and southern (Rõngu, Suur-Munamägi) Estonia. The till unit is up to 70 m thick (Suur-Munamägi) and often cemented with carbonates. According to its composition (absence of Vyborg rapakivi and quartz-porphyries from the Island of Suursaari), the till entrained by southeastward flowing ice. Practically all the clasts in the fore-klint area originate in the crystalline basement, but in other areas carbonate rocks prevail (65-80%). In southern Estonia, this till has the highest content of silt particles and the lowest content of Devonian material. Illite (65-80%) prevails and the content of kaolinite (10-20%) is low in the clay fraction.
The aqueoglacial deposits of the Lower (Puiestee 60 m) and Upper (Vääna-Jõesuu 60 m) Ugandi subformations are rather thick and variable in composition (Raukas 1978).
The Upper Pleistocene in Estonia begins with the well-known Eemian interglacial deposits in Western Europe and Mikulinan in Eastern Europe. In the Regional Scheme of the Baltic area this interglacial is called the Merkine Interglacial after the town in southeastern Lithuania. The Eemian (Mikulinan) deposits, both continental (Rõngu) and marine (Prangli), correlated on the basis of the pollen assemblage zones, are in good stratigraphical agreement (Liivrand 1991).
The continental Eemian deposits at Rõngu were investigated in particular detail about half a century ago (Orviku 1939, Thomson 1939a, 1941). Later, complementary investigations were carried out in several other sections (Küti, Kitse, Peedu) by Liivrand (1977) and Kajak (1995).
In the sixties, marine Eemian deposits were found on Prangli Island in the Gulf of Finland (Kajak 1961) and subject to palynological (Liivrand & Valt 1966, Liivrand 1974, 1987, 1990, 1991) and diatom (Cheremisinova 1961) studies. Later, marine deposits of the Prangli Formation were found in several places (Põhja-Lehtju, Väike Tütarsaar, Kihnu a.o.).
A stratotype section at a depth of 67.6-75.5 m in borehole 6 on Prangli Island and a parastratotype for the continental deposits in borehole 264 (2.3-7.8 m) and excavation II (2.0-5.8 m) on the lands of the Vaeva Farm, 2 km west of Rõngu, were established for the Prangli Formation (Raukas et al. 1993). The name for the formation was proposed by Kajak et al. (1976).
The name of the formation was proposed by Kajak et al. in 1976 after the Järva County in central Estonia where a typical grey till of the last glaciation is widespread in the drumlins and lowland near the Town of Paide. The Järva Formation is correlated with the Nemunas Formation in Lithuania, the Baltia in Latvia, the Valdaian Stage in Russia and the Weichselian in Western Europe. The Vääna-Jõesuu (13-70 m) and Kitse boreholes (0-31.1 m) were chosen for stratotype sections in northern and southern Estonia, respectively (Raukas et al. 1993).
The Kelnase beds were named after the village on Prangli Island. In the Prangli section, they are represented by clayey silts with the pollen spectra characterized by an increasing quantity of Betula nana (40-80%) and herbs (tundra species). Gramineae and Cyperaceae are common. Selaginella selaginoides, Lycopodium alpinum and Artemisia arctica are present. A cryophilous and hydrophilous vegetation refers to the approaching glacial advance (Kajak et al. 1976, Liivrand 1991).
The Valgjärve beds, named after the lake in southern Estonia, are represented by grey till in northern and purplish-grey till and related aqueoglacial deposits in southern Estonia. The purplish-grey till was proposed for a specific stratigraphical unit by Orviku (1939) and described lithologically by Orviku (1958a) and Raukas (1963a, 1978). In the Kitse borehole No. 19 near Lake Valgjärv at a depth of 4.2-31.1 m, the till of the Valgjärve bed covers the organogenous deposits of the Prangli (Rõngu) Formation (Kajak 1995).
The Savala beds named after the village in northeastern Estonia belong to the M i d d l e J ä r v a S u b f o r m a t i o n (Kajak et al. 1976). The type section (borehole 7854, depth 25.8-30.2 m) is situated in the Savala ancient valley about 120 km east of Tallinn. It is mainly filled with grey-coloured varved clays. The pollen and spore composition of the intermorainic layer suggests dry periglacial conditions (Liivrand 1985, 1991). The Savala interstadial warming was not accompanied by any substantial development of forests.
The Võrtsjärve beds, named after Lake Võrtsjärv, are represented mainly by tills of different colour of the last glaciation and aqueoglacial deposits above and beneath the till. In several places some till layers with thin intermorainic interstadial or interphasial sediments occur (Orviku 1939, Raukas 1963a). Tills of the last glaciation on the Cambrian blue clays, sand- and siltstones in the fore-klint area are bluish-grey, mostly clayey and contain mainly clasts from Finland and the bottom of the Gulf of Finland. On the crystalline basement, the till is brown or reddish-brown. Stony tills on the Ordovician and Silurian bedrock are enriched with the local carbonaceous material (Photo 29). The constituent clasts are mainly angular. Tills on the Devonian sand- and siltstones are reddish-brown. The rather well-rounded local carbonaceous and erratic crystalline material occurs in various ratios in the cobble and pebble fractions and are under the influence of the Devonian bedrock, comparatively rich in sand and silt fractions (Orviku 1958a, Raukas 1978). In the stratotype area - the basin of Lake Võrtsjärv, both grey carbonaceous (Valma) and reddish-brown (Tamme) tills are widespread.
The Upper Järva Late-glacial deposits are divided into Arctic (Oldest Dryas, Bølling, Older Dryas) and Subarctic (Allerød, Younger Dryas) chronozones (Table 14). According to the decision of the INQUA Congress in Paris in 1969, the Holocene/Pleistocene boundary is accepted as 10,000 14C years.
Traditionally, the Late-glacial interval in Estonia starts from the accumulation of Rauna interstadial deposits in central Latvia (Kajak et al. 1976). In the Raunis section, interstadial sands with alternating layers of silt and clay, which contain peat and plant remains, lie between two layers of till to the southeast from the Town of Cēsis, on the right bank of the Raunis River. Organic remains from the Raunis section have been dated in several laboratories (13,390±500: Mo-196; 13,250±160: TA-177; 13,320±250: RI-39) and the results obtained are in good agreement (Punning et al. 1968).
In mainland Estonia and on the islands of the Gulf of Finland, Eemian (Prangli) deposits or pre-Weichselian tills are in some places (e.g. Prangli Island) overlain by four till beds, the exact age of which is uncertain. The upper till beds are regarded as stadial ones of Late Weichselian age representing secondary oscillations of the ice sheet. It is also possible that a two- or three-layered till beds may consist of contemporaneous basal and ablation tills from a single glacial event (Raukas 1963a). Locally, Haanja/Otepää and Pandivere/Palivere tills are separated by terrigenous layers containing subfossil molluscs (Kameri, in Orviku 1939) and pollen assemblages of a cool character. Tills of the Haanja, Otepää, Pandivere and Palivere stadials have specific colour and lithological composition (Raukas 1963a, 1978) and can be regarded as lithostratigraphical units of the lowest taxonomical rank (beds).
In some places intermorainic layers have been dated by the 14C method, but the results are contradictory. In the Kurenurme section, southeastern Estonia, remains of Salix wood were taken from sandy loam overlying Haanja till (Ilves et al. 1974). Quite reliable radiocarbon datings of the wood (12,650±520: TA-57) and organic detritus (12,420±100: Tln-35) indicate that these deposits accumulated at the beginning of the Bølling Interstadial. Unfortunately, the process of the deposition of the organic material is not clear (Karukäpp 1986) and this hampers the usage of the section in the till stratigraphy. In the Kaagvere section southeast of Otepää, the dates obtained on the reddish-brown till (15,150±575: TA-50, >30,000: TA-36) suggest redeposition of older interglacial material. The palynological characteristics of interstadial layers between the stadial till beds are not clear either. Probably, these layers contain a lot of material redeposited from older interglacials (Liivrand 1991).
Palynological studies of pre-Allerød deposits above the till beds in Estonia give evidence of severe climatic conditions throughout the Arctic period. They do not permit the layers related to the Bølling amelioration to be distinguished. Such deposits may be present in southern but hardly in northern Estonia. In the section of Haljala (Männil & Pirrus 1963), a pollen assemblage suggesting a brief interval of warming, possibly Bølling, has been reported from a sandy interlayer at a depth of 10.5-11.2 m. However, as its redeposition in the sandy sediment is not excluded, the kind of pollen data interpretation must be taken with great caution (Pirrus & Raukas 1969), and the more that no corresponding warm interval is known from any other site in Estonia.
Deposits of Older Dryas age occur both in northern and in southern Estonia. The lower boundary of the Older Dryas is undefined in Estonia (Kajak et al. 1976), but probably it is the boundary between the Otepää and Kurenurme beds.
In the light of the pollen evidence, the retreat of the ice margin from the Haanja position started during the transitional from Oldest Dryas to Bølling time and the deglaciation of the Estonian territory was completed during the second half of the Allerød (Pirrus & Raukas 1969).
According to Reet Pirrus (Pirrus & Raukas 1996), some more or less clear trends in the vegetation history could be given (Table 15).
Artemisia - Chenopodiaceae palynozone
The Older Dryas Chronozone is represented by glaciolacustrial varved clays or rhythmically laminated silts and sands and overlain by lacustrine silts and clays. In the southern part of Estonia, minerogenic lacustrine sediments may contain minor amounts of plant remains. The thickness of deposits ranges from 1.3 to 11.3 m. In the Older Dryas about 12,000 years ago the Baltic Ice Lake formed and corresponding deposits started to accumulate.
This zone is characterized by high herb pollen percentages (Artemisia, Chenopodiaceae, Helianthemum, Cyperaceae, Gramineae, and several other species of primary vegetation) along with Betula nana L.
The Allerød Chronozone is represented by lacustrine clays and silts (0.15-1.85 m in thickness) with blackish-grey interlayers and the Baltic Ice Lake sediments. Scattered plant remains, mostly leaves and stalks of Bryales moss are common in lake sediments.
The Allerød Chronozone is subdivided into two pollen zones (Pirrus & Raukas 1996): a) Pinus-Betula Zone (ALa), b) Pinus Zone (ALb).
The lower boundary of AL Chronozone is fixed with a rather distinct increase of AP pollen and decrease of herbs (Artemisia, Chenopodiaceae) and Betula nana L.
Characteristic of AL Chronozone is the prevalence of tree pollen. Betula shows a rapid increase and towards the uppermost part of the zone Pinus increases distinctly and has its Late-glacial culmination. At the same time, herb pollen is at its minimum. Betula nana L. is constantly present in low percentages. The variety of bog and meadow species of terrestrial herbs and water plants has increased. Xerophytes, halophytes, heliophytes and tundra plants are continuously present, but in low values. Fine preservation and abundance of pollen as well as the regularity of pollen curves indicate their bedding in situ.
Artemisia - Betula nana Zone
The younger Dryas Chronozone is represented with the Baltic Ice Lake and Yoldia Sea sediments and by lacustrine silts and clays, often with fine-grained sand interlayers. Bryales remains are scattered or occur as thin layers, occasionally abounding in hydrotroilite. The thickness of lacustrine deposits ranges from 0.2 to 4.0 m.
