The quaternary uplift of the rhenish shield in germany

Z’ectonophysics, 61 (1979) 197-225 @ Elsevier Scientific Publishing Company, Amsterdam - Printed in The Netherlands

THE QUATERNARY

J. HENNING ILLIES, ARNO SEMMEL

UPLIFT OF THE RHENISH

CLAUS PRODEHL,

HANS-ULRICH

197

SHIELD IN GERMANY

SCHMINCKE

and

Geologisches Znstitut, Universitiit Karlsruhe, D 75 Karlsruhe (F.R. of Germany) Geophysikalisches Znstitut, Universitiit Karlsruhe, D 75 Karlsruhe (F. R. of Germany) Znstitut fiir Mineralogie, Ruhr-Universitiit Bochum, D 4630 Bochum (F.R. of Germany) Geographisches Znstitut, Johann Wolfgang Goethe-Universitiit, D 6000 Frankfurt/Main (F.R. of Germany) (Received May 3, 1979)

ABSTRACT Illies, J.H., Prodehl, C., Schminke, H.-U. and Semmel, A., 1979. The Quaternary uplift of the Rhenish shield in Germany. In: T.R. McGetchin and R.B. Merrill (Editors), Plateau Uplift: Mode and Mechanism. Tectonophysics, 61: 197-225. The Rhenish shield in Central Europe is made up by a series of slates, quartzites and limestones of dominantly Devonian age which were strongly folded by the Hercynian orogeny. The area has been uplifted to about 300 m since the Pliocene. Rates of uplift increased during the last 600,000 years as shown in the uplifted terraces of the Rhine river. The drainage pattern was forced to erode deep antecedent river valleys in the uplifted plateau. Two late Quaternary volcanic fields have developed on the Rhenish shield: The West Eifel, covering an area of some 500-600 km2 with c. 220 eruptive centers; and, about 25 km to the east, the smaller East Eifel (Laacher See area) with some 70 volcanic centers distributed over c. 400 km2. Magmas of both fields are moderately potasaic. Most West Eifel magmas are nephelinites and leucitites and have high CaO/A120a-ratios (up to about 1.45) and s7’86Sr-ratios (0.7039-0.7044); in contrast, intermediate (tephritic) and highly differentiated (phonolitic) magmas are common in the East Eifel, where magma chambers of different compositions have developed at varying depths in the crust, as evidenced by a variety of crustal xenoliths. Primitive East Eifel magmas are dominantly basanites with Mg numbers <67, and very high 87’s6Sr-ratios (0.7041-0.7047). The Quaternary Eifel magmas must have come from a different mantle area than the Tertiary magmas, whose 87’s6Sr-ratio ranges from 0.7035-0.7040. Most of the existing seismic-refraction profiles in the Rhenish shield have been reinterpreted using travel time and amplitude information. The general pattern of observed phases can be divided into three types: (1) Throughout the central shield a strong P-phase reflection from the crust-mantle boundary is recorded. The average crustal thickness is 28-29 km and the average crustal velocity is 6.2-6.3 km/set. (2) Beneath the southern part of the shield strongly reflected phases show clear intracrustal and Moho discontinuities. Along profiles crossing major volcanic features the Mdiscontinuity is disrupted and an intermediate intracrustal boundary at about 20 km depth forms the main reflector. Beneath this boundary the velocity increases gradually from about 7 km/set to upper-mantle velocities.

198 (3) For profiles crossing the northern Rhinegraben area as well as for a line east of the Lower Rhine embayment, the observed phases indicate only one major seismic boundary at a depth of about 23 km where the velocity increases to 7.3 km/set. Below this boundary the velocity increases gradually with depth reaching 8 km/set at 27-28 km. The Rhenish shield is part of an incoherent block mosaic of the Alpine foreland. The Alps are under rapid Holocene uplift and sideward extension. Between the Alps and the Rhenish shield there is a major foreland block which is shifted northwestward to be sheared along the Rhinegraben rift with a sinistral polarity. The northern apex of this block wedges in the Rhenish shield and provokes an axial deformation and, conditioned by that, a split-up of the northwestward frame of the shield. It is shown that the shield uplift evolved as a consequence of horizontal deformations of the unit. Subcrustal rift propagation under the tectonically incompetent crustal unit of the Rhenish shield had favored magmatic activity, thermal rise and subsequent plateau uplift. TECTONIC

SETTING

OF THE

SHIELD

(H. Illies)

