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Palaeoseismicity and De Geer Moraines
Quaternary International 68}71 (2000) 175}186
Palaeoseismicity and De Geer Moraines Jan Lundqvist* Department of Quaternary Research, Stockholm University, S-106 91 Stockholm, Sweden
Abstract De Geer moraines are small moraine ridges formed partly in crevasses at the base of glacier ice. There may be several di!erent origins of the crevasses. One possibility, proposed by Gerard De Geer who "rst studied these ridges, was seismic activity in connection with the isostatic rebound during the deglaciation. The present study shows a close connection between occurrences of De Geer moraines in Fennoscandia and seismic activity or fresh-looking fault lines and other features indicating neotectonic movements. This supports De Geer's theory but does not exclude formation in crevasses formed in other ways. There is no real proof that the tectonic movements took place in the Quaternary but the close connection between the moraines and tectonic features is evidence pointing in this direction. If this is correct it has important implications for the stability within cratons of Precambrian crystalline bedrock. 2000 Elsevier Science Ltd and INQUA. All rights reserved.
1. Introduction De Geer moraine is a term used for a type of small moraine ridge discussed by Gerard De Geer (e.g. De Geer, 1940). These ridges are usually a few 100 m in length, not more than some 10}20 m broad and a few metres high. Characteristically, they occur in large swarms or clusters within which the ridges show a general trend parallel to the ice margin. The orientation of individual moraines varies within wide limits. Ridges may even be perpendicular to the margin or winding and anastomosing. The following discussion relates only to small moraines that correspond to this de"nition * which actually applies to virtually all areas with small moraines in Fennoscandia. Fig. 1 shows all such areas within Sweden. Isolated small moraine ridges, probably formed in di!erent ways, are not included in the map, and this applies also to moraines in the mountain area and within the Younger Dryas and other end-moraine belts. In the latter case, they obviously belong to the endmoraine systems, although smaller than the typical end moraines. There has been many di!erent ideas as to the origin of these ridges. De Geer (1940) considered them end moraines, and deposited mainly annually at the ice margin. Individual ridges could also be formed between the an-
* Fax: #46-8-6747895. E-mail address: [email protected] (J. Lundqvist).
nual ridges due to calving, as later con"rmed by MoK ller (1962). Consequently, they have often in the literature been called annual moraines or just end moraines * even the term terminal moraines has been used. Hoppe (1957) and StroK mberg (1965), contrary to De Geer, interpreted them as formed some distance inside the ice margin. Di!erent ideas have also been put forward with respect to the mode of deposition. The internal structures show compact basal till as well as ablation till and waterlain sediments. Folded inclusions of clay and other proglacial sediments occur. These di!erent types of sediment indicate deposition by both push, dump, squeeze, and lodgement. A possible compromise suggests that the moraines started to form in basal crevasses. These crevasses initiated later break-up by calving. More material could then be added to the ridges when they were in a marginal position. The di!erent possible processes of formation have been discussed in many papers (see, e.g., Hoppe, 1948, 1957; StroK mberg, 1965; Zilliacus, 1987) and will not be treated here. In order to avoid these problems the non-genetic term De Geer moraines was proposed by Hoppe (1959). Reviews of the discussions about formation and terminology were presented by Lundqvist (1981) and Zilliacus (1987). In Fennoscandia, the classic De Geer moraines occur in #at lowland regions in the central and northern parts of Sweden and Finland and also in some restricted areas in Norway. Generally, these moraines were formed beneath water, in glacial lakes or, usually, the sea. Moraines
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authors (see references in Lundqvist, 1981). There is some doubt that all these types of small ridges are the result of the same process of formation, and a more systematic classi"cation is needed, as emphasised by Lundqvist (1981). In the following discussion only the occurrences in Fennoscandia which are in agreement with De Geer's original description will be treated. The processes of formation, the relation to the ice margin and a number of other aspects brie#y mentioned above will not be discussed. The aim of this paper is only to consider the possible relation between De Geer moraines and seismicity in Fennoscandia. According to De Geer (1940) fracturing of the basal part of the ice by seismic activity contributed to the formation of the ridges. The reason for this activity should be the isostatic rebound after the waning of the last (Weichselian) Scandinavian inland ice. This idea has not gained general acceptance, although supported by MoK rner in a number of papers (e.g. MoK rner, 1977, 1979; MoK rner et al., 1989). In the light of subsequent discoveries about neotectonic activity in connection with the deglaciation (Lundqvist and LagerbaK ck, 1976; MoK rner, 1977, 1985; LagerbaK ck, 1979, 1990) this aspect deserves further consideration.
2. Regional distribution of De Geer moraines 2.1. Sweden
Fig. 1. Seismically active areas in Sweden (dashed), De Geer moraine areas (black) and ice-marginal lines with dates (clay-varve years/calibrated years). P marks the PaK rvie fault and L the main LansjaK rv fault, referred to in the text.
cannot be formed if the ice is #oating, nor if it is too thick, which makes the pressure against the bottom comparable with supra-aquatic conditions. In the "eld this is noticed as the limitation of moraines to certain levels. Similar ridges have been described from other areas, notably in the region of the North American ice sheets. Terms like cross-valley moraines, washboard moraines, small moraines and others have been used by di!erent
The outlines of the regional distribution of De Geer moraines within Sweden are shown in Fig. 1. This picture is compiled from available geological maps but details can be found in many special papers referred to in the following. Very generally, De Geer moraines belong to areas that were situated below the level of the sea or large lakes during deglaciation, that is, subaquatic areas. They are also restricted to certain levels in the terrain * on higher as well as lower ground they are absent. The altitude varies from area to area, but mostly the level implies formation at water depths of 100}250 m, in each area depending on the local thickness of the ice. Fig. 1 shows all clusters of De Geer moraines, with two possible exceptions, marked with X. These occurrences, described by Lundqvist (1969) and BorgstroK m (1979) are situated near each other in mountain valleys and are too small to be shown in the scale. They may be exceptions from the general rule of distribution, but it is uncertain whether they should be called De Geer moraines * although this term was used by BorgstroK m (1979). As a matter of fact, they are also situated close to a seismic area in central Norway (Stephansson, 1979). The Ma( laren area: The main area of distribution is situated around Lake MaK laren in south-central Sweden. Its eastern part is the one described and mapped in detail by De Geer in his main paper (De Geer, 1940) (Fig. 2).
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Fig. 2. The main area with De Geer moraines in east-central Sweden. Black areas"De Geer moraine clusters; grey areas"lakes and (in the east) the Baltic Sea; dentated lines"ice margin with dates in clay-varve years/calibrated years (the line 11,230/10,430 runs 2 km north of the line 10,450/11,250); thin broken lines"isobases for isostatic uplift ("gures are an arbitrary scale * in order to obtain the altitude in metres above the sea-level at the time of deglaciation approximately 25 m should be added); thick lines"faults with indications of Holocene movements; thick broken lines"fault lines without clear indication of Holocene movements; dotted lines"hypothetical tectonic boundaries of area with di!erential isostatic rebound. BS"Baltic Sea; H"Lake HjaK lmaren; M"Lake MaK laren; S"Stockholm; U"Uppsala; V"VaK stera s; OG "OG rebro.
The moraines occur almost strictly between the ice-marginal lines dated at 10,450 and 10,200 clay-varve years BP (StroK mberg, 1989; Brunnberg, 1995). Considering possible errors in the clay-varve time scale (Wohlfarth et al., 1997) and calibration of radiocarbon dates according to Kitagawa and van der Plicht (1998) this can be estimated to roughly 11,250}11,000 calendar years BP. The zone ends abruptly towards a highland in the west (Fig. 2). East of the area shown in Fig. 2, according to personal communication by Christer Persson and Lars Rudmark at the Geological Survey of Sweden, there are some small occurrences (see also StroK mberg, 1965). This continuation is not included in the map since the area had not been mapped in su$cient detail when Fig. 2 was drawn. However, these occurrences are restricted in time like in the map area. Notably, a few small moraine groups occur north of the well-de"ned area, and corresponding in time to the northernmost clusters in the VaK nern area (see below). The water depth in the area is not known exactly since no part of the area reached above the sea-level during deglaciation but is estimated to be 130}160 m. The "gures in Fig. 2 are an arbitrary scale (in meters) used by
As se and BergstroK m (1984). In order to obtain the level of the sea when the moraines were formed we should add about 25 m to these "gures. Typical for this area as well as the other areas with De Geer moraines is the clustering of ridges. De Geer (1940) called such a cluster a `moraa, using a word from the Alps meaning a concentration of big boulders. One of the clusters is situated at Bromma in the western suburbs of Stockholm. Using this as an example De Geer developed his idea of a seismic origin of such moraines. The clustering, often with elliptic contours, is well visible in Fig. 2. Moraines do occur also outside the clusters but ridges there are few and scattered. Within at least some clusters there are also concentrations of very large boulders of local provenance. The Va( nern area: As mentioned above, the MaK laren area is interrupted westwards by a highland. West of this highland there is a continuation of moraines formed at about the same time as those to the east. The moraines are clustered in a similar fashion. One of these clusters is the well-known Kilsviken (or As ra s) Bay (Fig. 3) on the eastern side of Lake VaK nern, "rst described by De Geer (1896). A detailed map was published by Ericsson and
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Fig. 3. De Geer moraines in the Kilsviken Bay, Lake VaK nern. Photo J. Lundqvist, 1996.