The zone boundary AL/DR3 is placed at the strong and rapid increase of the content of herb pollen (particularly Artemisia) and Betula nana L. This zone is characterized by remarkably high frequency of herb pollen ranging from 40-60%. Maximum values of Betula nana L. pollen in different profiles range from 20 to 25%. The Late-glacial culmination of Picea is either in the lowermost (Võru, Visusti, Haljala) or uppermost (Remmeski, Loobu) part of the pollen zone.
The boundary DR3/PB is placed at a rapid increase of tree pollen, prevailingly Betula (about 80%, in SE Estonia up to 90%) and Pinus (about 20%).
Holocene deposits and their stratigraphical subdivision
The Holocene continental deposits, occasionally rather thick, occur practically everywhere above the Pleistocene deposits. Unfortunately, the offshore and nearshore marine deposits are characterized by numerous unconformities and rapid facies changes and in many sequences gaps cover longer time spans than the preserved strata. The main stages in the Baltic Sea history are known from the very beginning of the century, but they have never been properly defined as stratigraphical units (Hyvärinen & Raukas 1992). Therefore, the stratigraphical scheme of the Holocene deposits (Table 16) is mainly based on the continental deposits. The existence of the four major phases in the postglacial history of the Baltic – the Preboreal Yoldia “Sea”, the Ancylus Lake, the Litorina Sea and the Limnea Sea – is recognised.
In Estonia, there are 9836 peat bogs and about 1150 lakes greater than 1 ha in area (Mäemets 1976, Orru 1992). The peat is at its thickest (16.8 m) in the small Vällamäe kettle hole. The peat deposits are usually 8-10 m thick. The greatest thickness of organic lacustrine deposits is 18 m (Väimela-Alajärv), lake marl 6-7 m (Tapa, Kulina), travertine 5-6 m (Loosi, Rõuge), alluvial deposits 15 m (Väike-Emajõgi) and aeolian deposits 15-20 m (Sininõmme, Kõpu, Rannametsa). In the first half of the century, the palynological approach (Thomson 1925) was applied to the stratigraphical studies in Estonia, at the end of the fifties, physical dating methods were taken into use (Ilves et al. 1974). P. Thomson investigated lake and mire deposits in about 20 localities and modified his first (Thomson 1925) Estonian Holocene stratigraphical scheme in several high standard publications (Thomson 1926, 1929, 1930, 1933). Some 30 years later K. and L. Orviku published the next Holocene stratigraphical scheme (Orviku 1956, Orviku L. 1960).
The first official stratigraphical chart of the Estonian Holocene deposits was compiled under the leadership of Prof. K. Orviku and accepted in May 1976 (Kajak et al. 1976). The second official stratigraphical chart presented in this book (Table 16), was compiled by R. Pirrus, A. Raukas and S. Veski (Raukas et al. 1995b). The part of the scheme dealing with continental deposits is based on the studies by H. Kessel, R. Pirrus, A. Sarv, L. Saarse, K. Kimmel, T. Koff, L.-K. Königsson, S. Veski and A. Poska. The investigations of J. Lutt, H. Kessel and A. Raukas underlie the subdivision of marine deposits. The scheme was accepted at the session of the Estonian Stratigraphical Commission on May 6, 1993 and a week later it was approved at the Stratigraphical Conference of the Baltic States in Vilnius. The regional chart for the Baltic States was approved at the same time. They both followed the Scandinavian scheme (Mangerud et al. 1974). The charts have parallel subdivisions for the continental and marine deposits.
According to the international rules, stratigraphical charts should be based on the unit and boundary stratotypes. Unfortunately, up to now there are no officially accepted stratotype sections for Holocene deposits in Estonia or in the other Baltic States. This work is in progress.
Each site with its own local pollen assemblage biozones is effectively its own stratotype, but no stratotype can exist for the regional pollen assemblage biozones, which are artificial synthesis (Turner 1989). The same type of artificial synthesis is the proposed local stratigraphical chart (Table 16), based on the multiple sections throughout Estonia, all having their own characteristics. As the pollen zones are time transgressive, the boundaries between palynological chronozones have not been drawn and this makes the chart useful and applicable in Estonia as a whole.
R. Perens & L. Vallner
In terms of groundwater formation and circulation, the groundwater system in Estonia can be divided into three principal hydrostratigraphical units.
1. T h e Q u a t e r n a r y d e p o s i t s. The sandy and clayey Quaternary deposits and peat form porous aquifers with mainly unconfined groundwater which are directly influenced by meteorological conditions. The whole infiltration water percolates into the Quaternary cover and the greater part of groundwater discharge flows through it. The upper portion of the Quaternary cover or sporadically all Quaternary deposits belong to the aeration zone where a lot of water circulates by the agency of capillary force or evaporates, in addition to the filtration flow.
2. T h e b e d r o c k. The terrigeneous and carbonate Palaeozoic and Proterozoic rocks form porous, fissured and karstified, mostly confined aquifers, which are isolated from each other with aquitards of different isolation capacity. In the karst cavities near the ground, shallow groundwater flows very fast and its chemical composition is close to that of the surface water. However, the deeper strata contain pre-Quaternary groundwater, which is high in total dissolved solids (TDS) and moves very slowly under natural conditions.
3. T h e c r y s t a l l i n e b a s e m e n t. Predominantly pre-Quaternary groundwater in the fissures of igneous and metamorphic rocks contains a high rate of TDS and under natural conditions it is sporadically almost stagnant. The lower portion of the crystalline basement serves as an aqifuge for the whole overlying water-bearing formation in Estonia (to the exclusion of the hypothesis about water originating and arising from the depths of the Earth’s crust).
Aquifer and aquitard are the units of detailed hydrogeological stratification of the water-bearing formation. Aquifer is a relatively homogeneous water-bearing layer or rock with similar water conductivity and storage capacity yielding water in a useable quantity to a well.
According to the hydraulic conductivity value K*, the degree of the permeability of water-bearing strata is the following:
102 ≤ K < 1 very low
1 ≤ K < 3 low
3 ≤ K < 10 medium
10 ≤ K < 30 high
30 ≤ K < 70 very high
K > 70 extremely high
Permeability in a lateral direction can be up to 100 times higher than in a transversal direction.
Aquitards are the strata, the transversal conductivity Kt of which is generally less than 10-2 m/d. The following degrees of impermeability can be distinguished:
10-2 > K > 10-4 weak
10-4 ≥ Kt > 10-6 medium
10-6 ≥ Kt > 10-8 strong
Kt < 10-8 very strong
Not a single aquitard with the above-mentioned filtration characteristics has an absolute isolation ability. According to this gradation, even strong aquitards are permeable to up- or downward groundwater flows, the total amount of which in large areas can extend up to 104 m3/d.
Aquifers which lie one over another are not necessarily isolated with aquitards. The rocky complex consisting of aquifers and aquitards with similar hydraulic characteristics but with different rock types is termed aquifer system.
In terms of the real water supply, the aquifers and aquifer systems can be subdivided into sufficiently water yielding aquifers and aquifer systems (with specific capacity of wells correspondingly q > 0.1 l/(s×m) ≈ 10 m3/(d×m), K > 1 m/d) and weakly water yielding aquifers and aquifer systems (q < 0.1 l/(s×m), K < 1 m/d). Aquifers and aquifer systems can be sufficiently or weakly water yielding either sporadically (locally) or in the whole distribution area. According to the aforenamed criteria, at least 23 aquifers and 4 regional aquitards can be distinguished in the water-bearing formation of Estonia (Table 17, Figs. 94, 95).
The technogeneous deposits (tQIV) in settlements mostly consist of stuff and building waste. Besides, there are 495 dumps of different size in Estonia. In the mining region extensive areas are under spoil heaps and oil shale plateaus. The water in technogeneous deposits is usually highly polluted. The water leaking through ash hills and dumps is dangerous to the environment.
The boggy deposits (bQIV) are mostly represented by peat. Under natural conditions, the water generally lies at a depth of 0.1...0.5 m, being deeper only in the fields of milled peat. The conductivity of peat is 0.3...1 m/d. The inflow of water into experimental pits amounts to 1...10 m3/d per 0.5...1 m of drawdown. Bogs recharge from precipitation, while the replenishment to fens is also from lateral flows of unconfined groundwater and vertical flows of confined groundwater from deeper strata. The water of boggy deposits has a nasty taste and smell and is practically not used for the water supply.
The aeolian deposits (vQIV) are mainly represented by fine-grained and well- sorted sands of dunes on the northern coast of Lake Peipsi and on ancient and present beaches of the Baltic Sea. Due to its chain-like morphology, the upper and greater portion of a dune is usually dry; water occurs in the lower part at a depth of 10...15 m from the surface. The yield of wells does not usually exceed 1 to 5 m3/d. At the sea coast the occasional intrusion of brackish sea water can take place.
The lacustrine deposits (lQIV) occur in limited areas in intermittent strata of loamy sand, loam and sapropel mostly in the Alutaguse and Võrtsjärve lowlands. The strata are poorly permeable and not suitable for water supply.
The alluvial deposits (aQIV) are represented by gravel, sand, loamy sand and loam of river valleys with a total thickness of up to 15 m. Due to the limited distribution, they do not have any practical importance.
The marine deposits (mQIV) are up to six, occasionally even more metres thick and consist mostly of sand and coastal gravel which are found in Lower Estonia. In places, the water level can lie close to the ground. Water can sometimes be weakly confined due to the occurrence of clayey interlayers in sands. The discharge of 2...4-m-deep wells ranges from 10 to 60 m3/d and this water is used in many households.
The glaciolacustrine deposits (lgQIII), with a total thickness of 5...10 m, cover a large area and are represented by fine-grained sand, loamy sand and varved clay. Sands and light loamy sands are weakly or sufficiently water yielding with their conductivity varying from 0.1 to 5 m/d. Many wells with the discharge ranging from 0.5 to 20 m3/d have been sunk into these deposits.
Varved clay (lgQIII) with a thickness of up to 22 m and transversal conductivity less than 10-4 m/d, forms medium and strong local aquitards all over Estonia. The largest aquitard (30 km2) occurs in the catchment of the Kasari River in western Estonia. Varved clay effectively protects deeper aquifers from pollution.
The glaciofluvial deposits (fQIII-II) form frontal aprons, eskers and deltas and occur in some buried valleys. They consist mostly of sand and gravel, the conductivity of which is generally 5...10 m/d, in some places even up to 100 m/d. Due to this, the wells tapping the glaciofluvial deposits are generally high yielding. Glaciofluvial sediments in buried valleys are usually confined by clayey glaciolacustrine deposits and till. Public water intakes with a pumping rate of up to 10,000 m3/d tap glaciofluvial aquifers in the buried valleys at Vasavere near Jõhvi, at Raadi-Maarjamõisa in Tartu, and at Männiku-Pelguranna in Tallinn.