The “Rheinisches Schiefergebirge” (Rhenish Slate Mountains) in Central Europe is an uplifted plateau in which a dense drainage pattern has engraved several generations of erosional landforms. It is composed of a thick series of slates, graywackes, quartzites, limestones, and locally volcanics of mainly Devonian and Lower Carboniferous age. The series, generally some 4000 m but locally up to 10,000 m thick, was primarily the infill of the RhenoHercynian geosyncline. The strata were strongly folded and slightly metamorphosed during the Hercynian orogeny between 330 and 300 m.y. ago (Ahrendt et al., 1978), forming a part of the Variscan mountain range. The first phase of uplift and erosion occurred during Upper Carboniferous and Lower Permian. Subsequently, Upper Permian and Mesozoic seas marginally flooded the shield unit while its central part remained a flat emerged plateau. Tertiary marine transgressions temporarily had reached peripheral sections of the platform, but most of the area was covered by fluvial gravel accumulation (Quitzow, 1959; Liihnertz, 1978). Wide-spread dominantly basaltic eruptions pierced the block during Upper Eocene, Oligocene and Miocene (Cantarel and Lippolt, 1977). Since the Pliocene the ancestral Rhine and Mosel rivers traversed the platform and their fluvial terraces are preserved today up to 300 m above the present river levels. From the Upper Pliocene on, the whole unit was subjected to a regional uplift at varying rates. The uplifted plateau which we term the Rhenish shield, extends over an about 100 X 200 km wide area between Frankfurt and Bonn, Kassel and Luxemburg (Fig. 1). It includes the “Rheinisches Schiefergebirge” and some smaller marginal areas overlain by outliers of the peripheral Triassic sedimentary Fig. 1. The Alpine foreland of Western Europe is traversed by a belt of seismotectonic activity. Rift segments like the Rhinegraben (south of Frankfurt) and the Lower Rhine embayment (north of Bonn) serve as sites of strain release of a regional stress field that is generated by Holocene Alpine tectonics. The uplifted plateau of the Rhenish shield (between Frankfurt and Bonn) interrupts that rift belt physiographically, but clusters of seismic epicenters indicate that the intraplate belt is active under the coherent block of Hercynian folded slates.

199

0 Fault zone. reottw3ted 5!“C+? .. ..‘. J’-

Pimene

Earihquoke epIcenters ca 1200 - 1970 Alptne

fold

belr

50

100 km

,

_

contour of Ou&xnary of the

Rhentsh

shield

LIplIft

“00

cover. The Rhine and Mosel rivers and a dense pattern of their tributaries have cut deep antecedent river valleys during the uplift. Timing and rates of uplift are inferred from a sequence of well dated river terraces (Bibus and Semmel, 1977). This regional uparching was associated with wide-spread Quaternary volcanic action mainly in the western half of the shield which started about 600,000 years ago (Frechen and Lippolt, 1965; Windheuser, 1977). The youngest volcanic episode culminated about 11,000 years B.P. with the eruptions of the Eifel maars and the Laacher See volcano (Erlenkeuser et al., 1972; Duda and Schmincke, 1978). In terms of regional tectonics the uplifted unit is part of an incoherent mosaic of blocks in the northern foreland of the Alpine fold belt @lies and Greiner, 1978). The southern rim of the Rhenish shield lies about 300 km north of the northern margin of the Alps (Fig. 1). The shield is surrounded by rift valleys and fault troughs. Active rift segments are framing the shield at its southern and northern end: Rhinegraben and Lower Rhine embayment. Extinct rift segments are flanking the unit: the Luxemburg embayment in the west, and the Hessen depression in the east. The active rift segments are linked together seismo-tectonically along a northwest-trending “subplate boundary” which traverses the Rhenish shield (Illies and Greiner, 1979). Physiographically, the uplifted block of the Rhenish shield crosscuts the belt of active rifting. Perhaps, it was the tectonic~ly incompetent behavior of the slightly metamorphosed Devonian slates of the unit which impeded surficial rift valley progression. Like a front flap of a lederhosen protecting the fly of the trousers, the incompetent rock material of the upper crust covers axreactive rift belt underneath. Graben formation and up-arching of the Rhenish shield are interrelated in space and time. Periods of higher rates of uplift in the shield coincided with periods of accelerated faulting and subsidence in the grabens. The borders between both kinds of tectonic action are generally gradual rather than discrete. Flexure-like transitions dominate even at places where major fault zones are bordering the shield. Where the vertical motions are taken up by faulting, the displacement generally occurs over sets of splintering normal faults in about 3---5 km broad zones. A belt of earthquake epicenters traverses the shield about parallel with the Rhine river (Ahorner, 1975). In some areas the seismotectonic action seems to be associated with normal faulting of Pleistocene and sometimes Holocene age (Stengel-Rutkowski, 1976). This intra-shield deformation is diffuse because lengths and vertical throws of faults are mostly small. Only the Neuwied basin, centered between both ends of the active rift valleys, shows larger vertical motion up to about 300 m. THE TERRACES OF THE UPLIFT (A. Semmel)