LideH n (1982). This is considered the best developed area of De Geer moraines in Scandinavia. North of the lake old geological maps show several concentrations of De Geer moraines. Many of them, however, have been eliminated in later mapping (Lundqvist, 1958), but nevertheless several groups of moraines do occur. There is no good dating of this moraine area but according to Perhans (1981) the dated lines from the east continue as shown in Fig. 1. This means that the moraines here were formed between 11,500 and 10,900 calendar years BP. Using the levels of the highest post-glacial coastline of Lundqvist (1958) the water depth during the formation is estimated to be about 75}130 m. Gulf of Bothnia: The third region with De Geer moraines in Sweden extends along the coast of the northern part of the Gulf of Bothnia * called the Bothnian Bay. These moraines contain large volumes of sorted, coarseclastic sediments. They occur partly together with larger ridges, called Niemisel ridges by Lundqvist (1981). As shown by HaK ttestrand (1997) these ridges should actually be classi"ed as Rogen moraine, a type that was obviously subglacially formed (Lundqvist, 1989; HaK ttestrand, 1997). This co-existence supports the idea of De Geer moraines as subglacially deposited, as was also proposed by Hoppe (1948). The pattern of ridges is in this case more regular than in most other areas. Hoppe's (1948) interpretation was that the moraines were formed in basal crevasses when the thinning ice in the Bothnian Bay became a#oat. The group of De Geer moraines at Bothnian Bay has a continuation in the Vaasa area on the Finnish side of the bay. The Finnish occurrence may suggest that more
ridges are located on the bottom of Bothnian Bay, but AndreH n's (1990) seismic studies of the sediments there give no such indication. There is no accurate dating of this group of De Geer moraines, but AndreH n (1990) dated sediments in the Bay by means of clay-varve chronology. Calibration as described above gives an age between 10,000 and 10,300 BP for the moraines. The level of the highest coastline (Fromm, 1965) indicates that the water depth at the time of formation was 200}250 m. 2.2. Finland and Norway De Geer moraines in Finland occur in the southwestern part of the country where they are concentrated in a number of clusters (Fig. 4; Zilliacus, 1987, Fig. 37). The main clusters are situated in two areas on the coast of the Gulf of Bothnia. The best developed group is the one mentioned above, in the Vaasa region. Also farther south, around and NE of Uusikaupunki (Nystad) there are similar groups. A third main group is situated outside the First SalpausselkaK between Helsinki and Lahti. Together with these main groups there are smaller clusters. Altogether, the occurrences form a triangular area between Vaasa, the As land archipelago and Hamina (Fredrikshamn) in the southeast. Beyond this triangle there are just a few small occurrences in the extreme east and north of the Bothnian Bay. The clusters are similar to the ones described above from Sweden. Notably, in the Vaasa region, as on the Swedish side of Bothnian Bay De Geer moraines occur
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ridges, but since they occur outside the First SalpausselkaK , dated to Younger Dryas, they must antedate 12,600 calibrated yr BP (cf. BjoK rck et al., 1996). In accordance with the clay-varve chronology of NiemelaK (1971) they were formed from about 13,000 BP. The age of the youngest moraines can be estimated at 10,400 calibrated yr BP (cf. StroK mberg, 1990). In Finland the topographic conditions are similar to Sweden in the areas where De Geer moraines occur, that is, a rather #at, slightly undulating landscape below the coastline at the time of deglaciation. In Norway, the topography is quite di!erent, characterised by mountainous terrain with deep valleys and fjords. However, since a few moraine occurrences have been called De Geer moraines (Sollid and Torp, 1984) they should also be brie#y considered here. On the Atlantic coast in the As lesund area, western Norway, moraine ridges have been described as De Geer moraines (Sollid and S+rbel, 1984; Larsen et al., 1991). They occur mainly on the #oor of fjords outside the Younger Dryas moraines (see Sollid and S+rbel, 1982 with map) and must thus be at least as old as the moraines in southern Finland. The water depth at the time of formation can be estimated to just some tens of metres, certainly less than 100. Larsen et al. (1991) interpreted the ridges as true frontal formations. Although called De Geer moraines, it is questionable if they are compatible with De Geer moraines to be discussed in this paper. Similar uncertainty concerns an occurrence of moraines in the Pasvik valley, northernmost Norway (Fig. 4), studied in detail by Sollid and Carlson (1984). These moraines occur in valley position and were formed a few tens of metres below the contemporaneous sealevel. On the basis of the sedimentology in the ridges they were also in this case interpreted as frontal formations. All these Norwegian moraines could be described as cross-valley moraines, although according to the current less precise terminology (cf. Lundqvist, 1981) it is correct to include them among De Geer moraines.
3. Evidence of neotectonic activity
Fig. 4. De Geer moraine areas (black) in Finland according to Zilliacus (1987), isobases for the present isostatic uplift (broken lines, "gures in cm/yr) according to KaK aK riaK inen (1966), and fault scarps (indented lines) according to LagerbaK ck (1990). The dotted line shows the Younger Dryas ice margin.
together with larger ridges of Rogen moraine. The water depth during their formation varies from less than 100 m to more than 250 m. The time of formation is over a longer period than on the Swedish side. There is no exact dating of the oldest
The isostatic rebound after the last glaciation caused a slow and ongoing tectonic activity. The Precambrian basement in Scandinavia has been considered otherwise stable. In the last few years, however, a number of observations, partly in connection with investigations for repositories of nuclear waste, indicate that it is more subject to tectonic movements than earlier supposed. The most convincing evidence was presented in northernmost Sweden by Lundqvist and LagerbaK ck (1976), followed by detailed studies by LagerbaK ck (e.g. 1979, 1990). In Finnish Lapland similar evidence had been published by Kujansuu (1964). Although not generally accepted, indications of tectonic activity were also presented by
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MoK rner in a number of papers (e.g., 1977, 1979, 1985, 1996; MoK rner et al., 1989; also MoK rner and TroK ften, 1993; TroK ften, 1997) in the Stockholm area. In Swedish and Finnish Lapland and adjacent parts of Norway there are up to 30 m high fault scarps indicating seismic activity (Fig. 5). The area with faults can be followed from near the Caledonian mountain front in the northwest southwards to some 50 km west of the Bothnian coast at Skelleftea (Figs. 1 and 4). Some of the scarps can be very precisely dated to the time of the local deglaciation, although there is so far no possibility to obtain an age in absolute years. This applies in particular to the PaK rvie fault in northernmost Lapland (P in Figs. 1 and 4). Along this line the fault scarp is well developed. In parts it cuts beaches of ice-dammed lakes and meltwater channels formed at the ice margin. In adjacent parts such features intersect the scarp una!ected. Liqui"ed sediments and faults in the till from the last deglaciation as well as landslides are clear evidence that faulting at least partly took place after the local deglaciation. This is especially clear in the LansjaK rv area (L in Figs. 1 and 4, LagerbaK ck, 1979, 1990). The features of liquefaction and landslides even in coarse-textured till as well as the fault scarps themselves allow an estimation of the magnitude of the earthquakes related to the faulting. After a discussion of these features, LagerbaK ck (1988) concluded that faulting was accompanied by earthquakes with a magnitude of M 7.0}7.5 or probably higher. A tentative extension of known lines of ice recession (Fig. 1) based on our present knowledge of the deglaciation (e.g. Kleman et al., 1997) allows an approximate dating of these tectonic events to
10,000}10,300 calibrated yr BP. In Finnish Lapland faults may be older than this. All the North-Swedish postglacial faults, a few of which are shown in Figs. 1 and 4, are located within or close to zones of recent seismicity (Ba th, 1979; Stephansson, 1979). These zones extend from northwestern Lapland to the northern coast of the Bothnian Bay and further southward along the coast (Fig. 1). The region in Sweden where the seismic activity is strongest at present is the area around Lake VaK nern in the southwest (Fig. 1; Ba th, 1979), with modern earthquakes detected. The lake basin is bounded by fault lines in the west and east. Sandegren (1916) observed shorelines on the eastern side which he interpreted as a!ected by faulting in early Holocene time. On the northwestern side of the basin the bedrock is fractured and traversed by numerous tectonic valleys. There mostly is no evidence that movements took place after the deglaciation, but from dates of the isolation of small lake basins Risberg et al. (1996) concluded that bedrock blocks had moved di!erently after the isolation. Lundqvist (1994) pointed out that strong fracturing of the bedrock in certain zones could be an e!ect of these movements. We can conclude that the area around and north of Lake VaK nern has been subject to tectonic movements in postglacial time. The VaK nern zone extends southwards to southernmost Scandinavia and is also closely related to the Oslo graben in Norway (Huseby et al., 1978). The Stockholm area and the region around and west of Lake MaK laren is the region where De Geer (1940) proposed his theory of seismic moraines, in particular, the Bromma area just west of downtown Stockholm. The evidence upon which the theory was based was one small
Fig. 5. The PaK rvie fault in northern Swedish Lapland. Photo J. Lundqvist, 1975.
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fault scarp interpreted as recently active, and a number of very large boulders (Fig. 6) and concentrations of boulders (Fig. 7) which were obviously displaced by the gla-
Fig. 6. Large boulder transported a very short distance by the inland ice. Bromma, western Stockholm. Photo J. Lundqvist, 1981.