The glaciogeneous sediments (gQIII) cover almost 2/3 of Estonian territory. Weakly or sufficiently water yielding are the loamy-sandy varieties of till and sporadically spread lenses of sand and gravel in till with a thickness of a couple of metres. The conductivity of loamy-sandy till ranges from 0.01 to 1.0 m/d. The majority of up-to-10-m-deep dug wells all over Estonia get water from till. The discharge of these wells is predominantly 0.2...2 m3/d. Usually the water table is at a depth of 1.5...3 m from the surface, quite often it is at a depth of 8...12 m, in the Otepää and Haanja heights occasionally even 20 m below the ground. In late summer, shallow (2 - 5 m) wells in loamy-sandy till often run dry. The loamy-sandy till with the conductivity of 10‑3...10-4 m/d is considered a weak or medium aquitard.
The Upper Devonian aquifer system (D3) consists of karstified and fissured dolomites and dolomitized limestones of the Dubniki and Pļaviņas stages. The total thickness of this aquifer system is 17...25 m and it covers some 500 km2 in southeastern Estonia (Fig. 94). The siltstone of the Snetnaja Gora Member with interlayers of clay in the lower portion of the Pļaviņas Stage forms an aquitard with medium isolation ability. The aquifer system is overlain by the Quaternary cover with a thickness of 40...80 m. Groundwater is mostly confined and its potentiometric surface lies at a depth of 3...8 m from the ground. Big sink-holes through which melt- and rain-water percolates fast into the karstified bedrock strata occur at Rõuge, Meremäe, Meeksi and some other places. Conductivity of karstified carbonate rocks varies between 1...50 m/d. According to this, the specific capacity of wells ranges from 0.1 to 6.0 l/(s×m), predominantly it is 1 l/(s×m). Due to its limited occurrence, the Upper Devonian aquifer system is used for the public water supply in a few places only.
The Middle Devonian aquifer system (D2) is extending in southern Estonia (Fig. 94) between the Gulf of Liivi (Riga) and Lake Peipsi. It consists of sand- and siltstones with interlayers and lenses of clay of the Amata, Gauja, Burtnieki and Aruküla stages. Clayey material prevails in the Amata Stage, forming with the Snetnaja Gora Member a weak or medium aquitard between the Upper and Middle Devonian. One third of the volume of this aquifer system includes clayey rocks which serve as weak or medium aquitards and, for this reason, probably form several confined aquifers of local distribution (Verte 1965). The occurrence of the latter has not been sufficiently proved yet.
The northern boundary of the distribution of the Middle Devonian aquifer system lies approximately on the Häädemeeste - Mustvee line. To the south from this line, the thickness of the aquifer system increases up to 250 m in the Haanja Heights. The aquifer system outcrops only occasionally in deeper river valleys, elsewhere it is covered with Quaternary deposits, ranging 5...80 m in thickness. In uplands the potentiometric surface lies at a depth of 10...15 m from the surface, while in lower areas a lot of flowing wells are encountered (Tõrva, Valga, Antsla, Võru, etc.). The absolute height of the potentiometric surface ranges from 80 to 130 m in the Otepää and Haanja heights, in the Sakala Upland it is between 50...80 m.
The lateral conductivity of aquifer system is rather equable: predominantly 1...3 m/d. Transmissivity reaches 200...500 m2/d in the Sakala Upland, Otepää and Haanja heights, elsewhere it is usually less than 100 m2/d. The storage coefficient amounts to 5×10-5...10-3. The discharge of wells changes between 200...700 m3/d per 3...7 m of drawdown. The specific capacity of wells is predominantly 0.4...1 l/(s×m). The Middle Devonian aquifer system is used for the public water supply mainly in the areas south of the Häädemeeste - Põlva line, but also in the towns of Tartu, Viljandi, Elva and Kallaste.
The Narva regional aquitard (D2nr) consists of layers of siltstone, dolomite, marl and clay of the Narva Stage with a total thickness of up to 90 m. In southern Estonia, these layers form the uppermost effective bedrock aquitard, the transversal conductivity of which is 10-5...10-4 m/d, in places 10-6 m/d or even less. The clayey layers of the Narva Stage serve as a regional aquitard for the whole Baltic Artesian Basin (Juodkazis 1989). The rocks in the upper portion of the stage supply water for the area between Tartu and Mustvee and for Ruhnu Island. The specific capacity of wells is 0.06...0.2 l/(s×m). The Narva aquitard separates the Middle Devonian aquifer system from the underlying water-bearing strata.
The Middle-Lower-Devonian aquifer system (D2-1). The Narva aquitard is underlain by the water-bearing layers of the Pärnu Stage (Middle Devonian) and Rēzekne and Tilžė stages (Lower Devonian) which consist of fine-grained weakly cemented sand- and siltstones with interlayers of clayey and dolomitized sandstone. Together with the underlying Silurian strata the layers are used for the public water supply in Pärnu, Viljandi and Tartu. The association of water-bearing strata has been named the Middle-Devonian-Silurian aquifer system and the united account has been kept of its water extraction and water resources (Savitski et al. 1996). However, in view of the collector characteristics of the rocks, it would be more correct to treat the complexes of terrigeneous and carbonate rocks separately.
In southern Estonia, the Middle and Lower Devonian aquifer system with a thickness of up to 100 m lies at a depth of more than 200 m below sea level (Fig. 95). The water is predominantly confined. In lowlands, where the potentiometric surface extends above the ground, flowing wells occur. In the uplands, the potentiometric surface is at a depth of 10...20 m below the ground.
Due to its good collector characteristics, the water yielding capacity of sandstone is relatively high. The discharge of wells is predominantly between 260...700 m3/d by drawdown of 4...10 m. The specific capacity of wells ranges from 0.25 to 1.0 l/(s×m). Conductivity of sandstones is mostly 2...6 m/d, rarely 8...10 m/d. Transmissivity of the aquifer system is 50...500 m2/d, the storage coefficient ranges from 0.001 to 0.15.
The Silurian-Ordovician aquifer system (S-O) is an important source of water supply in the regions north of the Pärnu - Põlva line and on the islands of the West-Estonian Archipelago (Fig. 96). It consists of diverse limestones and dolomites with clayey interlayers. The upper portion of the rocky complex with a thickness of 30 m is extremely cavernous, with numerous cracks and fissures (Heinsalu 1977). Karst cavities form some half-a-metre-high canals trending in the direction of bedrock fissures. Caverns are especially abundant in dolomites and dolomitized limestones. Close to the ground, bigger karst cavities, a couple of metres high and some twenty or thirty metres long, occur in some places. Karst phenomena and fissures are most abundant in the carbonate rocks forming the upper part of the bedrock (Photo 32) - the weathering zone, usually 1...3 m, rarely 5...10 m in thickness. The deeper the lying depth of carbonate rock, the less fissures and cavities; such rocks generally turn into an aquitard. Southward from the Ikla - Elva - Mehikoorma line the Silurian-Ordovician rocks practically yield no water.
Besides traditional aquifer pumping tests, the impeller method (flowmeter logging) has been widely used in studying the filtration characteristics of Estonia’s bedrock (Perens 1984, Perens & Paltanavičius 1989, Perens et al. 1994). The results indicate (Fig. 97) that the Silurian-Ordovician carbonate rocks have fragmentary water conducting zones with parallel lamination and an abundance of fissures. In these 1...2-m-thick zones groundwater flows in a lateral direction (wells included). Water conductivity zones are separated from each other by 5...10-m-thick layers in which groundwater flows predominantly in vertical fissures. Only about 13% of the whole length of the rock complex is covered by lateral water zones (Perens 1984). As an average, there are about 5 water conducting zones per 100 m of vertical cross-section. In a lateral direction the water conducting zones are fragmentary and their stratigraphical level may be more or less the same only within a couple of kilometres. The water zones of quite different stratigraphical levels can be found in wells lying only a few hundred metres from each other.
According to the logging data gathered in more than 300 wells, about half of water in these wells is provided by the upper portion of cross-section with a thickness of 15 m and average transmissivity of 400 m2/d. Downwards the total transmissivity of carbonate rocks is evenly increasing and at depths of 50 and 75 m it is 630 and 700 m2/d, respectively. As new water conducting zones are very rare in deeper layers, the depth of 75 m can be considered the lowest border of sufficiently water yielding layers of the whole Silurian-Ordovician aquifer system. In western Estonia and on islands, the thickness of the main water yielding portion of the Silurian-Ordovician aquifer system is only 30...40 m. In most of Estonia, the total transmissivity of the Silurian-Ordovician aquifer system usually varies between 100...500 m2/d, depending essentially on the distribution of zones of tectonic disturbances. In those areas transmissivity increases and often exceeds 1000 m2/d.
The lateral conductivity in the carbonate bedrock is variable: 10...50m/d in the topmost 20 m, 5...8 m/d at a depth of 20...50 m, and only 1...2 m/d at a depth of 50...100 m. The lateral conductivity of deeper strata does not usually exceed 1 m/d, although occasionally strata with considerably higher conductivity can be found even as deep as 200 m from the surface (Fig. 97). The data of 235 pumping tests in the oil shale region of northeastern Estonia have shown that the lateral conductivity of Ordovician carbonate rocks near the ground is predominantly between 5...300 m/d, while at a depth of 80...100 m it is only 0.1 m/d (Riet 1974, 1976).
According to the water budget calculations, the transversal conductivity of the layers between the lateral water conducting zones is 10-5...10-2 m/d (Vallner 1980, Jõgar 1983). These interlayers serve as weak or medium aquitards confining the local aquifers of different range, the distribution of which is not yet completely clear. In northeastern Estonia (Verte 1965, Gazizov 1971, Jõgar 1977, Savitski et al. 1996, a.o.), including the region of oil-shale mines with plentiful experimental data, the following aquifers and aquitards can be distinguished (from top downwards): Nabala-Rakvere aquifer (O3-2nb-O2rk), Oandu aquitard (O2on), Keila-Jõhvi aquifer (O2kl-O2jh), Jõhvi-Idavere aquitard (O2jh-O2id), Idavere-Kukruse aquifer (O2id-O2kk), Uhaku aquitard (O2uh) and Lasnamäe-Kunda aquifer (O2ls-O1kn). The occurrence of aquitards and aquifers in the Silurian-Ordovician aquifer system has also been proved by hydrogeological modelling. The satisfactory calibration of the more extensive filtration models is impossible without this distinction (Vallner 1996b).
The average lateral conductivity of the Silurian-Ordovician aquifer system is 8.1 m/d. Higher values of conductivity have been recorded in the limestones of the Raikküla Stage (24.2 m/d) and in the Adila Formation (17.0 m/d). The average conductivity of the Idavere and Uhaku stages is only 0.3 and 0.7 m/d, respectively (Perens 1989).
Water in the fissure systems and karst cavities of the carbonate bedrock flows relatively fast. In the outcrop area of aquifer system it recharges from Quaternary deposits and, for this reason, can easily become polluted in the areas with a thin Quaternary cover (Photo 26). The upper portion of aquifer system enfolds an aeration zone which in North-Estonian uplands is up to 20 m and in alvars only a couple of metres thick. Elsewhere, the Silurian-Ordovician aquifer system is more or less confined by the Quaternary cover and the uppermost aquitards of the carbonate bedrock. On the foot of uplands and in river valleys the potentiometric surface can often be 0.5...2 m above the ground which is the reason for many springs and flowing wells (Heinsalu & Vallner 1995). The North-Estonian watershed serves as the main area of head generating. In the Pandivere Upland, the groundwater level is at a height of up to 110...120 m above sea level.