RHINE

RIVER:

REFERENCE

LEVEL

FOR

THE SHIELD

The question of whether or not an area has been influenced by young tectonic movements to a great extent can often be answered by the study of

201

fluvial terraces. Frequently the Middle Rhine Valley has provided an examplary model for this. The term middle Rhine Valley comprises that section of the river, which is situated in the “Rheinisches Schiefergebirge”. As in all formerly periglacially influenced valleys of Central Europe, the Middle Rhine Valley formed many older fluvial terraces, indicating frequent alternations of accumulation and erosion during the Pleistocene. For many years attempts have been made to reconstruct these numerous older valley bottoms, which were dissected, but obviously were also exposed to stronger postsedimentary upwarping. Respective gradient profiles can be found in many publications, e.g., Quitzow (1959) and Kopp (in: Schmidt-Thom~, 1972, p. 20). An extremely strong upwarping was postulated for the Early Pleistocene Main Terraces (“Hauptterrassen”). In the area of Andernach and north of Bingen (Fig. 2) uplift supposedly reached almost 80 m, thus showing an opposite river gradient. According to these authors the younger Middle Terraces showed a distinctly smaller amount of upwarping, and the Late Pleistocene Lower Terraces showed no upwarping at all. In a more recent publication Birkenhauer (1971) expressed contrary opinions. He came to the conclusion that the formerly assumed upwarping of the Rhine terraces arose from an incorrect correlation of the former valley bottoms. The question, as to which of these diverging opinions is correct, is of great impo~~ce for the research programme “Vertical uplift and its causes, e.g., in the area of the Rhenish shield”. Consequently, a new mapping of the remaining terraces was carried out as part of the research programme mentioned above. Emphasis was placed on those terraces situated above the Rhine gorge, as only in this topo~aphic~ position were former valley bottoms preserved to such an extent that a definite correlation was possible. The essential results of this costly and painstaking research (as numerous drillings were involved), are reported elsewhere (Bibus and Semmel, 1977). Therefore, the following conclusions can be drawn: In the Middle Rhine areas, situated above the narrow Rhine gorge, five to six Early Pleistocene river terraces can be reconstructed, With the exception of the Neuwied and Linz basins, the terraces do not show any considerable postsedimentary dislocations (Fig. 2). Strangely enough, in contrast to the younger river terraces, these terraces do not possess any gradient. Gradients, however, must have existed at the time of their formation in order to transport the gravel masses. Possibly the gradients were eliminated by a stronger uplift in the northern part of the “Rheinisches Schiefergebirge” block, which began after the development of the old terraces. This uplift presumably also caused the incision of the narrow Rhine gorge. The youngest of the terraces mentioned above (tRS according to the nomenclature of Bibus and Semmel, 1977), in many places identical with the Younger Main Terrace (“Jiingere Hauptterrasse”) of older studies, shows, in contrast to the older terraces, a distinct gradient beneath the estuary of the Brohl valley. Bibus considers this as a result of stronger tectonic subsidence in the Lower Rhine area, which also triggered an increased incision in the northern part of the “Rheinisches

202

203

Fig. 3. Tectonic dislocation in the tR4-gravel near St. Goar. 1 = tRe-gravel; 2 = clay-layer in the tR4-gravel; 3 = fluvial loams; 4 = loess; 5 = tuff layer (not dislocated) in the loess.

Schiefergebirge”, before the formation of tR5. This terrace is regarded to be an equivalent of the HT,-terrace described by Schniitgen (1974) from the Lower Rhine area. In Fig. 2 the altitudes of the Early Pleistocene terraces t~4 and tR5 (“Main Terraces” in the sense of some older publications), and the recent valley bottom are shown. With the exception of the above mentioned areas of subsidence, faults can be recognized only in few instances within the fluvial Pleistocene sediments. They also show that tectonic movements in later times were extremely limited. One of the few examples of tectonic dislocations (apart from the Neuwied basin area) can be seen in Fig. 3: North of St. Gear, gravels of the t,,-terrace (which is partly an equivalent of the Older Main Terrace (&tere Hauptterrasse”)), and covering sediments are dislocated. Whereas the top of the gravel accumulation shows a displacement of c. 0.5 m, a basaltic tuff layer, which is embedded in the hanging loess, was displaced only a few centimeters. The deposition of this tuff coincides roughly with the youngest parts of the Upper Middle Termce (“Obere Mittelterrasse”). THE LATE QUATERNAR~

EIFEL VOLCANIC

FIELDS (H.-U. Sehmincke)