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cier, albeit a very short distance * just a few metres. The boulders are striated even on their basal side and so fractured and fragile that they could not possibly be transported, perhaps just tilted. New evidence has been put forward by MoK rner (1977, 1979, 1985) and MoK rner et al. (1989). Liqui"ed sediments from the time of deglaciation have been interpreted as a result of seismic activity (MoK rner, 1996; TroK ften, 1997; TroK ften and MoK rner, 1997). Seismic events were dated around 10,400 clay-varve yr BP with a strong event at 10,430, which corresponds to 11,230 calibrated yr BP (MoK rner, 1996; cf. Fig. 2). Liquefaction of clay, especially if it is rich in silt as the studied sediments, can take place due to several reasons, however, and Brunnberg (1995) did not "nd traces of such an event in his detailed clay-varve chronology. Possibly the sediments were disturbed only within narrow zones. New observations of faults give additional support to the hypothesis of seismic instability. Several fault scarps in the Stockholm area look very fresh and una!ected by glacial erosion (Fig. 8). Di!erential uplift in postglacial time has been demonstrated by Gardemeister (1999). Isobases for the isostatic rebound, ongoing as well as total from the time of deglaciation, show a marked irregularity from Stockholm some 130 km westwards (Fig. 2; MoK rner, 1977; Ussisoo, 1977; As se and BergstroK m, 1984; MoK rner, 1977, Fig. 2). New investigations by Ekman (1996) indicate that this feature is within the margin of error but if it is a reality it means that a block has moved di!erentially between fractures or #exures trending about NW}SE. Like the basin of Lake VaK nern discussed above it appears that a block of the bedrock has tilted northwards round an axis at an oblique angle to the ice margin and isobases. As such, perhaps the block
Fig. 7. Concentration of huge angular boulders on the fault line 10 km north of Stockholm centre (cf. Fig. 2). Photo J. Lundqvist, 1971.
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Fig. 8. 12 m high fault scarp, 10 km north of Stockholm (see Henkel, 1988). Photo J. Lundqvist, 1999.
was left behind during the isostatically induced general tilting southwards of the land surface. The eastern boundary of this block coincides with the area with fresh-looking faults and the best-developed concentrations of fractured rocks. In Fig. 2 the continuation northwestwards of this zone is marked with a dotted line. The western boundary has been studied in connection with the present work. The area in general is rather rich in boulders, slightly varying with the types of bedrock. However, west to south of VaK stera s (Fig. 2) there are very distinct valleys of tectonic origin. Most of them are certainly pre-Quaternary (Ahlbom et al., 1991; Bergman et al., 1996) The outcrops on the valley sides are scoured by ice. Several parallel lines are clearly visible on topographic and geologic maps. Notably, smaller fracture lines parallel to the topographically most distinct valleys show fresher, although smaller, scarps and a concentration of angular boulders that is markedly higher than at the main lines. This is most clearly observed south of the western end of Lake MaK laren (Fig. 2). There is no conclusive evidence that these faults and fractures were active in late glacial or Holocene time, but this is suggested by the similarity with the Bromma area. The zone with such lines runs about 10 km north to east of the FjaK llveden area studied by Ahlbom et al. (1991). Thus their results do not exclude more recent movements and irregularities in the isostatic rebound similar to those in VaK rmland (Risberg et al., 1996). East-westerly faults are also common in the Precambrian bedrock around Lake MaK laren (Lidmar-BergstroK m, 1994) but there is little evidence of Holocene
movements along these lines (see, however, MoK rner et al., 1989; TroK ften and MoK rner, 1997). The entire zone with faulting forms an elliptic area around the central Swedish lowland in which the big lakes MaK laren and HjaK lmaren are situated (Stephansson, 1979; Fig. 1). The central part of this area is seismically stable but the peripheral parts are a!ected by recent seismic activity.
4. Discussion De Geer (1940) proposed that the moraines in the Stockholm area were formed due to fracturing of the base of the ice as a consequence of seismic activity. Isostatic rebound generally takes place as a slow, elastic uplift of the Earth's crust but may at times have been released stepwise, causing fracturing of the bedrock. These movements resulted in earthquakes that a!ected the ice. Since the craton of Precambrian crystalline rocks was considered a tectonically stable region this idea was not generally accepted. Other interpretations of the moraines were favoured. However, the new evidence brie#y described above justi"es a reconsideration of this issue. Due to the preconditions necessary for formation of De Geer moraines, their absence can not be used as negative evidence in the following discussion about possible relationships between De Geer moraines and seismic activity. Therefore, only the actual occurrence of moraines will be considered. The northern area of De Geer moraines is well restricted to the seismically active area from the inland and
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along the coast of the Gulf of Bothnia (Fig. 1). Inland these moraines disappear, certainly due to the restrictions mentioned. Towards the south they are absent, although most probably topographic conditions would be favourable. There is no reason to believe that the ice thickness would prevent moraine formation. However, towards the south the indications of tectonic movements in connection with the deglaciation also disappear. The zone marked in Fig. 1 (after Stephansson, 1979) is based on the recent seismic activity. The southernmost moraines are found where the southernmost early Holocene scarps occur. On the other hand, similar moraines occur also on the eastern side of the Gulf. According to Talvitie (1979) this belongs to a zone of relatively high seismicity. In other parts of Swedish and Finnish Lapland De Geer moraines are absent due to a generally supraaquatic environment. Exceptions locally are found in areas where water was ponded up in front of the ice margin (BorgstroK m, 1979; Zilliacus, 1987). In Finland, the tectonic activity in connection with the deglaciation is well documented (Kujansuu, 1964; LagerbaK ck, 1990). Possibly this zone continues eastwards to the Rybachyi Peninsula (Fig. 4; Tanner, 1930). The only occurrences of De Geer moraines there are situated close to the zone with Holocene faults (Olesen, 1988) and also to a centre of recent seismicity (SahlstroK m, 1930). Most important among these occurrences are the Pasvik valley in northeasternmost Norway (Sollid and Carlson, 1984) and the small PeaK ldujaK vri "eld mentioned by Zilliacus (1987). The classical region of De Geer moraines in the Stockholm area and westwards in the MaK laren lowland is where Gerard De Geer came to the conclusion that seismic activity triggered the moraine formation. This is not a seismically active area today (Ba th, 1979) but conditions may have been di!erent at the time of deglaciation and earlier. Fault scarps are pronounced in the "eld (Sta lhoK s, 1969; Lidmar-BergstroK m, 1994), although in most cases there is no evidence of movements along them in Quaternary time. However, disturbances in the glacial sediments were interpreted as evidence of seismicity during the deglaciation (MoK rner, 1979, 1985, 1996; MoK rner et al., 1989; MoK rner and TroK ften, 1993; TroK ften, 1997). In these papers the occurrence of fresh-looking fault scarps was emphasised. Some scarps west to northwest of the city of Stockholm are so well preserved and una!ected by glacial erosion that it is reasonable to believe that they were formed after the ice left the location although there is no unequivocal evidence (see Henkel, 1988). All these occurrences are in the vicinity of the city where typical De Geer moraines are frequent. Notably, the southern boundary of the De Geer moraine clusters runs close to the ice margin 11230/10430, the time of a big earthquake according to MoK rner (1996). Together with the moraines are occurrences of huge boulders of local provenance mentioned above as well as remarkable concentrations of big angular blocks. Sometimes cave-
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like hollows occur beneath and in connection with such occurrences (cf. SjoK berg, 1994). The largest clusters of De Geer moraines are found in the western part of the MaK laren lowland. It may be signi"cant that the contours of the clusters are elliptic. This can sometimes be explained by topographic conditions, but in many cases this form seems to be independent of topography. It might indicate a seismic origin (cf. Smith, 1978; Adams, 1981) although this must be considered very speculative. So far no clear evidence of late tectonism has been shown there although MoK rner (1996) is of the opinion that the activity included this region. However, the distinct irregularity of the isobases over the central MaK laren lowland described above indicates that a bedrock block has been tilted towards the NW between faults or tectonic lines. There is no dating of this process, but notably the largest concentration of moraines is within the block and fairly well bounded by these lines. Although this area with De Geer moraines is nowadays seismically stable it is surrounded by one of the active zones in Sweden (Fig. 1). The idea that the entire elliptic area was more active in deglacial time with stronger isostatic rebound is speculative, but plausible. At least the coincidence between De Geer moraine distribution and seismicity is a fact that should not be neglected. The western area with De Geer moraines, at Lake VaK nern, is closely related to the area of seismic activity (Fig. 1). This is the most active area in Sweden today, where the strongest and most frequent earthquakes occur (Stephansson, 1979; Ba th, 1979). In the eastern margin of that area moraines are especially well developed (De Geer, 1896; Ericsson and LideH n, 1982). The boundary is the fault scarp at which Sandegren (1916) identi"ed Holocene irregularities in the land uplift. The western boundary coincides with the area where von Post (in Magnusson and von Post, 1929) and Risberg et al. (1996) found evidence for strong tectonism after the deglaciation. Most of the De Geer moraine area is situated within a graben between these tectonic zones, and partly occupied by the eastern part of Lake VaK nern. The isobases of the rebound show the same irregularity here as in the MaK laren region (e.g. Lundqvist, 1958, 1998; MoK rner, 1977). They give the same indication, namely, that a block between tectonic lines in NNW}SSE has been tilted northwards. Conditions in the other Nordic countries will not be discussed in detail here, but a brief review is appropriate. From Finland, Lapland and the Vaasa area were already considered above. The area northeast of Helsinki is tectonically stable (PenttilaK , 1964; Talvitie, 1979). However, it is possible that tectonic movements a!ected the area at the time of deglaciation. In the 18th and 19th centuries several epicenters were located in that area (MaK ntyniemi et al., 1993). Active zones traverse Finland northwest}southeast on both sides of the De Geer moraine
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area (PenttilaK , 1964; Talvitie, 1979). The isobases between these zones show a similar irregularity as in the Stockholm and Lake VaK nern areas (Fig. 4; KaK aK riaK inen, 1966), indicating a tilting of a block towards the northwest (cf. above). The third main group of De Geer moraines is situated on the western boundary of this block, also an area characterised by recent seismic activity (PenttilaK , 1964). The northwestern side of the block is limited by an active zone on the Bothnian coast (Tuominen et al., 1973). The southern side is situated close to the eastern boundary of activity in southern Finland de"ned by Talvitie (1979). Although delineated in a di!erent way the De Geer moraines appear to be situated inside an elliptic seismo-tectonic area. The occurrence of at least these three blocks, to which possibly the Oslo graben in southern Norway could be added, being tilted towards a centre somewhere near the Atlantic coast of central Scandinavia may be important for seismo-tectonic discussions. Similar irregularities on the Bothnian coast of Sweden (see MoK rner, 1977, Fig. 2) are oriented towards the same centre. This issue, however, is beyond the scope of the present study. The De Geer moraines in Norway are not very informative from the point of view discussed here. The Pasvik moraines were mentioned above. Those on the western coast (Larsen et al., 1991) are in a position where moraines can be formed in di!erent ways, but notably the Atlantic coast is one of the tectonically most active areas in Norway (Huseby et al., 1975, 1978; Stephansson, 1979; Anundsen, 1993). Thus this occurrence does not contradict the relationship proposed in this paper. There are small moraines also in the seismo-tectonically active Oslo graben. However, they have not been described as De Geer moraines and can naturally be explained solely by glacio-geological aspects.