The specific capacity of wells tapping the upper portion of the aquifer system ranges from 1 to 3 l/(s×m), being 3...5 l/(s×m) on average. The specific capacity of wells deriving water from deeper strata does not usually exceed 1 l/(s×m). The average yield of wells by drawdown of 5...10 m is usually 400...900 m3/d. The average specific yield of unconfined aquifer system is 0.02...0.06, depending on the degree of fissuration and karstification. The storage coefficient changes between 10-6...10-3.
The Silurian-Ordovician regional aquitard (S-O) consists of limestones, marls, siltstones, clays and argillites of the Toila, Leetse, Varangu and Türisalu formations (Lower Ordovician), extending at a length of 30 km southward from the North-Estonian Klint. Farther in the south, the aquitard includes all Silurian and Ordovician rocks. Its thickness increases from a couple of metres in the vicinity of the klint up to 200...350 m near the southern border of Estonia. Conductivity is very variable: the lateral conductivity changes between 0.001...1 m/d, extending sometimes up to 2...5 m/d (Fig. 97), the transversal conductivity ranges from 10-7 to 10-5 m/d (Vallner 1980, 1996a).
The Ordovician-Cambrian aquifer system (O-C) underlies the Silurian-Ordovician regional aquitard extending under most of Estonia, except the North-Estonian coastal region, Mõniste-Lokno uplift area and the islands of the West-Estonian Archipelago (Fig. 98). The aquifer system includes the Kallavere-Tiskre aquifer (O1kl- C1ts) in mainland Estonia which consists of fine-grained sandstone and siltstone of the Lower Ordovician Kallavere Formation and the Lower Cambrian Tiskre Formation. The thickness of the aquifer system is 20...60 m, it increases from north to south. The depth from the ground increases from 10...20 m at the North-Estonian Klint up to 500 m on Estonia’s southern border. The main recharge area is the Pandivere Upland where water from Ordovician strata leaks downward through the Silurian-Ordovician regional aquitard and disperses in radial directions as confined filtration flows. There the absolute altitude of potentiometric surface is up to 70 m under natural conditions. In Lower Estonia, Võrtsjärve, Alutaguse, Valga and Varnja-Värska lowlands and in the Väike-Emajõgi Valley, the potentiometric surface of the Ordovician-Cambrian aquifer system extends above the ground. The aquifer system is an important source of public water supply in northern Estonia and it is also intensively used in the towns of Pärnu, Viljandi and Tartu. This has caused several depressions of potentiometric surface (Fig. 98).
The lateral conductivity of the Ordovician-Cambrian aquifer system mostly ranges from 1 to 3 m/d. Transmissivity tends to decrease southwards. However, due to the thickness of water-bearing strata it is 80...130 m2/d in central and southeastern and only 25...50 m2/d in northern Estonia. The yields of wells are predominantly 430...600 m3/d per 10...15 m of drawdown. The specific capacity of wells changes between 0.2...0.4 l/(s×m). The storage coefficient is 2.5×10-5...6×10-3; the specific yield of the aquifer drained is 0.12...0.14.
On the islands of the West-Estonian Archipelago, the Ruhnu (C1rh) and Soela-Tiskre (C1sl-ts) aquifers, isolated by the Irben (C1ir) aquitard, belong to the Ordovician-Cambrian aquifer system. They consist of the Lower Cambrian sand- and siltstone (Ruhnu and Soela formations).
The Lükati-Lontova regional aquitard (C1lk-C1ln), spread in most of mainland Estonia, is represented by siltstones and clays of the Lower Cambrian Lükati and Lontova formations. The total thickness of layers decrease from 90...100 m in northern Estonia to pinching out on the Mõisaküla - Vastseliina line in southern Estonia and in western Estonia (Fig. 99). This aquitard has a strong isolation capacity; the transversal conductivity is predominantly between 10‑7...10‑5 m/d (Vallner 1980, 1996a, unpublished report). In the West-Estonian Archipelago, the Ordovician-Cambrian aquifer system is isolated from the underlying Cambrian-Vendian aquifer system with clays and clayey siltstones of the Lower Cambrian Lükati and Sõru formations.
The Cambrian-Vendian aquifer system (C-V). Cambrian-Vendian terrigeneous rocks occur all over Estonia, except the Mõniste-Lokno uplift area. The water yielding portion consists of sand- and siltstones with interlayers of clay. The difference between the cross-sections of western and eastern Estonia is obvious (Figs. 94, 95, 99). East of the Rakvere - Põltsamaa - Otepää line, the up-to-53m-thick clays of Kotlin Formation (V2kt) divide the aquifer system into two aquifers. The upper, Voronka aquifer, consists of quartzose sand- and siltstone with a thickness of up to 45 m in northeastern Estonia. The lower, Gdov aquifer is formed of up-to-68m-thick complex of mixed-grained sand- and siltstone.
In northern Estonia, the aquifer system is confined by 60...90m-thick clays of the Lontova Formation. Westwards from the Tallinn - Pärnu-Jaagupi line, the Lontova Formation is gradually replaced by interbedding clay and sandstone of the Voosi Formation, which attain a thickness of 90 m in southwestern Estonia. On the West-Estonian islands, the Vendian deposits have also been pinched out and the water-bearing terrigeneous rocks consist only of Cambrian sand- and siltstones with interlayers of clay.
The Cambrian-Vendian aquifer system is the most important source of public water supply in northern Estonia. Intensive water extraction has led to the formation of two extensive depressions of potentiometric level (Fig. 99).
The Voronka aquifer (V2vr) in eastern Estonia consists mainly of quartzose sand- and siltstones of the Voronka Formation, up to 40 m in thickness.The conductivity of rocks ranges from 0.6 to 12.5 m/d, being 2...6 m/d on average. Transmissivity decreases from 100...150 m2/d in northern Estonia to 50 m2/d and even less in the southern direction. In northern Estonia, the specific capacity of wells ranges from 0.2 to 5.8 (average 2) l/(s×m). In central and southern Estonia, it is 0.1...0.3 l/(s×m). Under natural conditions, the potentiometric level in the coast of the Gulf of Finland was 1.5...5.5 m above sea level.
The Gdov aquifer (V2gd) consisting of mixed-grained sand- and siltstone with the thickness predominantly between 40...65 m lies directly on the Precambrian basement. The clay of the Kotlin Formation serves as an upper confining unit. In northern Estonia, the conductivity of water-bearing rocks is 0.5...9.2, average 5...6 m/d. Transmissivity in northeastern Estonia is 300...350 m2/d; it decreases in a southerly and westerly direction to 100 m2/d or less. Specific capacities of wells differ, the average value being 1.5...2.5 l/(s×m). Since most of wells tap both the Gdov and Voronka aquifers, their specific capacity is considerably higher. The potentiometric surface was 3...5 m above sea level under natural conditions in the coastal area of northern Estonia .
Westward from the line where the Kotlin clays are pinching out (Fig. 99), the Cambrian and Vendian water-bearing rocks form the steady Lontova-Gdov aquifer. On the West-Estonian islands, the Vendian sediments are absent and all the rocks deeper than the Lükati-Sõru aquitard (C1lk-sr) form the Voosi aquifer (C1vs). The productivity of wells on the islands of the West-Estonian Archipelago is 3...4 times lower than in northern Estonia.
The crystalline basement (PR1) comprises groundwater in its upper weathered and fissured portion only. The specific capacity of wells does not exceed 0.1...0.2 l/(s×m). Natural potentiometric surface refracts the heads of the overlying Cambrian-Vendian aquifer system. Flowing wells occur in some places. Currently, the water stored in the crystalline basement is not used for water supply.
In extensive regional hydrogeological reviews the Estonian water-bearing formation has been treated as the northern slope of the Baltic Artesian Basin extending from the Gulf of Finland up to Minsk and Warsaw (Juodkazis 1989). This viewpoint, although based on structural and geological aspects and a certain unity of formation of palaeohydrogeological conditions of deep groundwater, is still a theoretical construction. From the applicational point of view, the Estonian water-bearing formation should be considered as an independent artesian basin because the exchange of underground water with the neighbouring areas is less than 0.1% of the total annual groundwater recharge (Vallner 1980).
L. Vallner
In Estonia, groundwater is recharged from rain- and melt-water percolating through unsaturated soils. Mean annual precipitation is in the range of 500...750 mm or 1370...2060 m3/(d×km2), and is lower on the coast and some 10...20% higher on uplands.
The net infiltration (I) for Estonia (total groundwater recharge minus evaporation from the zone of saturation or ca-pillary fringe) has been calculated preliminary from the budget equation comprising the main components of groundwater flow (Vallner 1976, 1980):
I = R + Q + M - W ± V ± S
where R is the groundwater discharge (base flow) m2/d to streams; Q is the pumpage from layers; M is the direct seepage of groundwater to the sea; W is the flux from streams into aquifers (induced recharge); V is the subsurface exchange of groundwater between Estonia and surrounding areas and S is the storage change.
The long-term groundwater discharge to streams (R) has been estimated on the basis of observations carried out during several decades at more than 100 hydrological gauging stations all over Estonia. Apart from the gauging stations, many irregular measurements of the low flow have been made approximately in 1000 stream cross-sections. The gained sporadic low flow data were modified to average base flow value by statistical methods using regular observations of gauging stations (Vallner 1980). The pumpage data (Q) were obtained from state institutions checking the groundwater use. The subsurface fluxes to the sea (M) and groundwater exchange with adjacent areas (V) were calculated by Darcy’s formula.
The total groundwater discharge to the channel network of Estonia (R) is approximately 7,700,000 m3/d, but its intensity varies with regions (Fig.100). Average pumpage from wells and mines (Q) reaches 1,000,000 m3/d causing the inverse fluxes (W) in an amount of ca. 500,000 m3/d from surface bodies of water into aquifers. Direct groundwater seepage to the sea (M) averages 800,000 m3/d. In the west and north, Estonia is bounded by the sea. In the east, its border runs along the Narva River and central portion of Lake Peipsi, in the south the border generally coincides with the divides of major streams. As a result, the exchange of groundwater with adjacent territories (V) is quite insignificant - not more than 10,000 m3/d. For a long-term period the storage change S = 0. The total net infiltration (I) calculated by the budget equation averages 9,000,000 m3/d.
Groundwater discharge to the channel network (the base flow) and the instrumentally checked pumpage make up about 90% of the sum of the right-side members in the above water budget equation. Therefore, the value of net infiltration estimated by budget equation is probably more authentic than that based on indirect data, such as the air temperature and atmospheric humidity, evapotranspiration, etc. (Lerner et al. 1990). After completing the general budget of groundwater flow main components, the distribution of heads was simulated for the whole Estonian water-bearing formation (Vallner & Tobias 1984). Thereby the net infiltration was estimated by the trial and error method pursuing an optimum coincidence between the modelled and measured data.