5. Conclusions From the distribution of De Geer moraines in Fennoscandia in relation to the post-Weichselian seismic activity we can hardly draw any indisputable conclusions. There is, however, a striking correlation between De Geer moraine areas and areas characterised by postglacial faulting or recent seismicity. All occurrences are located in areas that are or have been seismically active, especially along their margins. Small moraines, referred to as cross-valley moraines, washboard moraines and similar terms, occur in other parts of the world and might not be relevant in this connection. We must observe that the terminology is somewhat unclear. The discussion here has only concerned those moraines that correspond very strictly to the moraines originally described by Gerard De Geer. Obviously, there are small moraine ridges to which this de"nition is not strictly applicable.
Even if we apply the term De Geer moraine very strictly, the formation of the ridges can be by more than one mechanism. The classic De Geer moraines are according to most studies formed in basal crevasses in glacier ice, probably when calving changed the environment to marginal. The only conclusion that can be drawn from the present study is that in many cases the origin of the basal crevasses may well be seismic. Seismic activity is one of several factors favouring formation of such crevasses. Thus it supports Gerard De Geer's old idea of these moraines as `seismic morainesa. Basal crevasses can, however, also originate for other reasons, for instance, break-up by calving or topographic factors. The idea of surges as a contributing factor, proposed by Zilliacus (1987), is noteworthy and may well be combined with a theory of seismic origin. These conclusions, although somewhat speculative, have even more important implications for our knowledge about the seismic stability of cratons of Precambrian crystalline rocks. They also draw the attention to the necessity to scrutinize other areas with small moraines, for instance in North America, for this aspect of their genesis.
Acknowledgements Some expenses for this project have been covered by a grant from the Swedish Natural Science Research Council. Additional information has been provided by Curt FredeH n and Lars Rudmark at the Geological Survey of Sweden. For valuable information and discussion I am grateful to Michael Stephens and Lars Brunnberg. Dave Liverman and an anonymous referee helped to improve the text. Drawings were made by Laszlo Madarasz, and for help with their completion by means of a computer programme I am indebted to Anders Borgmark and Hanna Dittrich.
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J. Lundqvist / Quaternary International 68}71 (2000) 175}186 Bergman, T., Isaksson, H., Johansson, R. (Ed.), LindeH n, A.H., Persson, Ch., Stephens, M., 1996. FoK rstudie NykoK ping. Jordarter, bergarter och deformationszoner. SKB DjupfoK rvar, Projekt Rapport PR D96-013, 82pp. #appendices. BjoK rck, S., Kromer, B., Johnsen, S., Bennike, O., Hammarlund, D., Lemdahl, G., Possnert, G., Lander Rasmussen, T., Wohlfarth, B., Hammer, C.U., Spurk, M., 1996. Synchronized terrestrial}atmospheric deglacial records around the North Atlantic. Science 274, 1155}1160. BorgstroK m, I., 1979. De Geer moraines in a Swedish mountain area? Geogra"ska Annaler 61A, 35}42. Brunnberg, L., 1995. Clay-varve chronology and deglaciation during the Younger Dryas and Preboreal in the easternmost part of the Middle Swedish Ice Marginal Zone. Quaternaria, Stockholm University, A : 2, 94pp. De Geer, G., 1896. AG ndmoraK ner vid Venerns ostligaste bugt. Geologiska FoK reningens i Stockholm FoK rhandlingar 18, 110}111. De Geer, G., 1940. Geochronologia Suecica Principles. Kungl. Svenska Vetenskapsakademiens Handlingar III : 18 : 6, 367pp. Ekman, M., 1996. A consistent map of the postglacial uplift of Fennoscandia. Terra Nova 8, 158}165. Ericsson, B., LideH n, E., 1982. Beskrivning till jordartskartan Karlskoga SV. Sveriges Geologiska UndersoK kning Ae 50, 73pp. Fromm, E., 1965. Beskrivning till jordartskarta oK ver Norrbottens laK n nedanfoK r lappmarksgraK nsen. Sveriges Geologiska UndersoK kning Ca 39, 236pp. Gardemeister, O., 1999. Strandniva er omkring 8000 BP i Tyresta Nationalpark, sydoK stra Sverige. Quaternaria, Stockholm University B16, 40pp. HaK ttestrand, C., 1997. Ribbed moraines in Sweden * distribution pattern and palaeoglaciological implications. Sedimentary Geology 111, 42}56. Henkel, H., 1988. Locality 6: The Vaxmora Fault scarp. In: MoK rner, N.-A. (Ed.), Proceedings of the Lejondal Symposium on NEOTECTONICS, ILP Report 1988, pp. 97}98. Hoppe, G., 1948. Isrecessionen fra n Norrbottens kustland i belysning av de glaciala formelementen. Geographica 20, 112pp. Hoppe, G., 1957. Problems of glacial morphology and the Ice Age. Geogra"ska Annaler 39, 1}18. Hoppe, G., 1959. Glacial morphology and inland ice recession in northern Sweden. Geogra"ska Annaler 41, 193}212. Huseby, E.S., Gj+ystdal, H., Bungum, H., Eldholm, O., 1975. The seismicity of the Norwegian and Greenland Seas and adjacent continental shelf areas. Tectonophysics 26, 55}70. Huseby, E.S., Bungum, H., Fyen, J., Gj+ystdal, H., 1978. Earthquake activity in Fennoscandia between 1497 and 1975 and intraplate tectonics. Norsk Geologisk Tidsskrift 58, 51}68. KaK aK riaK inen, E., 1966. Land uplift in Finland computed with the aid of precise levellings. Annales Academiv Scientarium Fennicv A III 90, 187}190. Kitagawa, H., van der Plicht, J., 1998. Atmospheric radiocarbon calibration to 45,000 yr BP: late glacial #uctuations and cosmogenic isotope production. Science 279, 1187}1190. Kleman, J., HaK ttestrand, C., BorgstroK m, I., Stroeven, A., 1997. Fennoscandian palaeoglaciology reconstructed using a glacial geological inversion model. Journal of Glaciology 43, 283}299. Kujansuu, R., 1964. Nuorista siirroksista Lapissa [Summary: Recent faults in Finnish Lapland]. Geologi 16, 30}36. LagerbaK ck, R., 1979. Neotectonic structures in northern Sweden. Geologiska FoK reningens i Stockholm FoK rhandlingar 100, 263}269. LagerbaK ck, R., 1988. Postglacial faulting and paleoseismicity in the LansjaK rv area, northern Sweden. SKB Technical Report 88}25, 37pp. LagerbaK ck, R., 1990. Late Quaternary faulting and paleoseismicity in northern Fennoscandia, with particular reference to the LansjaK rv area, northern Sweden. Geologiska FoK reningens i Stockholm FoK rhandlingar 112, 333}354.
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Larsen, E., Longva, O., Follestad, B.A., 1991. Formation of De Geer moraines and implications for deglaciation dynamics. Journal of Quaternary Science 6, 263}277. Lidmar-BergstroK m, K., 1994. Morphology of the bedrock surface. In: FredeH n, C. (Ed.), National Atlas of Sweden, Geology, pp. 44}54. Lundqvist, J., 1958. Beskrivning till jordartskarta oK ver VaK rmlands laK n. Sveriges Geologiska UndersoK kning Ca 38, 229pp. Lundqvist, J., 1969. Beskrivning till jordartskarta oK ver JaK mtlands laK n. Sveriges Geologiska UndersoK kning Ca 45, 418pp. Lundqvist, J., 1981. Moraine morphology * Terminological remarks and regional aspects. Geogra"ska Annaler 63A, 127}138. Lundqvist, J., 1989. Rogen (ribbed) moraine * identi"cation and possible origin. Sedimentary Geology 62, 281}292. Lundqvist, J., 1994. Traces of neotectonic movements in Southwestern VaK rmland? 21:a Nordiska Geologiska VintermoK tet, Lulea 1994. Abstracts. Tekniska HoK gskolan i Lulea , 121. Lundqvist, J., 1998. Istider och mellanistider. In: FredeH n, C. (Ed.), Sveriges Nationalatlas, Berg och Jord. Andra utga van, pp. 120}133. Lundqvist, J., LagerbaK ck, R., 1976. The PaK rve Fault: a late-glacial fault in the Precambrian of Swedish Lapland. Geologiska FoK reningens i Stockholm FoK rhandlingar 98, 45}51. Magnusson, N.H., von Post, L., 1929. Beskrivning till kartbladet SaK %e. Sveriges Geologiska UndersoK kning Aa 167, 96pp. MaK ntyniemi, P., WahlstroK m, R., Lindholm, C., Kijko, A., 1993. Seismic hazards in Fennoscandia: a regionalized study. Tectonophysics 227, 205}213. MoK ller, H., 1962. Annuella och interannuella aK ndmoraK ner. Geologiska FoK reningens i Stockholm FoK rhandlingar 84, 134}143. MoK rner, N.-A., 1977. Past and present uplift in Sweden: glacial isostasy, tectonism and bedrock in#uence. Geologiska FoK reningens i Stockholm FoK rhandlingar 99, 48}54. MoK rner, N.-A., 1979. Earth movements in Sweden, 20,000 BP to 20,000 AP. Geologiska FoK reningens i Stockholm FoK rhandlingar 100, 279}286. MoK rner, N.-A., 1985. Paleoseismicity and geodynamics in Sweden. Tectonophysics 117, 139}153. MoK rner, N.-A., 1996. Liquefaction and varve deformation as evidence of paleoseismic events and tsunamis. The autumn 10,430 BP case in Sweden. Quaternary Science Reviews 15, 939}948. MoK rner, N.-A., Somi, E., Zuchiewicz, W., 1989. Neotectonics and Paleoseismicity within the Stockholm intracratonal region in Sweden. Tectonophysics 163, 289}303. MoK rner, N.-A., TroK ften, P.-E., 1993. Palaeoseismotectonics in glaciated cratonal Sweden. Zeitschrift fuK r Geomorphologie N.F.94 (Suppl.Bd) 107}117. NiemelaK , J., 1971. Die quartaK re Stratigraphie von Tonablagerungen und der RuK ckzug des Inlandeises zwischen Helsinki und HaK meenlinna in SuK d"nnland. Geological Survey of Finland, Bulletin 253, 79pp. Olesen, O., 1988. The Stuoragurra Fault, evidence of neotectonics in the Precambrian of Finnmark, northern Norway. Norsk Geologisk Tidsskrift 68, 107}118. PenttilaK , E., 1964. Some remarks on earthquakes in Finland. Fennia 89, 25}28. Perhans, K.-E., 1981. Lervarvskronologin mellan Borensberg och Vinga ker. IGCP project 73/1/24, Symposium 12}13 januari 1981, Den senaste nedisningens foK rlopp med saK rskild haK nsyn till deglaciationen i Sverige. Sammanfattningar av foK redragen, pp. 66}68. Risberg, J., Sandgren, P., AndreH n, E., 1996. Early Holocene shore displacement and evidence of irregular isostatic uplift northwest of Lake VaK nern, western Sweden. Journal of Paleolimnology 15, 47}63. SahlstroK m, K.E., 1930. A seismological map of northern Europe. Sveriges Geologiska UndersoK kning C 364, 8pp. Sandegren, R., 1916. En postglacial strandlinje vid oK stra sidan av VaK nern. Sveriges Geologiska UndersoK kning C 270, 18pp.
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SjoK berg, R., 1994. Bedrock caves and fractured rock surfaces in Sweden. Occurrence and origin. Doctoral Thesis, Paleogeophysics and Geodynamics, Stockholm University, pp. 1}86. Smith, W.D., 1978. Spatial distribution of felt intensities for New Zealand earthquakes. New Zealand Journal of Geology and Geophysics 21, 293}311. Sollid, J.L., Carlson, A.B., 1984. De Geer Moraines and Eskers in Pasvik, North Norway. Striae 20, 55}61. Sollid, J.L., S+rbel, L., 1982. Kort beskrivelse til glasialgeologisk kart over Midt-Norge 1 : 500 000. Norsk geogra"sk Tidsskrift 36, 225}232. Sollid, J.L., S+rbel, L., 1984. Distribution and Genesis of Moraines in Central Norway. Striae 20, 63}67. Sollid, J. L., Torp, B., 1984. Glasialgeologisk kart. Nasjonalatlas for Norge. Map 2.3.2. Sta lhoK s, G., 1969. Beskrivning till Stockholmstraktens berggrund. Sveriges Geologiska UndersoK kning Ba 24, 190pp. Stephansson, O., 1979. Seismo-tectonics in Fennoscandia. Geologiska FoK reningens i Stockholm FoK rhandlingar 100, 239}245. StroK mberg, B., 1965. Mappings and geochronological investigations in some moraine areas of south-central Sweden. Geogra"ska Annaler 47A, 73}82. StroK mberg, B., 1989. Late Weichselian deglaciation and clay varve chronology in east-central Sweden. Sveriges Geologiska UndersoK kning Ca 73, 70pp. StroK mberg, B., 1990. A connection between the clay varve chronologies in Sweden and Finland. Annales Academiv Scientiarum Fennicv, Series A, III, Geologica-Geographica 154, 32pp.
Talvitie, J., 1979. Seismo-tectonics in Finland. Geologiska FoK reningens i Stockholm FoK rhandlingar 100, 247}253. Tanner, V., 1930. Studier oK ver kvartaK rsystemet i Fennoskandias nordliga delar. IV. Om niva foK raK ndringarna och grunddragen av den geogra"ska utvecklingen efter istiden i Ishavs"nland samt om homotaxin av Fennoskandias kvartaK ra marina avlagringar. Bulletin de la Commission GeH ologique de Finlande 88, 589pp. TroK ften, P.E., 1997. Neotectonics and Paleoseismicity in Southern Sweden with emphasis on sedimentological criteria. Doctoral Thesis, Paleogeophysics and Geodynamics, Stockholm University, 124pp. TroK ften, P.-E., MoK rner, N.-A., 1997. Varved clay chronology as a means of recording paleoseismic events in southern Sweden. Journal of Geodynamics 24, 249}258. Tuominen, H.V., Aarnisalo, J., SoK derholm, B., 1973. Tectonic patterns in the central Baltic Shield. Bulletin of the Geological Society of Finland 45, 205}217. Ussisoo, I., 1977. Computation of land uplift and mean sea level in Sweden. Geologiska FoK reningens i Stockholm FoK rhandlingar 99, 42}48. Wohlfarth, B., BjoK rck, S., Cato, I., Possnert, G., 1997. A new middle Holocene varve diagram from the river As ngermanaK lven, northern Sweden: indications for a possible error in the Holocene varve chronology. Boreas 26, 347}353. Zilliacus, H., 1987. De Geer moraines in Finland and the annual moraine problem. Fennia 165, 145}239.
Palaeoseismicity and De Geer Moraines Jan Lundqvist* Department of Quaternary Research, Stockholm University, S-106 91 Stockholm, Sweden
Abstract De Geer moraines are small moraine ridges formed partly in crevasses at the base of glacier ice. There may be several di!erent origins of the crevasses. One possibility, proposed by Gerard De Geer who "rst studied these ridges, was seismic activity in connection with the isostatic rebound during the deglaciation. The present study shows a close connection between occurrences of De Geer moraines in Fennoscandia and seismic activity or fresh-looking fault lines and other features indicating neotectonic movements. This supports De Geer's theory but does not exclude formation in crevasses formed in other ways. There is no real proof that the tectonic movements took place in the Quaternary but the close connection between the moraines and tectonic features is evidence pointing in this direction. If this is correct it has important implications for the stability within cratons of Precambrian crystalline bedrock. 2000 Elsevier Science Ltd and INQUA. All rights reserved.
1. Introduction De Geer moraine is a term used for a type of small moraine ridge discussed by Gerard De Geer (e.g. De Geer, 1940). These ridges are usually a few 100 m in length, not more than some 10}20 m broad and a few metres high. Characteristically, they occur in large swarms or clusters within which the ridges show a general trend parallel to the ice margin. The orientation of individual moraines varies within wide limits. Ridges may even be perpendicular to the margin or winding and anastomosing. The following discussion relates only to small moraines that correspond to this de"nition * which actually applies to virtually all areas with small moraines in Fennoscandia. Fig. 1 shows all such areas within Sweden. Isolated small moraine ridges, probably formed in di!erent ways, are not included in the map, and this applies also to moraines in the mountain area and within the Younger Dryas and other end-moraine belts. In the latter case, they obviously belong to the endmoraine systems, although smaller than the typical end moraines. There has been many di!erent ideas as to the origin of these ridges. De Geer (1940) considered them end moraines, and deposited mainly annually at the ice margin. Individual ridges could also be formed between the an-
* Fax: #46-8-6747895. E-mail address: [email protected] (J. Lundqvist).
nual ridges due to calving, as later con"rmed by MoK ller (1962). Consequently, they have often in the literature been called annual moraines or just end moraines * even the term terminal moraines has been used. Hoppe (1957) and StroK mberg (1965), contrary to De Geer, interpreted them as formed some distance inside the ice margin. Di!erent ideas have also been put forward with respect to the mode of deposition. The internal structures show compact basal till as well as ablation till and waterlain sediments. Folded inclusions of clay and other proglacial sediments occur. These di!erent types of sediment indicate deposition by both push, dump, squeeze, and lodgement. A possible compromise suggests that the moraines started to form in basal crevasses. These crevasses initiated later break-up by calving. More material could then be added to the ridges when they were in a marginal position. The di!erent possible processes of formation have been discussed in many papers (see, e.g., Hoppe, 1948, 1957; StroK mberg, 1965; Zilliacus, 1987) and will not be treated here. In order to avoid these problems the non-genetic term De Geer moraines was proposed by Hoppe (1959). Reviews of the discussions about formation and terminology were presented by Lundqvist (1981) and Zilliacus (1987). In Fennoscandia, the classic De Geer moraines occur in #at lowland regions in the central and northern parts of Sweden and Finland and also in some restricted areas in Norway. Generally, these moraines were formed beneath water, in glacial lakes or, usually, the sea. Moraines
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authors (see references in Lundqvist, 1981). There is some doubt that all these types of small ridges are the result of the same process of formation, and a more systematic classi"cation is needed, as emphasised by Lundqvist (1981). In the following discussion only the occurrences in Fennoscandia which are in agreement with De Geer's original description will be treated. The processes of formation, the relation to the ice margin and a number of other aspects brie#y mentioned above will not be discussed. The aim of this paper is only to consider the possible relation between De Geer moraines and seismicity in Fennoscandia. According to De Geer (1940) fracturing of the basal part of the ice by seismic activity contributed to the formation of the ridges. The reason for this activity should be the isostatic rebound after the waning of the last (Weichselian) Scandinavian inland ice. This idea has not gained general acceptance, although supported by MoK rner in a number of papers (e.g. MoK rner, 1977, 1979; MoK rner et al., 1989). In the light of subsequent discoveries about neotectonic activity in connection with the deglaciation (Lundqvist and LagerbaK ck, 1976; MoK rner, 1977, 1985; LagerbaK ck, 1979, 1990) this aspect deserves further consideration.