The total amount of groundwater in the cracks and pores of Estonia’s water-bearing strata is nearly 2000 km3. Both the distribution of the groundwater head and the direction of subsurface flows depend on spacial relations of the topography, surface bodies of water, and impermeable crystalline bedrock. Besides, human impact on the groundwater flow has been continuously increasing during the last five decades. Based on the total effect of the topography and geological structures, three main groundwater flow systems can be recognized in Estonia (Vallner 1980, Tóth 1995)
The local flow system enfolds chiefly the unconfined or locally confined shallow groundwater moving from its recharge area toward the nearest ditches, creeks, rivers, and discharging directly to lakes or to the sea in coastal areas (Fig. 101). The length of the upper branches of the local flow system usually does not exceed a few kilometres, but lower branches which are not drained by small surface waterbodies and discharge to the middle courses of rivers can reach 10...15 km in length. The vertical thickness of the local flow system mostly ranges from 10 to 30 m.
The intermediate flow system takes its rise from Upper Estonia (Fig. 2) where the height of the terrain above sea level is more than 40 m. In the Harju Plateau, Otepää and Haanja heights, Pandivere, Jõhvi, Sakala and Karula uplands and in the Vooremaa area, the local maxima of the groundwater potential energy occur. In the Haanja and Otepää heights the groundwater table is 120...280 m above sea level, while in uplands and plateaus its height mostly ranges from 50 to 100 m. The groundwater head declines down from above in uplands; the difference of heads in some aquitard can be up to 50...60 m, and the corresponding head gradient up to 2.5...3. Consequently, a downward groundwater flow recharges the underlying aquifers from the overlying ones in highlands.
Groundwater moves from Upper Estonia towards Lower Estonia where the surface altitudes are 0...40 m above sea level. Branches of the intermediate flow system discharge to surface waterbodies between absolute heights of 30...40 m in the Võrtsjärv, Alutaguse, Valga and Varnja-Värska lowlands, and in the Väike Emajõgi Valley. The portions of the intermediate flow system formed in the Sakala Upland and in the western part of the Harju Plateau are drained by the lower course of the Kasari River between altitudes 0...30 m, while the lower branches of this flow system discharge to the offshore Baltic Sea. The intermediate flow system is completely confined. Its branches, approaching the discharge areas, bend up and recharge the local flow system from the underside. It is proved by decreasing of the groundwater head from below upward observed in lake depressions and lower course valleys of most significant rivers where a lot of deep flowing wells occur. A portion of the intermediate flow system formed in the Harju and Viru plateaus, Pandivere and Jõhvi uplands discharges through springs or seeps out from layers on the North-Estonian Klint. The lateral length of the intermediate flow system is up to 100 km.
The regional flow system takes its rise in those parts of the Haanja and Otepää heights where the groundwater table is 180 to 280 m above sea level. There the head declines down from above attesting to the existence of a downward groundwater flow. This flow reaching the absolutely waterproof portion of the crystalline basement changes its direction and bends towards the discharge areas which are situated in the depressions of the Baltic Sea and the Gulf of Finland. The regional flow system underlies the intermediate flow system enfolding the portion of the lithosphere where the confined groundwater moves directly into the sea under natural (predevelopment) conditions. In western Estonia, the regional flow system includes the strata of the Ordovician-Cambrian and Cambrian-Vendian aquifer systems lying higher than 250 m below sea level. The Cambrian-Vendian aquifer system belongs completely to the regional flow system west- and northward from lakes Võrtsjärv and Peipsi. The length of lower branches of the regional flow system can reach 300 km between the Haanja Heights and the central part of the Baltic Sea.
The diversity of rock permeability complicates and modifies the head distribution in the subsurface hydrosphere. Therefore, the groundwater movement is mainly parallel to bedding in aquifers, but transversal in aquitards. The latter are often disconnected at tectonic disturbances where relatively intense vertical groundwater flow may occur. The configuration of the potentiometric surface and pumpage data show that all aquitards are more or less permeable in Estonia. Evidence is derived from declining of the head of the Ordovician-Cambrian aquifer system in the opposite directions - from the Otepää Heights and the Pandivere Upland towards the Emajõgi Valley. There the groundwater flows rise up and discharge to the river regardless of the Narva and Silurian-Ordovician confining aquitards with a total thickness of up to 300 m. The transversal permeability of regional aquitards is an eminent, but not completely clarified evidence until now (Tissot & Welte 1978, Brace 1980, Tóth 1995). Apart from porosity, it may well be caused by microfissurization of rocks.
While the potential direction of the groundwater movement is marked by the flow systems described above, the actual quantity and velocity of subsurface flows depend on permeability of layers. Various co-influences of head distribution and permeability are expressed by the vertical zoning of the groundwater flow. In general, the velocity of the groundwater movement decreases with the increase of the flow depth.
The z o n e o f a c t i v e w a t e r e x c h a n g e enfolds the upper portion of the water-bearing formation (Table 18, Fig. 101). All aquifer systems overlying the Lükati-Lontova aquitard north of the Pärnu - Tartu line belong to this zone. Southward the zone is limited from underneath by the Ordovician aquitards or the altitude of 250 m below sea level. In the zone of active water exchange, an upper subzone of fast groundwater flow may be distinguished which includes the Quaternary cover and Devonian aquifers less than 50 m below sea level and the Silurian-Ordovician aquifer system, as a whole. Farther down, a subzone of moderate groundwater flow occurs comprising the portion of the Middle Devonian aquifer at a depth of 50-250 m below sea level and the Ordovician-Cambrian aquifer system down to the lower boundary of the zone of active water exchange.
The s u b z o n e o f f a s t f l o w is immediately influenced by climate. In this subzone the infiltration is accumulating and evaporation from the aeration zone takes place. From the subzone of fast flow all underlying strata are recharged and also the whole groundwater flow discharging to stream channels permeates through this subzone. As mentioned above, in Estonia the total net infiltration reaches 9,000,000 m3/d, which makes an average of 200 m3/(d×km2) or 73 mm per year, but its actual value significantly varies with areas (Fig. 100). Infiltration is most intensive (350 to 770 m3/(d×km2) in the uplands of northern Estonia, where the carbonate bedrock abounds in karst phenomena. In an area of about 1000 km2 in the central part of the Pandivere Upland, where the channel network is entirely lacking, most of rain- and meltwater, which has not been removed by evapotranspiration, percolates into karst interstices. Owing to the high permeability of the karsted carbonate bedrock, a significant portion of this newly formed groundwater quickly discharges to adjacent streams. That’s why the groundwater table is relatively deep in the upland, being often 7...20 m below the ground surface. With such a depth, the discharge by evaporation from the aeration zone is negligible and the net infiltration reaches its maximum in Estonia - up to 900 m3/(d×km2).
In the uplands of southern Estonia, the net infiltration is twice as high as its average value. Highlands form only 16% of Estonia’s total area but, nevertheless, about 40% of the total groundwater recharge takes place just on the uplands and their slopes. On the plateaus of northern and southern Estonia, the net infiltration is 170...260 m3/(d×km2) but in lowlands it is less than 170 m3/(d×km2).
Mires impede the recharge of deeper groundwater by infiltration. Fens are recharged from rising groundwater flows and bogs are usually located above an effective local aquitard which restricts the downward filtration. Similar conditions occur in the domains of glaciolacustrine varved clays covering an area of some thousand square kilometres.
The intensity of groundwater discharge to the channel network varies in a wide range, but its mean value is 170 m3/(d×km2). On the Harju and Viru plateaus, the groundwater discharge into streams averages 130 m3/(d×km2). The groundwater discharge to the channel network is at its highest, reaching 400...900 m3/(d×km2) in some topographic catchments around the Pandivere Upland where a lot of springs occur. In this area, the groundwater discharge averages 300 m3/(d×km2).
S p r i n g s are essential sources of recharge for many streams in northern Estonia. Commonly, they are associated with tectonic faults in the carbonate bedrock and occur in groups. The central part of the Pandivere Upland where the permanent channel network is lacking, is surrounded by a belt of gravity springs at an height of 80 to 100 m above sea level. Unconfined groundwater discharges through these, often intermittent springs (Heinsalu & Vallner 1995). A significant portion of confined groundwater formed in divides is discharged through ascension springs which are situated farther away from the upland in places where tectonic faults cross the river valleys. There groundwater moves upward along tectonic disjunctions of aquitards. About half of rising groundwater discharges directly to the streams dispersively or through subaquatic springs.
The total discharge of some spring groups draining the karsted bedrock is up to 50,000 m3/d during wet periods, but mostly it ranges from 500 to 10,000 m3/d. Besides the Pandivere Upland, there are a lot of springs in the upper course of rivers flowing out from the Harju Plateau and also in the upper course of the Navesti River. All in all, some 1500 springs occur in North-Estonian uplands and plateaus, but predominantly their discharge is less than 50 m3/d. The abundance of springs points to both the high permeability and vulnerability of the carbonate bedrock.
In the uplands and plateaus of southern Estonia, the subzone of fast flow includes mostly till, but also Devonian sand- and siltstones. The intensity of the groundwater discharge into the channel network ranges from 130 to 350 m3/(d×km2). Of an abundance of springs occurring in this area, the vast majority have a discharge of up to 10 m3/d, and only a few springs 1000 to 1500 m3/d. In Lower Estonia (Fig. 2), the intensity of groundwater runoff usually ranges from 50 to 100 m3/(d×km2), i.e. less than the average. This is due to the low and weakly dissected topography and low permeability of upper aquifers. In the Fore-Klint Lowland, the groundwater runoff reaches 200 m3/(d×km2) which is higher than the average value. The klint abounds in springs, but their discharge usually does not outnumber 10...20 m3/d during the dry period. The total discharge of springs averages 500,000 m3/d in Estonia forming about 7% of the total groundwater discharge to the channel network.
The total groundwater discharge from the subzone of fast flow directly to the sea is 700,000 m3/d. In Estonia, the total lengths of the channel network and coastline are 31,000 and 3800 km, respectively (Hang & Loopman 1995). Thus, the intensity of groundwater discharge per a unit length of the channel network is about 250 m3/(d×km) and per a coastline unit - 180 m3/(d×km). These parameters are in good accordance and prove the reliability of the performed water budget calculations.
In northeastern Estonia, the subzone of fast flow enfolds the oil shale mines where the groundwater head has withdrawn up to 60 m. As a result of dewatering of mines, drawdown occurs in an area of about 700 km2 and the induced recharge of Ordovician aquifers from the channel network reaches 400,000 m3/d.
Of the total amount of groundwater formed in Upper Estonia, 430,000 m3/d flows laterally into Lower Estonia; of that ca. 170,000 m3/d into the Fore-Klint Lowland. The groundwater discharge directly to Lake Peipsi is 150,000 m3/d.
Actual velocity of groundwater movement is variable in the subzone of fast flow. During wet periods, in the outcrop of the carbonate bedrock a lot of intermittent gravity springs, recharged from flows of shallow groundwater, come into being. Experiments carried out by means of dyestuffs have shown that the velocity of such groundwater movement is up to 5000 m/d. In deeper layers the actual velocity of groundwater v is much smaller and it is calculated by the formula v = IK/n, where I is hydraulic gradient, K is conductivity, and n is porosity.