2. Regional distribution of De Geer moraines 2.1. Sweden
Fig. 1. Seismically active areas in Sweden (dashed), De Geer moraine areas (black) and ice-marginal lines with dates (clay-varve years/calibrated years). P marks the PaK rvie fault and L the main LansjaK rv fault, referred to in the text.
cannot be formed if the ice is #oating, nor if it is too thick, which makes the pressure against the bottom comparable with supra-aquatic conditions. In the "eld this is noticed as the limitation of moraines to certain levels. Similar ridges have been described from other areas, notably in the region of the North American ice sheets. Terms like cross-valley moraines, washboard moraines, small moraines and others have been used by di!erent
The outlines of the regional distribution of De Geer moraines within Sweden are shown in Fig. 1. This picture is compiled from available geological maps but details can be found in many special papers referred to in the following. Very generally, De Geer moraines belong to areas that were situated below the level of the sea or large lakes during deglaciation, that is, subaquatic areas. They are also restricted to certain levels in the terrain * on higher as well as lower ground they are absent. The altitude varies from area to area, but mostly the level implies formation at water depths of 100}250 m, in each area depending on the local thickness of the ice. Fig. 1 shows all clusters of De Geer moraines, with two possible exceptions, marked with X. These occurrences, described by Lundqvist (1969) and BorgstroK m (1979) are situated near each other in mountain valleys and are too small to be shown in the scale. They may be exceptions from the general rule of distribution, but it is uncertain whether they should be called De Geer moraines * although this term was used by BorgstroK m (1979). As a matter of fact, they are also situated close to a seismic area in central Norway (Stephansson, 1979). The Ma( laren area: The main area of distribution is situated around Lake MaK laren in south-central Sweden. Its eastern part is the one described and mapped in detail by De Geer in his main paper (De Geer, 1940) (Fig. 2).
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177
Fig. 2. The main area with De Geer moraines in east-central Sweden. Black areas"De Geer moraine clusters; grey areas"lakes and (in the east) the Baltic Sea; dentated lines"ice margin with dates in clay-varve years/calibrated years (the line 11,230/10,430 runs 2 km north of the line 10,450/11,250); thin broken lines"isobases for isostatic uplift ("gures are an arbitrary scale * in order to obtain the altitude in metres above the sea-level at the time of deglaciation approximately 25 m should be added); thick lines"faults with indications of Holocene movements; thick broken lines"fault lines without clear indication of Holocene movements; dotted lines"hypothetical tectonic boundaries of area with di!erential isostatic rebound. BS"Baltic Sea; H"Lake HjaK lmaren; M"Lake MaK laren; S"Stockholm; U"Uppsala; V"VaK stera s; OG "OG rebro.
The moraines occur almost strictly between the ice-marginal lines dated at 10,450 and 10,200 clay-varve years BP (StroK mberg, 1989; Brunnberg, 1995). Considering possible errors in the clay-varve time scale (Wohlfarth et al., 1997) and calibration of radiocarbon dates according to Kitagawa and van der Plicht (1998) this can be estimated to roughly 11,250}11,000 calendar years BP. The zone ends abruptly towards a highland in the west (Fig. 2). East of the area shown in Fig. 2, according to personal communication by Christer Persson and Lars Rudmark at the Geological Survey of Sweden, there are some small occurrences (see also StroK mberg, 1965). This continuation is not included in the map since the area had not been mapped in su$cient detail when Fig. 2 was drawn. However, these occurrences are restricted in time like in the map area. Notably, a few small moraine groups occur north of the well-de"ned area, and corresponding in time to the northernmost clusters in the VaK nern area (see below). The water depth in the area is not known exactly since no part of the area reached above the sea-level during deglaciation but is estimated to be 130}160 m. The "gures in Fig. 2 are an arbitrary scale (in meters) used by
As se and BergstroK m (1984). In order to obtain the level of the sea when the moraines were formed we should add about 25 m to these "gures. Typical for this area as well as the other areas with De Geer moraines is the clustering of ridges. De Geer (1940) called such a cluster a `moraa, using a word from the Alps meaning a concentration of big boulders. One of the clusters is situated at Bromma in the western suburbs of Stockholm. Using this as an example De Geer developed his idea of a seismic origin of such moraines. The clustering, often with elliptic contours, is well visible in Fig. 2. Moraines do occur also outside the clusters but ridges there are few and scattered. Within at least some clusters there are also concentrations of very large boulders of local provenance. The Va( nern area: As mentioned above, the MaK laren area is interrupted westwards by a highland. West of this highland there is a continuation of moraines formed at about the same time as those to the east. The moraines are clustered in a similar fashion. One of these clusters is the well-known Kilsviken (or As ra s) Bay (Fig. 3) on the eastern side of Lake VaK nern, "rst described by De Geer (1896). A detailed map was published by Ericsson and
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Fig. 3. De Geer moraines in the Kilsviken Bay, Lake VaK nern. Photo J. Lundqvist, 1996.
LideH n (1982). This is considered the best developed area of De Geer moraines in Scandinavia. North of the lake old geological maps show several concentrations of De Geer moraines. Many of them, however, have been eliminated in later mapping (Lundqvist, 1958), but nevertheless several groups of moraines do occur. There is no good dating of this moraine area but according to Perhans (1981) the dated lines from the east continue as shown in Fig. 1. This means that the moraines here were formed between 11,500 and 10,900 calendar years BP. Using the levels of the highest post-glacial coastline of Lundqvist (1958) the water depth during the formation is estimated to be about 75}130 m. Gulf of Bothnia: The third region with De Geer moraines in Sweden extends along the coast of the northern part of the Gulf of Bothnia * called the Bothnian Bay. These moraines contain large volumes of sorted, coarseclastic sediments. They occur partly together with larger ridges, called Niemisel ridges by Lundqvist (1981). As shown by HaK ttestrand (1997) these ridges should actually be classi"ed as Rogen moraine, a type that was obviously subglacially formed (Lundqvist, 1989; HaK ttestrand, 1997). This co-existence supports the idea of De Geer moraines as subglacially deposited, as was also proposed by Hoppe (1948). The pattern of ridges is in this case more regular than in most other areas. Hoppe's (1948) interpretation was that the moraines were formed in basal crevasses when the thinning ice in the Bothnian Bay became a#oat. The group of De Geer moraines at Bothnian Bay has a continuation in the Vaasa area on the Finnish side of the bay. The Finnish occurrence may suggest that more
ridges are located on the bottom of Bothnian Bay, but AndreH n's (1990) seismic studies of the sediments there give no such indication. There is no accurate dating of this group of De Geer moraines, but AndreH n (1990) dated sediments in the Bay by means of clay-varve chronology. Calibration as described above gives an age between 10,000 and 10,300 BP for the moraines. The level of the highest coastline (Fromm, 1965) indicates that the water depth at the time of formation was 200}250 m. 2.2. Finland and Norway De Geer moraines in Finland occur in the southwestern part of the country where they are concentrated in a number of clusters (Fig. 4; Zilliacus, 1987, Fig. 37). The main clusters are situated in two areas on the coast of the Gulf of Bothnia. The best developed group is the one mentioned above, in the Vaasa region. Also farther south, around and NE of Uusikaupunki (Nystad) there are similar groups. A third main group is situated outside the First SalpausselkaK between Helsinki and Lahti. Together with these main groups there are smaller clusters. Altogether, the occurrences form a triangular area between Vaasa, the As land archipelago and Hamina (Fredrikshamn) in the southeast. Beyond this triangle there are just a few small occurrences in the extreme east and north of the Bothnian Bay. The clusters are similar to the ones described above from Sweden. Notably, in the Vaasa region, as on the Swedish side of Bothnian Bay De Geer moraines occur
J. Lundqvist / Quaternary International 68}71 (2000) 175}186
179
ridges, but since they occur outside the First SalpausselkaK , dated to Younger Dryas, they must antedate 12,600 calibrated yr BP (cf. BjoK rck et al., 1996). In accordance with the clay-varve chronology of NiemelaK (1971) they were formed from about 13,000 BP. The age of the youngest moraines can be estimated at 10,400 calibrated yr BP (cf. StroK mberg, 1990). In Finland the topographic conditions are similar to Sweden in the areas where De Geer moraines occur, that is, a rather #at, slightly undulating landscape below the coastline at the time of deglaciation. In Norway, the topography is quite di!erent, characterised by mountainous terrain with deep valleys and fjords. However, since a few moraine occurrences have been called De Geer moraines (Sollid and Torp, 1984) they should also be brie#y considered here. On the Atlantic coast in the As lesund area, western Norway, moraine ridges have been described as De Geer moraines (Sollid and S+rbel, 1984; Larsen et al., 1991). They occur mainly on the #oor of fjords outside the Younger Dryas moraines (see Sollid and S+rbel, 1982 with map) and must thus be at least as old as the moraines in southern Finland. The water depth at the time of formation can be estimated to just some tens of metres, certainly less than 100. Larsen et al. (1991) interpreted the ridges as true frontal formations. Although called De Geer moraines, it is questionable if they are compatible with De Geer moraines to be discussed in this paper. Similar uncertainty concerns an occurrence of moraines in the Pasvik valley, northernmost Norway (Fig. 4), studied in detail by Sollid and Carlson (1984). These moraines occur in valley position and were formed a few tens of metres below the contemporaneous sealevel. On the basis of the sedimentology in the ridges they were also in this case interpreted as frontal formations. All these Norwegian moraines could be described as cross-valley moraines, although according to the current less precise terminology (cf. Lundqvist, 1981) it is correct to include them among De Geer moraines.
3. Evidence of neotectonic activity
Fig. 4. De Geer moraine areas (black) in Finland according to Zilliacus (1987), isobases for the present isostatic uplift (broken lines, "gures in cm/yr) according to KaK aK riaK inen (1966), and fault scarps (indented lines) according to LagerbaK ck (1990). The dotted line shows the Younger Dryas ice margin.
together with larger ridges of Rogen moraine. The water depth during their formation varies from less than 100 m to more than 250 m. The time of formation is over a longer period than on the Swedish side. There is no exact dating of the oldest
The isostatic rebound after the last glaciation caused a slow and ongoing tectonic activity. The Precambrian basement in Scandinavia has been considered otherwise stable. In the last few years, however, a number of observations, partly in connection with investigations for repositories of nuclear waste, indicate that it is more subject to tectonic movements than earlier supposed. The most convincing evidence was presented in northernmost Sweden by Lundqvist and LagerbaK ck (1976), followed by detailed studies by LagerbaK ck (e.g. 1979, 1990). In Finnish Lapland similar evidence had been published by Kujansuu (1964). Although not generally accepted, indications of tectonic activity were also presented by
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MoK rner in a number of papers (e.g., 1977, 1979, 1985, 1996; MoK rner et al., 1989; also MoK rner and TroK ften, 1993; TroK ften, 1997) in the Stockholm area. In Swedish and Finnish Lapland and adjacent parts of Norway there are up to 30 m high fault scarps indicating seismic activity (Fig. 5). The area with faults can be followed from near the Caledonian mountain front in the northwest southwards to some 50 km west of the Bothnian coast at Skelleftea (Figs. 1 and 4). Some of the scarps can be very precisely dated to the time of the local deglaciation, although there is so far no possibility to obtain an age in absolute years. This applies in particular to the PaK rvie fault in northernmost Lapland (P in Figs. 1 and 4). Along this line the fault scarp is well developed. In parts it cuts beaches of ice-dammed lakes and meltwater channels formed at the ice margin. In adjacent parts such features intersect the scarp una!ected. Liqui"ed sediments and faults in the till from the last deglaciation as well as landslides are clear evidence that faulting at least partly took place after the local deglaciation. This is especially clear in the LansjaK rv area (L in Figs. 1 and 4, LagerbaK ck, 1979, 1990). The features of liquefaction and landslides even in coarse-textured till as well as the fault scarps themselves allow an estimation of the magnitude of the earthquakes related to the faulting. After a discussion of these features, LagerbaK ck (1988) concluded that faulting was accompanied by earthquakes with a magnitude of M 7.0}7.5 or probably higher. A tentative extension of known lines of ice recession (Fig. 1) based on our present knowledge of the deglaciation (e.g. Kleman et al., 1997) allows an approximate dating of these tectonic events to
10,000}10,300 calibrated yr BP. In Finnish Lapland faults may be older than this. All the North-Swedish postglacial faults, a few of which are shown in Figs. 1 and 4, are located within or close to zones of recent seismicity (Ba th, 1979; Stephansson, 1979). These zones extend from northwestern Lapland to the northern coast of the Bothnian Bay and further southward along the coast (Fig. 1). The region in Sweden where the seismic activity is strongest at present is the area around Lake VaK nern in the southwest (Fig. 1; Ba th, 1979), with modern earthquakes detected. The lake basin is bounded by fault lines in the west and east. Sandegren (1916) observed shorelines on the eastern side which he interpreted as a!ected by faulting in early Holocene time. On the northwestern side of the basin the bedrock is fractured and traversed by numerous tectonic valleys. There mostly is no evidence that movements took place after the deglaciation, but from dates of the isolation of small lake basins Risberg et al. (1996) concluded that bedrock blocks had moved di!erently after the isolation. Lundqvist (1994) pointed out that strong fracturing of the bedrock in certain zones could be an e!ect of these movements. We can conclude that the area around and north of Lake VaK nern has been subject to tectonic movements in postglacial time. The VaK nern zone extends southwards to southernmost Scandinavia and is also closely related to the Oslo graben in Norway (Huseby et al., 1978). The Stockholm area and the region around and west of Lake MaK laren is the region where De Geer (1940) proposed his theory of seismic moraines, in particular, the Bromma area just west of downtown Stockholm. The evidence upon which the theory was based was one small
Fig. 5. The PaK rvie fault in northern Swedish Lapland. Photo J. Lundqvist, 1975.
J. Lundqvist / Quaternary International 68}71 (2000) 175}186
fault scarp interpreted as recently active, and a number of very large boulders (Fig. 6) and concentrations of boulders (Fig. 7) which were obviously displaced by the gla-
Fig. 6. Large boulder transported a very short distance by the inland ice. Bromma, western Stockholm. Photo J. Lundqvist, 1981.
181
cier, albeit a very short distance * just a few metres. The boulders are striated even on their basal side and so fractured and fragile that they could not possibly be transported, perhaps just tilted. New evidence has been put forward by MoK rner (1977, 1979, 1985) and MoK rner et al. (1989). Liqui"ed sediments from the time of deglaciation have been interpreted as a result of seismic activity (MoK rner, 1996; TroK ften, 1997; TroK ften and MoK rner, 1997). Seismic events were dated around 10,400 clay-varve yr BP with a strong event at 10,430, which corresponds to 11,230 calibrated yr BP (MoK rner, 1996; cf. Fig. 2). Liquefaction of clay, especially if it is rich in silt as the studied sediments, can take place due to several reasons, however, and Brunnberg (1995) did not "nd traces of such an event in his detailed clay-varve chronology. Possibly the sediments were disturbed only within narrow zones. New observations of faults give additional support to the hypothesis of seismic instability. Several fault scarps in the Stockholm area look very fresh and una!ected by glacial erosion (Fig. 8). Di!erential uplift in postglacial time has been demonstrated by Gardemeister (1999). Isobases for the isostatic rebound, ongoing as well as total from the time of deglaciation, show a marked irregularity from Stockholm some 130 km westwards (Fig. 2; MoK rner, 1977; Ussisoo, 1977; As se and BergstroK m, 1984; MoK rner, 1977, Fig. 2). New investigations by Ekman (1996) indicate that this feature is within the margin of error but if it is a reality it means that a block has moved di!erentially between fractures or #exures trending about NW}SE. Like the basin of Lake VaK nern discussed above it appears that a block of the bedrock has tilted northwards round an axis at an oblique angle to the ice margin and isobases. As such, perhaps the block
Fig. 7. Concentration of huge angular boulders on the fault line 10 km north of Stockholm centre (cf. Fig. 2). Photo J. Lundqvist, 1971.
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Fig. 8. 12 m high fault scarp, 10 km north of Stockholm (see Henkel, 1988). Photo J. Lundqvist, 1999.
was left behind during the isostatically induced general tilting southwards of the land surface. The eastern boundary of this block coincides with the area with fresh-looking faults and the best-developed concentrations of fractured rocks. In Fig. 2 the continuation northwestwards of this zone is marked with a dotted line. The western boundary has been studied in connection with the present work. The area in general is rather rich in boulders, slightly varying with the types of bedrock. However, west to south of VaK stera s (Fig. 2) there are very distinct valleys of tectonic origin. Most of them are certainly pre-Quaternary (Ahlbom et al., 1991; Bergman et al., 1996) The outcrops on the valley sides are scoured by ice. Several parallel lines are clearly visible on topographic and geologic maps. Notably, smaller fracture lines parallel to the topographically most distinct valleys show fresher, although smaller, scarps and a concentration of angular boulders that is markedly higher than at the main lines. This is most clearly observed south of the western end of Lake MaK laren (Fig. 2). There is no conclusive evidence that these faults and fractures were active in late glacial or Holocene time, but this is suggested by the similarity with the Bromma area. The zone with such lines runs about 10 km north to east of the FjaK llveden area studied by Ahlbom et al. (1991). Thus their results do not exclude more recent movements and irregularities in the isostatic rebound similar to those in VaK rmland (Risberg et al., 1996). East-westerly faults are also common in the Precambrian bedrock around Lake MaK laren (Lidmar-BergstroK m, 1994) but there is little evidence of Holocene
movements along these lines (see, however, MoK rner et al., 1989; TroK ften and MoK rner, 1997). The entire zone with faulting forms an elliptic area around the central Swedish lowland in which the big lakes MaK laren and HjaK lmaren are situated (Stephansson, 1979; Fig. 1). The central part of this area is seismically stable but the peripheral parts are a!ected by recent seismic activity.