In the above lateral water-conducting zones of the carbonate bedrock, the actual groundwater velocity v predominantly ranges from 1 to 10 m/d under natural conditions. In transversal fissures connecting these zones it is mostly 0.001...1 m/d. Consequently, in the karstified carbonate bedrock of high permeability it may take a month for the polluted groundwater to reach from the ground surface to a depth of 30 m, where usually the cased portion of a well ends. During another month, it may cover some 300 m in a lateral direction. After a year, the pollution may be found at a distance of about one kilometre from the pollution source. Approximately such velocities of groundwater movement have been observed in cases of the oil pollution on the outcrop of carbonate bedrock. In extremely permeable aquifers, groundwater may flow from the surface to a depth of 30...60 m even in a few days. However, in the Silurian-Ordovician carbonate bedrock the actual lateral velocity of groundwater movement ranges from 0.1 to 3 m/d. It is lower in Lower Estonia and in deeper layers.
In South-Estonian highlands, the hydraulic gradient of downward groundwater flows is between 0.01...0.4, but the gradient of lateral flows ranges from 0.0001 to 0.01. The lateral velocity of groundwater is 0.02...0.2 m/d in sandstones; the transversal velocity ranges from 0.001 to 0.005 m/d. In loamy till, the velocity of groundwater movement usually does not exceed 0.001 m/d, but in glaciolacustrial or glacial sandy loam it is up to 0.1 m/d. The velocity of groundwater movement is 0.0005...0.001 m/d in peat, 0.001...0.15 m/d in sand and 10...15 m/d in gravel.
A s i n g l e w a t e r e x c h a n g e may take place in some short and highly permeable branches of local flow system during a couple of months. Sometimes such branches with a length of up to 0.5 km occur between the local divide and the nearest stream. If the permeability of flow system is moderate or low, then under the same conditions a couple of years are required for a single water exchange. In the case of very low permeability of local flow system, when conductivity is 0.01...0.1 m/d (loam, peat, fine sand), it may take 50...150 years.
The time needed for a single water exchange has been calculated assuming that the existing groundwater will be replaced gradually along the branches of flow system which are isolated from one another. Actually, a more or less intensive water exchange takes place between all adjacent branches of groundwater flow. The most intensive water exchange occurs across the groundwater table directly affected by infiltration and evaporation. If an unconfined aquifer consists of Quaternary deposits or sandstone, the average thickness of which is 10 m and porosity 0.2, then at a common infiltration rate of 200 m3/(d×km2), the single water exchange takes about 25 years. Under the same conditions, in the carbonate bedrock with the porosity of 0.02 it takes only 3...4 years.
The s u b z o n e o f m o d e r a t e f l o w is recharged from the overlying subzone of fast flow. The amount of this downward flux is 400,000 m3/d, i.e. only about 4% of total net infiltration. Of that, 160,000 m3/d leaks through the Silurian-Ordovician regional aquitard with average intensity of 9.5 m3/(d×km2) in northern Upper Estonia. The intensity of downward fluxes penetrating the tectonic faults and recharging the underlying Ordovician-Cambrian aquifer system reaches 25...50 m3/(d×km2) or even more in the Pandivere Upland. The portions of the Middle-Lower Devonian aquifer system and the underlying Silurian aquifers which both belonging to the subzone of moderate flow are recharged through the Narva regional aquitard in South-Estonian highlands. In that case, the total amount of downward fluxes is 240,000 m3/d and average intensity 20 m3/(d×km2). The lateral groundwater flux from Upper to Lower Estonia along the subzone of moderate flow is 220,000 m3/d. Of that, 200,000 m3/d returns into the overlying subzone of fast flow by uprising filtration. The flow to shelf deposits and from there into the sea averages 100,000 m3/d, almost an equal amount goes to the underlying subzone of slow flow. Pumpage from the subzone of moderate flow was about 74,000 m3/d in 1995.
The length of lateral groundwater flow branches belonging to the intermediate flow system ranged from 50 to 250 km under natural conditions, the hydraulic gradient was mostly 0.0003...0.0004. The actual velocity of groundwater movement was probably 0.005 m/d in Devonian and Cambrian sandstones and 0.05 m/d in Silurian carbonate rocks. At such velocities, a complete water exchange could have taken place only in the Silurian layers during the last 10,000 years, i.e. since the time the ice sheet retreated from Estonia’s territory (Raukas & Rõuk 1995). During the same period, in sand- and siltstones the groundwater could have moved forward only some thirty or forty kilometres.
Owing to the intensive pumpage, a piezometric depression has formed in the subzone of moderate flow. Local cones occur in Haapsalu, Paldiski, Vasalemma, Tallinn, Kohtla-Järve, Pärnu, and Tartu where the drawdown ranges from 20 to 94 m. At the time being, the groundwater moves in many different directions towards the intakes in the subzone of moderate flow. The hydraulic gradient is 0.005...0.01 or even higher (Fig. 98). The actual velocity of the groundwater movement predominantly ranges from 0.1 to 2 m/d.
The z o n e o f p a s s i v e w a t e r e x c h a n g e under the subzone of moderate flow enfolds the Silurian-Ordovician regional aquitard and all underlying strata south of the Tartu latitude (Fig. 101). Farther in the north, it comprises the Lükati-Lontova regional aquitard, the Cambrian-Vendian aquifer system, the water-bearing portion of the crystalline basement, and in the West-Estonian Archipelago partially also the Ordovician-Cambrian aquifer system. In the zone of passive water exchange, the water moves at a considerably lower actual velocity than in the zone of active water exchange. Therefore, the above-mentioned relatively thick regional aquitards belong to the zone of passive water exchange. Under natural conditions, the zone of passive water exchange recharges from overlying strata in South-Estonian highlands only, in the Pandivere Upland such recharge is negligible.
The s u b z o n e o f s l o w f l o w embraces the upper portion of the zone of passive water exchange to a depth of 350 m below sea level. Under natural conditions, the downward filtration from the overlying subzone of moderate flow into the subzone of slow flow averaged 30,000 m3/d which was only 0.3% of the total net infiltration. The uppermost branches of the intermediate flow system with a length of up to 50 km, which belonged to the subzone of slow flow rised up in the region of Lake Peipsi recharging the subzone of moderate flow from below. The lower branches headed along the regional flow system towards discharge areas in the depressions of the Gulf of Finland or central Baltic Sea. The lateral hydraulic gradient of deep groundwater flows ranged from 0.0001 to 0.0003.
The calculated velocities of deep groundwater movement are between 0.0005...0.005 m/d under the above-described conditions, which means that during the last 10,000 years the deep groundwater could have move forward only by some twenty or thirty kilometres and a complete water exchange along flow branches was impossible. This viewpoint is proved by the isotopic composition of Cambrian-Vendian fresh water which shows values of d18O from -1.8 to 2.2% (Vaikmäe & Vallner 1989). Such water must have originated from the thawing of the ice sheet at the end of the Pleistocene and its preservation in the subzone of slow flow is only due to an extremely low velocity of groundwater movement.
At present, pumpage from the subzone of slow flow is about 110,000 m3/d. Pumping wells are mostly situated in the coastal area of northern Estonia within 20 km from the sea. Pumpage is most intensive in Tallinn and Kohtla-Järve where the local centres of piezometric depression have formed and the maximum drawdowns reach 25 and 50 m, respectively (Figs. 99, 102). At the present time, the water moves to the centres of piezometric depressions in the subzone of slow flow. North of groundwater intakes the direction of flows is from the sea to the mainland, i.e. contrary to that in predevelopment conditions. Therefore, an encroachment of brackish sea water into coastal aquifers is taking place in the near-shore area of northern Estonia.
The s u b z o n e o f v e r y s l o w f l o w enfolds the lower portion of the zone of passive water exchange at a depth greater than 300 m below sea level. This subzone includes the lower strata of the Ordovician-Cambrian and Cambrian-Vendian aquifer systems (Fig. 101) south of Elva latitude which comprise water with TDS ranging from 1 to 22 g/l. Due to the lack of experimental data, the velocity of groundwater movement in this subzone is not yet clear. In the lowermost portion of the Estonian water-bearing formation which lies at a depth of 500...700 m below sea level in the vicinity of Ruhnu Island (Fig. 94), the water may be stagnant (Mazor 1995, Mazor et al. 1995). In any case, the water has not become fresh in the subzone of very slow flow during the postglacial period though the TDS of water might have decreased.
In Estonia, the total net infiltration enters first the Quaternary cover (Fig. 103). The downward flow from the Quaternary deposits into the underlying bedrock averages 5,300,000 m3/d, while the direct discharge into the channel network is 3,000,000 m3/d. Discharge through springs is 43,000 m3/d and 320,000 m3/d flows directly to the sea. The flux of bedrock water rising upward and recharging the Quaternary cover from below is 4,100,000 m3/d.
The amount of water flowing from the Quaternary cover into the Upper and Middle Devonian aquifer systems reaches 960,000 m3/d and the recharge from the underlying Middle Devonian aquifer system is 510,000 m3/d. About ¾ of inflow discharges in the form of lateral flows to the channel network across the streambeds. The remaining amount discharges through springs, is extracted by pumping or leaks into deeper layers.
The Middle-Lower Devonian aquifer system is recharged from the Quaternary deposits on its outcrop and farther in the south from the overlying Middle Devonian aquifer system. The total inflow is 350,000 m3/d of which some 80% rises upwards into the overlying Middle Devonian aquifer system or discharges into streams through the Quaternary deposits. Approximately 15% is pumped out or drained directly by the sea, while 5% goes into the underlying Silurian-Ordovician aquifer system.
Of the total downward flow formed in the Quaternary cover, the Silurian-Ordovician aquifer system receives 75% or 4,000,000 m3/d. The induced recharge of the carbonate bedrock due to the dewatering of mines is 500,000 m3/d. The upward recharge from the underlying Ordovician-Cambrian aquifer system reaches 70,000 m3/d in Lower Estonia. Discharges to the channel network through Quaternary deposits and springs are 3,000,000 and 450,000 m3/d, respectively. Pumpage from the Silurian-Ordovician aquifer system, including mine water, averages 760,000 m3/d and discharge into the sea through the shelf is 300,000 m3/d.
The Ordovician-Cambrian aquifer system is recharged mainly from downward flow reaching 150,000 m3/d which comes from the overlying Silurian-Ordovician aquifer system. Of the above amount of inflowing water, one third recharges the underlying Cambrian-Vendian aquifer system, one third rises up into the overlying Silurian-Ordovician aquifer system in Lower Estonia and one third is pumped out or flows into the sea.
Pumpage from the Cambrian-Vendian aquifer system was approximately 110,000 m3/d in 1995. About half of it leaks through the Lükati-Lontova aquitard from the Ordovician-Cambrian aquifer system, and another half has formed on account of lateral flows coming from the side of the Gulf of Finland and central Estonia.
Natural fluctuation of the groundwater table has been investigated in many places all over Estonia. The statistical analyses of observation data have shown (Vallner 1982) that the groundwater table in the Quaternary deposits covering the outcrop of the Devonian rocks is 0.1 m higher in the beginning of the year than the annual average level of the water table. Later on, the groundwater table is lowering more or less evenly until the beginning of March when it is below the mean level by 0.05 m. When the air temperature rises above 0oC, the melt-water will percolate into the soil and the spring phase of intensive infiltration starts, lasting until snow has melted in the last decade of April. The amplitude of the spring rise of the groundwater table is 0.4 m whereby the maximum point exceeds the annual mean level by 0.35 m. In late spring and in summer, the amount of groundwater mostly decreases due to the intensive evapotranspiration of the soil moisture. As a result, the groundwater table will lower during about 140 days until the first or second decade of September when the decrease of evapotranspiration caused by lowering of the air temperature will be balanced with infiltration. The amplitude of groundwater level lowering is 0.6...0.7 m in the warm period. The minimum point of the groundwater table is below the annual mean level by 0.3 m in September. The summer lowering of the groundwater table can be retarded or even changed to rising because of occasional rain periods. Intensive infiltration recurs in autumn when a systematic rainfall starts and evapotranspiration is low. Then the groundwater table rises until the soil will freeze in December. During a cold winter the amount of groundwater predominantly decreases due to restricted infiltration.
The fluctuation of the groundwater table in western Estonia is very similar to that in southern Estonia as described above. Only the fluctuation amplitudes in southern Estonia are by 0.1 m less in winter and spring. It may be explained with the higher air temperature in winter, owing to which a portion of melt-water percolates into the soil or discharges directly to stream channels before the main thawing period starts in spring.
The annual amplitude of the groundwater table fluctuation is significantly greater in the karstified and fissured carbonate bedrock of northern Estonia. The spring rising reaches 0.7 m and the summer lowering is about 0.9 m. In the areas of abundant karst phenomena, the annual amplitude of fluctuation can range from 4 to 6 m, exceeding occasionally even 10 m. The spring maximum point and the summer minimum point arrive by one decade earlier than in southern Estonia. Such relatively great fluctuation amplitudes are caused by the karst cavities in the carbonate bedrock which accumulate a significant amount of water in spring, but this water quickly discharges to streams in summer.
Seasonal fluctuations of the water table are remarkably small in peat. The spring rising amplitude does not exceed 0.1 m and the summer lowering averages 0.2 m. The winter lowering lasts until the third decade of March and the spring maximum point arrives in the middle of May.
In confined bedrock aquifers the seasonal fluctuation of the head is commonly similar to the fluctuation of the groundwater table, but the amplitude decreases with depth. Owing to the intensive pumpage, the character of natural seasonal fluctuation can be more or less perverted.
V. Karise
Infiltration water, comprised in the active water exchange zone of the Estonian groundwater system, obtains the chemical composition typical of groundwater mostly in the aeration zone (Table 19, Fig. 95). The upper 30...50 metres of the active water exchange zone are characterised by oxidized state, while in the lower part a transition from oxidizing to reducing conditions takes place (Põllumajanduslik ... 1994). The passive water exchange zone is entirely under reduced state.
In the active water exchange zone, calcium and magnesium carbonates are practically the only dissolved compounds. Therefore, regardless of the lithological compositon and the redox state of the groundwater the HCO3-Ca-Mg (frequently also HCO3-Mg-Ca) type of groundwater is formed, with the content of dissolved mineral salts under natural conditions being 0.1...0.6 g/l, most commonly 0.3-0.4 g/l. The content of free CO2 in the upper part of the active water exchange zone is prevailingly 20...30 mg/l (the boundary content limits are 0.5...50 mg/l, occasionally even 100 mg/l). With the pH values 7.2...7.6, the concentration of balanced HCO3- in the water is 200...400 mg/l. When carbonates dissolve, then together with HCO3- also Ca2+ and Mg2+ reach water in proportional amounts (most frequently 40...95 and 11...30 mg/l, respectively). Besides, the water is enriched with Na+, Cl- and SO42- (2...20, average 10 mg/l) originating from precipitation or soil.
In the active water exchange zone, neither the groundwater composition nor the amount of dissolved mineral salts is controlled by the lithological composition of rocks. Evidence is derived from the uniform chemical composition of water in springs and the amount of mineral salts dissolved in water. As an exception serves the water stored in the Quaternary sands, in which the proportion of quartz reaches 90%, and also the water stored in Devonian sandstones in southern Estonia in the areas where the Devonian is covered by sand, not by till. In that case the total content of dissolved salts (TDS) may reach 0.1...0.2 mg/l. This kind of water is unsaturated with carbonates, and has maintained the potential ability of dissolving carbonates. In practice, the cases are known, when groundwater in sandy areas has corroded concrete well curbs. This kind of water is used for removing scale from steam boilers.
Groundwater in the active water exchange zone is generally weakly alkaline (pH = 6.8...7.6). The unsaturated water in sandy regions is slightly acid (pH = 5.5...6.5). The water in bogs is acid (pH = 3.0... 5.0).
In the upper part of the active water exchange zone, the groundwater always contains free oxygen (O2). Its content is much the same as in the surfce water, which in Estonia is 8...12 mg/l, as an average (Simm 1975). Due to the presence of free oxygen, the redox potential (Eh) of the water is always positive, because even with a small amount of free oxygen available, Eh cannot be less than +0.17...+0.18V (Posokhov 1975). The higher the concentration of free oxygen in water, the higher the redox potential (maximum +0.7V). Generally, the water under oxidized state does not contain iron, because Fe-oxides and -hydroxides are insoluble in water, which means that Fe comprised in water-bearing rocks does not reach groundwater. However, in several cases Fe2+ (0.7...5.0 mg/l) has been determined in the water of bored wells tapping the Devonian sandstones or Silurian and Ordovician carbonate rocks, and in several cases also in the water of springs flowing out from Devonian sandstones which is indicative of reduced state at that depth. Aqueous environment, where pH = 7.0 and Eh < +0.3 V, is reducing in respect of iron. The latter stays dissolved under such conditions, although there is a small amount of free oxygen available in the water (Carrels & Christ 1965, Shvartsev 1982).
In natural, uncontaminated groundwater which is in oxidized state the content of NO3- is commonly 5...6 mg/l. In the spring water flowing out from Devonian sandstones the concentration of NO3- is only 1...3 mg/l, in the water of excessively damp areas and peatlands it is less than 1 mg/l. This is due to reducing conditions under which, as a result of denitrification, part of the initially dissolved NO3- has been reduced to free oxygen (N2) which volatalizes (Põllumajanduslik... 1994).
In northern Estonia, in the areas with a thick (up to 100 m) Quaternary cover where sediments contain buried organic matter, emissions of burning gas from bored wells have sometimes been recorded (Keri, Prangli and Mohni islands, Viinistu, Püssi). The gas comprises methane, hydrogen, nitrogen, hydrogen sulphide and, to a lesser extent, also helium, argon, oxygen and carbon dioxide and other compounds (Voytov et al. 1982). Frequently, there are gas emissions from the bored wells tapping the bedrock. In this gas the main component is free nitrogen (N2) which accounts for 80...90% of gas volume. Sometimes, hydrogen sulphide is emitted together with the gas. In places, particularly in northern and southwestern Estonia, the water of borings contains helium in a rather high concentration (up to 0.6 ml/l) which originates in the crystalline basement and reaches groundwater through tectonic disturbances (Tibar 1987).
A great part of the Estonian water-bearing formation is situated in the passive water exchange zone with reduced state. Due to combined effect of several factors, such as the very slow groundwater flow, connate water occasionally present in rocks, soluble mineral salts, etc., the chemical composition and TDS in groundwater in the passive water exchange zone differ with regions. The HCO3-Mg-Ca water with the TDS ranging from 0.5 to 0.6 g/l is of limited distribution and occurs in the Middle-Lower-Devonian aquifer system along Pärnu - Tartu - Viljandi line. The HCO3-Cl-Na-Mg-Ca water with the TDS in between 0.3...1.5 g/l is more widespread. It occurs in central and southern Estonia in the Middle-Lower Devonian, Silurian-Ordovician and Ordovician-Cambrian aquifer systems. The Cl-HCO3-Na-Ca and Cl-HCO3-Ca-Na water with the content of dissolved mineral salts 0.4...1.0 g/l is stored in the Cambrian-Vendian aquifer system in northern Estonia.
In several places around Tallinn, the content of δ18O in this water is -18...-22‰, but in the area of Loksa, Võsu, Kunda, Toila and Salutaguse it is -11.9...-16‰ (Vaikmäe & Vallner 1989, Savitskaja & Viigand 1994). In all likelihood, in the former case we have a typical glaciogenic palaeowater, in the latter case the glaciogenic water seems to have mixed with infiltration water formed under the conditions of moderate climate.
The Cl-Na, Cl-Na-Ca and Cl-Ca-Na water with the TDS from 2 to 22 g/l is widespread in the passive water exchange zone. This type of water has been established in the Ordovician strata on Hiiumaa Island (Kärdla), Ordovician-Cambrian aquifer system in southern, Cambrian-Vendian aquifer system in northeastern, southwestern and southeastern Estonia, on the islands of Saaremaa and Ruhnu and in the Lower Proterozoic strata in northern and southwestern Estonia.
The SO4-Cl-Ca-Na water with the TDS up to 4.6 g/l, which very rarely occurs under such conditions as prevailing in Estonia, is found at a depth of 260 m at Värska in the Middle-Lower Devonian aquifer system (Fig. 94). Formation of sulphate-rich groundwater is due to the occurrence of gypsum in those layers south of Värska. Since the depth interval 250...800 m is charactericed by reduced state, the water there does not contain O2 or NO3-; SO42- is absent or present in a very low amount (1...5 mg/l). The content of Fe2+ is occasionally very high (up to 10 mg/l) and pH of deep water is usually 7.8...8.5. One reason is the low concentration of free CO2 (often 0...2 mg/l, occasionally up to 6 mg/l). The content of HCO3- is also low, commonly 90...120 mg/l, seldom more.
Total hardness of Estonian groundwater differs in a wide range: it is soft in non-carbonaceous sediments, hard or very hard in carbonate-rich strata. Total hardness of bog water is 1...25 mg/l as CaCO3, the hardness of water stored in sands under the influence of bog water is 20...85 mg/l, in extensive sand areas 65...195 mg/l and in sedimentary rocks 145...360 mg/l. Total hardness of mineral water with the TDS reaching 22 g/l, may be up to 4250 mg/l.
The concentration of microelements in the water stored in oxidized state in the active water exchange zone is generally very low. The low content of some physiologically important microelements, such as fluor and iodine in drinking water, may cause health troubles. The incidents of endemic struma and caries, in some regions more numerous than in others, are associated with the low content of iodine and fluor in drinking water, respectively (Kuik 1961). The groundwater, which occurs under reducing conditions in the passive water exchange zone, is richer in microelements. In some places the content of microelements is even in excess of the optimum value established for drinking water. For instance, in the Silurian-Ordovician aquifer system the content of fluorides (F-) is 5.5...7.2 mg/l in western and southwestern Estonia, 3.2 mg/l in Tartu and 2.4 mg/l at Abja, being well in excess of the standard established for the drinking water in Estonia and exceeding the level permitted by the World Health Organisation, which is 1.5 mg/l (Kuik 1963, Viigand & Vatalin 1992, Guidelines... 1993). This explains the incidences of fluorosis in these areas (Kuik 1961). In Pärnu, the content of fluorides in the Silurian-Ordovician aquifer system is almost optimal - 0.8...1.0 mg/l (Boldõreva et al. 1993), being elsewhere well below it.