4. Discussion De Geer (1940) proposed that the moraines in the Stockholm area were formed due to fracturing of the base of the ice as a consequence of seismic activity. Isostatic rebound generally takes place as a slow, elastic uplift of the Earth's crust but may at times have been released stepwise, causing fracturing of the bedrock. These movements resulted in earthquakes that a!ected the ice. Since the craton of Precambrian crystalline rocks was considered a tectonically stable region this idea was not generally accepted. Other interpretations of the moraines were favoured. However, the new evidence brie#y described above justi"es a reconsideration of this issue. Due to the preconditions necessary for formation of De Geer moraines, their absence can not be used as negative evidence in the following discussion about possible relationships between De Geer moraines and seismic activity. Therefore, only the actual occurrence of moraines will be considered. The northern area of De Geer moraines is well restricted to the seismically active area from the inland and
J. Lundqvist / Quaternary International 68}71 (2000) 175}186
along the coast of the Gulf of Bothnia (Fig. 1). Inland these moraines disappear, certainly due to the restrictions mentioned. Towards the south they are absent, although most probably topographic conditions would be favourable. There is no reason to believe that the ice thickness would prevent moraine formation. However, towards the south the indications of tectonic movements in connection with the deglaciation also disappear. The zone marked in Fig. 1 (after Stephansson, 1979) is based on the recent seismic activity. The southernmost moraines are found where the southernmost early Holocene scarps occur. On the other hand, similar moraines occur also on the eastern side of the Gulf. According to Talvitie (1979) this belongs to a zone of relatively high seismicity. In other parts of Swedish and Finnish Lapland De Geer moraines are absent due to a generally supraaquatic environment. Exceptions locally are found in areas where water was ponded up in front of the ice margin (BorgstroK m, 1979; Zilliacus, 1987). In Finland, the tectonic activity in connection with the deglaciation is well documented (Kujansuu, 1964; LagerbaK ck, 1990). Possibly this zone continues eastwards to the Rybachyi Peninsula (Fig. 4; Tanner, 1930). The only occurrences of De Geer moraines there are situated close to the zone with Holocene faults (Olesen, 1988) and also to a centre of recent seismicity (SahlstroK m, 1930). Most important among these occurrences are the Pasvik valley in northeasternmost Norway (Sollid and Carlson, 1984) and the small PeaK ldujaK vri "eld mentioned by Zilliacus (1987). The classical region of De Geer moraines in the Stockholm area and westwards in the MaK laren lowland is where Gerard De Geer came to the conclusion that seismic activity triggered the moraine formation. This is not a seismically active area today (Ba th, 1979) but conditions may have been di!erent at the time of deglaciation and earlier. Fault scarps are pronounced in the "eld (Sta lhoK s, 1969; Lidmar-BergstroK m, 1994), although in most cases there is no evidence of movements along them in Quaternary time. However, disturbances in the glacial sediments were interpreted as evidence of seismicity during the deglaciation (MoK rner, 1979, 1985, 1996; MoK rner et al., 1989; MoK rner and TroK ften, 1993; TroK ften, 1997). In these papers the occurrence of fresh-looking fault scarps was emphasised. Some scarps west to northwest of the city of Stockholm are so well preserved and una!ected by glacial erosion that it is reasonable to believe that they were formed after the ice left the location although there is no unequivocal evidence (see Henkel, 1988). All these occurrences are in the vicinity of the city where typical De Geer moraines are frequent. Notably, the southern boundary of the De Geer moraine clusters runs close to the ice margin 11230/10430, the time of a big earthquake according to MoK rner (1996). Together with the moraines are occurrences of huge boulders of local provenance mentioned above as well as remarkable concentrations of big angular blocks. Sometimes cave-
183
like hollows occur beneath and in connection with such occurrences (cf. SjoK berg, 1994). The largest clusters of De Geer moraines are found in the western part of the MaK laren lowland. It may be signi"cant that the contours of the clusters are elliptic. This can sometimes be explained by topographic conditions, but in many cases this form seems to be independent of topography. It might indicate a seismic origin (cf. Smith, 1978; Adams, 1981) although this must be considered very speculative. So far no clear evidence of late tectonism has been shown there although MoK rner (1996) is of the opinion that the activity included this region. However, the distinct irregularity of the isobases over the central MaK laren lowland described above indicates that a bedrock block has been tilted towards the NW between faults or tectonic lines. There is no dating of this process, but notably the largest concentration of moraines is within the block and fairly well bounded by these lines. Although this area with De Geer moraines is nowadays seismically stable it is surrounded by one of the active zones in Sweden (Fig. 1). The idea that the entire elliptic area was more active in deglacial time with stronger isostatic rebound is speculative, but plausible. At least the coincidence between De Geer moraine distribution and seismicity is a fact that should not be neglected. The western area with De Geer moraines, at Lake VaK nern, is closely related to the area of seismic activity (Fig. 1). This is the most active area in Sweden today, where the strongest and most frequent earthquakes occur (Stephansson, 1979; Ba th, 1979). In the eastern margin of that area moraines are especially well developed (De Geer, 1896; Ericsson and LideH n, 1982). The boundary is the fault scarp at which Sandegren (1916) identi"ed Holocene irregularities in the land uplift. The western boundary coincides with the area where von Post (in Magnusson and von Post, 1929) and Risberg et al. (1996) found evidence for strong tectonism after the deglaciation. Most of the De Geer moraine area is situated within a graben between these tectonic zones, and partly occupied by the eastern part of Lake VaK nern. The isobases of the rebound show the same irregularity here as in the MaK laren region (e.g. Lundqvist, 1958, 1998; MoK rner, 1977). They give the same indication, namely, that a block between tectonic lines in NNW}SSE has been tilted northwards. Conditions in the other Nordic countries will not be discussed in detail here, but a brief review is appropriate. From Finland, Lapland and the Vaasa area were already considered above. The area northeast of Helsinki is tectonically stable (PenttilaK , 1964; Talvitie, 1979). However, it is possible that tectonic movements a!ected the area at the time of deglaciation. In the 18th and 19th centuries several epicenters were located in that area (MaK ntyniemi et al., 1993). Active zones traverse Finland northwest}southeast on both sides of the De Geer moraine
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area (PenttilaK , 1964; Talvitie, 1979). The isobases between these zones show a similar irregularity as in the Stockholm and Lake VaK nern areas (Fig. 4; KaK aK riaK inen, 1966), indicating a tilting of a block towards the northwest (cf. above). The third main group of De Geer moraines is situated on the western boundary of this block, also an area characterised by recent seismic activity (PenttilaK , 1964). The northwestern side of the block is limited by an active zone on the Bothnian coast (Tuominen et al., 1973). The southern side is situated close to the eastern boundary of activity in southern Finland de"ned by Talvitie (1979). Although delineated in a di!erent way the De Geer moraines appear to be situated inside an elliptic seismo-tectonic area. The occurrence of at least these three blocks, to which possibly the Oslo graben in southern Norway could be added, being tilted towards a centre somewhere near the Atlantic coast of central Scandinavia may be important for seismo-tectonic discussions. Similar irregularities on the Bothnian coast of Sweden (see MoK rner, 1977, Fig. 2) are oriented towards the same centre. This issue, however, is beyond the scope of the present study. The De Geer moraines in Norway are not very informative from the point of view discussed here. The Pasvik moraines were mentioned above. Those on the western coast (Larsen et al., 1991) are in a position where moraines can be formed in di!erent ways, but notably the Atlantic coast is one of the tectonically most active areas in Norway (Huseby et al., 1975, 1978; Stephansson, 1979; Anundsen, 1993). Thus this occurrence does not contradict the relationship proposed in this paper. There are small moraines also in the seismo-tectonically active Oslo graben. However, they have not been described as De Geer moraines and can naturally be explained solely by glacio-geological aspects.
5. Conclusions From the distribution of De Geer moraines in Fennoscandia in relation to the post-Weichselian seismic activity we can hardly draw any indisputable conclusions. There is, however, a striking correlation between De Geer moraine areas and areas characterised by postglacial faulting or recent seismicity. All occurrences are located in areas that are or have been seismically active, especially along their margins. Small moraines, referred to as cross-valley moraines, washboard moraines and similar terms, occur in other parts of the world and might not be relevant in this connection. We must observe that the terminology is somewhat unclear. The discussion here has only concerned those moraines that correspond very strictly to the moraines originally described by Gerard De Geer. Obviously, there are small moraine ridges to which this de"nition is not strictly applicable.
Even if we apply the term De Geer moraine very strictly, the formation of the ridges can be by more than one mechanism. The classic De Geer moraines are according to most studies formed in basal crevasses in glacier ice, probably when calving changed the environment to marginal. The only conclusion that can be drawn from the present study is that in many cases the origin of the basal crevasses may well be seismic. Seismic activity is one of several factors favouring formation of such crevasses. Thus it supports Gerard De Geer's old idea of these moraines as `seismic morainesa. Basal crevasses can, however, also originate for other reasons, for instance, break-up by calving or topographic factors. The idea of surges as a contributing factor, proposed by Zilliacus (1987), is noteworthy and may well be combined with a theory of seismic origin. These conclusions, although somewhat speculative, have even more important implications for our knowledge about the seismic stability of cratons of Precambrian crystalline rocks. They also draw the attention to the necessity to scrutinize other areas with small moraines, for instance in North America, for this aspect of their genesis.
Acknowledgements Some expenses for this project have been covered by a grant from the Swedish Natural Science Research Council. Additional information has been provided by Curt FredeH n and Lars Rudmark at the Geological Survey of Sweden. For valuable information and discussion I am grateful to Michael Stephens and Lars Brunnberg. Dave Liverman and an anonymous referee helped to improve the text. Drawings were made by Laszlo Madarasz, and for help with their completion by means of a computer programme I am indebted to Anders Borgmark and Hanna Dittrich.
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