In the Ordovician-Cambrian aquifer system, the content of microelements is low in the areas where the TDS is less than 1.0 g/l. South of the Pärnu - Viljandi - Tartu line where the Ordovician-Cambrian aquifer system stores Cl-Na-(Ca) water with the content of dissolved mineral salts up to 14 g/l (Ikla), the content of microelements is higher. Thus the contents of cadmium (Cd2+), lead (Pb2+) and lithium (Li+), in excess of the norms established by the drinking water standard in Estonia, have been registered at Värska (Fig. 94). Since the water derived at Värska is used as a mineral water, and half of the amount produced is dilluted, then it does not pose any threat to human health.
In the groundwater stored in the Cambrian-Vendian aquifer system east of Tallinn, the concentration of iodides (J-) is heightened, reaching 120...280 μg/l (Kuik 1961). This water, if used for drinking, will cover the need for iodine in the population. Generally, the content of all microelements increases in northern Estonia towards the east in this aquifer system. In northeastern Estonia, in some bored wells the content of cadmium (Cd2+), lead (Pb2+) and lithium (Li+) is slightly in excess of the norms. In this aquifer system the concentration of microelements is at its highest at Värska in southeastern Estonia where in Cl-Na-Ca water the TDS ranges from 6.0 to 19.0 mg/l. In this water the content of cadmium (Cd2+), lithium (Li+), manganese (Mn6+) and lead (Pb2+) is in excess of permitted boundary limits. This type of Värska water is not used directly for drinking, but due to its medicinal effect it is used in baths and as curative drinking water. The origin of the microelements in the water of Ordovician-Vendian and Cambrian-Vendian aquifer systems is not yet unambiguously clear. They may partly reach the water from the steel casing of wells, continuously corroded by salt water stored in reduced state, however, part of microelements is evidently of natural origin (Savitskaja & Viigand 1994).
In several regions of Estonia, the heightened concentrations of bromides (Br-) have been detected (Fig. 94, Table 20):
- in the Ordovician-Cambrian aquifer system: 31 mg/l at a depth of 645...658 m at Ikla and 50...54 mg/l at a depth of 707...784 m on Ruhnu Island;
- in the Cambrian-Vendian aquifer system: 13 mg/l at a depth of 540...555 m at Kuressaare, 16...17 mg/l at a depth of 520...535 m and 51...56 mg/l at a depth of 540...600 m at Värska;
- in the water of the crystalline basement 51...61 mg/l at Hirvli and Pudisoo.
The groundwater in which the TDS is 2 g/l or more is rated as mineral water in Estonia. A. Verte was the first to predict the existence of different types of mineral water in Estonia (Photo 33). The first mineral water deposit was discovered at Pärnu in 1959 by the researchers of the Geological Survey of Estonia. For the purposes of structural geological investigations a well, deeper than 500 m, was sunk in Pärnu. It tapped the Lower Proterozoic crystalline bedrock and yielded Cl-Na water with the TDS about 22 g/l (Vingisaar 1978). Mineral water has been found in 16 different sites all over Estonia. At Värska, it occurs in four aquifers, at Kuressaare and Arumetsa in two aquifers (Table 20, Fig. 94).
In Estonia the bottling and marketing of mineral water was started in 1968. Currently, the Cl-Na-Ca water of the Ordovician-Cambrian aquifer system obtained at Värska with the content of TDS ranging from 2.0 to 2.2 g/l is bottled. The mineral water derived from Värska and Kuressaare is used for curative purposes both for drinking and in baths. In earlier years, the mineral water derived at Arumetsa, Häädemeeste, Ikla, Kuressaare and Kärdla was also bottled and sold. During the bottling, the water is often enriched with carbon dioxide (5...7 g/l). The reserves of Estonian mineral waters are estimated at about 6000 m3/d. Total reserves of salty and salt water are very large in Estonia and may amount to hundreds of cubic kilometres.
The temperature fluctuations caused by meteorological factors occur in the upper part of Estonia’s water-bearing formation with an average thickness of some 18 m (Jürima 1984). The maximum thickness of this zone, marking the depth of the so-called neutral layer, reaches 30 metres in the Pandivere Upland. In western Estonia and on islands, the thickness of the zone subjected to annual temperature fluctuations is 10...15 m and in the coastal plain of northern Estonia it is 5...11 m. The temperature of water in springs and up-to-30-m-deep wells in the Pandivere Upland ranges from +4oC to +6oC, elsewhere it is +6...+7oC. The water is coldest in March- April when a great quantity of melt-water percolates into the ground, and warmest in September - October. The temperature in the neutral layer is stable, being prevailingly +7oC, in some places also +6oC. Downwards the temperature rises steadily (Fig. 104). At a depth of 50 m in the area of Silurian and Ordovician carbonate rocks, the temperature is +6.2...+6.3oC, in southwestern Estonia under Devonian sandstones +9.4...+9.5oC and in Lontova clays on Estonia’s north coast +7...+8oC. At a depth of 100 m the average temperature is +7.6oC. In the uplands with intensive groundwater recharge the temperature is lower than the average (+6.5...+7.2oC), whereas in northeastern Estonia and on the Pärnu Lowland it is higher, being +9.2...+9.5oC and +9.8...+10.2oC, respectively. At a depth of 200 m in the carbonate bedrock of central Estonia, the temperature is only +8.0oC which shows that the subsurface groundwater percolates quickly to a greater depth. At the same depth in northeastern Estonia, the temperature is +14.0...+14.5oC and in southwestern Estonia +10.3...+10.6oC. At this depth, the highest temperature (+15...+16oC) has been measured in the crystalline basement of northeastern Estonia (Jürima 1984).
The mean value of geothermal gradient in Estonia is 1.2oC/100 m, and it increases with depth being 1.0oC/100 m in Silurian-Ordovician carbonaceous rocks and Devonian sandstones, 2.0...3.5oC in deeper bedded Cambrian sandstones, and 5.0...6.0oC (average 4.0oC) per 100 m in the underlying Cambrian clays and silts (Lükati-Lontova regional aquitard) (Jürima 1984).
L. Savitskaja
As a result of extensive economical activity and high vulnerability of the uppermost aquifer, the shallow groundwater is in places heavily polluted and therefore unfit for drinking. The point-pollution sources are different constructions and pipelines in poor condition, such as boilerhouses, fuel storages, storages of chemicals and manure, settling basins, sewerage, leaching beds, gas stations, landfills, burial places of domestic animals, etc. Extensive fuel leakages have occurred on military airfields (Fig. 105) and in railway junctions. The majority of gas stations that were state-owned during the Soviet period contaminate environment with oil products. The asphalt concrete plants using primitive equipment and technology (Tiitso, Riisipere, etc.) are also sources of extensive pollution of groundwater.
In the Ida-Viru County, essential point-pollution sources are the spoil heaps of oil shale mines and ash plateaus of thermal power plants (Fig. 105). The groundwater leaching from oil shale ash is polluted with phenols and compounds of heavy metals, its pH-value is 12 and even more. Extensive surface water and groundwater pollution has been caused by cracking processes accompanying fires in oil shale mines (Vallner & Sepp 1993, Vallner 1994).
Non-point pollution is caused by the misuse of mineral fertilizers on arable lands and the underutilization of slurry from pig farms, but also by the treatment of fields with toxic chemicals. In the Ida - Viru County powerful thermal power plants annually eject into the atmosphere 120,000 tonnes of fly-ash and 80,000 tonnes of aerosol fractions containing harmful elements (S, F, Cl, V, Cr, Ni, Br, Sb, Cd, Pb, As) and radioactive isotopes (Õispuu & Rootamm 1994). The ash and gaseous pollutants transported by wind over large areas percolate together with rainwater into the groundwater. Higher concentrations of these pollutants in the atmosphere promote formation of sulphuric and nitric acids. The impact of acid rains is mitigated by fly-ash from thermal power plants which may spread to a distance of 100...150 km (Frey et al. 1987).
In 1994, about 14,000,000 tonnes of solid waste was generated in Estonia; of that amount 46 % by the oil-shale-based energy production, 40 % by oil shale mining, 8 % by the chemical industry and only 6 % in other spheres (Keskkond... 1995). Of 1,962,000,000 m3 of waste water produced, 70 % was the cooling water heated through use in an industrial process which raised the temperature of both surface and groundwater. The amount of waste water needing purification was 380,000,000 m3. Approximately 1% of waste water (1,800,000 m3) was discharged directly into soil and groundwater. Compared to 1989-90, the total industrial capacity has decreased and the technology has improved. As a result, the amount of solid waste and waste water has decreased by about 20 and 40 %, respectively.
In 1994, a total of 41,000 tonnes of mineral fertilizers (as N) and 1,100,000 tonnes of manure were used in Estonia, which is about 6.5 times less than during 1988-89. The average pollution load of solid waste was of 320 t/km2, and that of untreated waste water 8,700 m3/km2. The amount of mineral fertilizers (as N) was 0,04 t/ha and that of manure 1.1 t/ha. The pollution load was highest in the Ida - Viru County due to oil shale enrichment waste and ash of thermal power plants. In 1994, the average pollution load of solid waste in this region was 3,840 t/km2 (1.6 times less than in 1990), while the load of waste water in need of purification was 69,000 m3/km2.
In northeastern Estonia, the Ordovician carbonate rocks covered with a thin Quaternary mantle are intensively polluted with shale oil and phenols. In the undermined area (about 200 km2) the lowering of groundwater level has resulted in oxidation of pyrite in the aeration zone, due to which the content of sulphates in groundwater has increased up to 650 mg/l (under natural conditions it is less than 20 mg/l).
Groundwater is polluted with oil products at Tapa (Fig. 105) and in its surroundings. Since 1966, several big accidents of fuel tanks and constant leakage of fuel pipelines took place on the former military objects of this region. Particularly intensive pollution with oil products has been recorded on the former military airfield, where the stopcocks, valves and pipelines of fuel tanks were leaking continuously. Near the railway oil receiving centre, an oil lake had formed on the surface. Due to carelessness or leakage from boiler-houses and fuel tanks, groundwater is polluted with oil products also in Tallinn, Tartu, Rakvere, Kohila, Rapla, Tamsalu, Aruküla and several other places. In these regions, the water supply is more or less disturbed.
In the area of the former galvanic departments and in the surroundings of landfills, groundwater contains heavy metals. Due to the diffusion of filtrates from Tuula landfill near Keila, the concentrations of hazardous compounds in the neighbouring wells are the following: Pb2+ – up to 0.13 mg/l, Mn2+ – up to 1.4 mg/l, Cd2+ – up to 0.0035 mg/l (Tennokesse et al. 1992); the maximum permissible concentrations of these elements in drinking water are 0.01, 0.1 and 0.003 mg/l, respectively (Eesti standard...1995).
In the regions of agricultural activity, groundwater is contaminated mainly by nitrogen compounds (Fig. 106), but the concentration of chlorides and sulphates has also increased to 40...60 mg/l (under natural conditions 20 mg/l), as an average. In the surroundings of Viiratsi piggery near Viljandi, the concentrations of ions are the following: NH4+ – up to