Introduction The Quaternary volcanoes of the Eifel represent the youngest volcanism in Central Europe. They have been studied for over 200 years but modem work focusing on currently relevant volcanologic and petrologic problems is

scarce. The present report is a rather preliminary summary of work done chiefly in connection with the recent research program on vertical movements of the Rhenish shield. Physical characteristics

of volcanic fields

The Eifel area, which lies west of the Rhine river and north of the Mosel river, contains two volcanic fields of Quaternary age, separated by areas of Tertiary volcanism. The basement consists chiefly of Devonian slates and sandstones with some Devonian limestones and outliers of a Triassic sandstone cover in the western part. The East Eifel, centering around the Laacher See, is c. 30 km in diameter and covers an area of approximately 400 km?. The larger West Eifel volcanic field covers an area of about 500--600 km” and extends over a distance of c. 50 km in a NW-SE direction from Bad Bertrich at the Mosel river to Ormont near the Belgian border (Fig. 4). The two fields are about 25 km apart with a few isolated eruptive centers in between. Eruptive centers are defined here as volcanic vents - or several closely grouped vents - that form one volcanic cone or a composite cone. Exogeneous and endogeneous domes within calderas are considered as separate centers. About 220 eruptive centers occur in the West Eifel and c. 70 centers in the East Eifel (Schminke, 1977a, b; Duda and Schminke, 1978). Spacing between centers is generally less than 1 km but becomes

Fig. 4. Late Quaternary West Eifel and East Eifel volcanic fields, each dot representing one eruptive center. Data for East Eifel from Duda and Sehmincke (1978) and for West Eifel from H. Mertes (unpublished).

larger and more irregular towards the margins of the fields. Both the volume of erupted material and the number of volcanoes increase towards the center of both fields. Likewise, the proportion of intermediate and highly differentiated products is highest in the central part of the fields. This suggests that the melting anomaly in the mantle and/or the structural conditions in the crust are such that most magma is produced - and/or is stored in the crust resp, is released to the surface - predominantly near the central areas of the fields. The total volume of ~~~~c magmas in the West Eifel is roughly 0.5 km3 and is probably less in the East Eifel. The volume of ~~ono~~~~cto tephritic magma erupted from the Laacher See volcano alone, however, may be about 1.5 km3 (Schmincke et al., 1978a). Age

of volcanism

The West Eifel “Quaternary” volcanism extends from Late Pliocene (2.7 m.y.) throughout the Quaternary (Mertes and Schminke, 1979) with the youngest eruption being about 8000 years old (Lorenz and Biiehel, 1978). Data are not sufficient to define separate phases of volcanism, or spatial or compositional relationships. Previous K/Ardata obtained on East Eifel lavas and subvolcanic intrusions combined with the occurrence of characteristic heavy minerals in Rhine terrace conglomerates led Frechen (1971) to postulate 3 phases of magmatic activity: an older one to the west, of very alkalic volcanic rocks lacking plagioclase, one older than about 0.4 m.y. ago and younger phases, with much volcanic activity occurring between about 25,000 and 75,000 years ago (Freehen, 1971, pp. 30-31). Stratigraphic studies by Windheuser (1977) suggest that many volcanoes are somewhat older (between about 0.2 and 0.3 m,y. ago). Recent K/Ar-determinations indicate that very little - if any - mafic volcanism took place less than about 0.19 m.y. ago with most volcanoes having erupted between 0.2 and 0.5 m.y. ago. It is clear, therefore, that volcanism started in the Late Pliocene (West Eifel) and culminated during the Late Quatemary, and that the Eifel can be regarded as an active volcanic area, considering that the latest activity occurred less than 10,000 years ago. Eruptive mechanisms The West Eifel volcanic field is the type area of the maar volcano - an enigmatic, sometimes water-filled crater surrounded by a low rim of pyroelastics in which country rock fragments often predominate. Traditionally thought to result from eruptions of rn~ati~ gases, chiefly CO*, (Frechen, 1971), there are now many lines of evidence for their phreatomagmatic origin (Lorenz, 1973; Schmincke et al., 1973). Moreover, many cinder cones that are deeply dissected by quarries show evidence for one or more phreatomagmatic phases during their evolution, usually in their beginning (Schmincke, 1977b).

nephelineleucitite

Steinbeuel

Lot. Mosenberg

olivinenephelinite Buerberg

melilitite

99.25

38.62 2.85 11.93 6.07 5.69 0.23 9.47 15.42 4.29 3.26 0.94 0.32 0.16

41.28 2.49 11.46 3.95 7.05 0.20 14.85 12.48 3.62 1.90 0.89 0.20 0.05

100.46

M 92

M 77

Gossberg

leucitenephelinite

99.65

43.04 2.48 12.04 9.93 0.27 0.18 10.15 13.73 2.62 3.57 0.56 0.93 0.15

NM 353

Kyller Hohe

melilitenephelinite

99.74

40.07 2.41 11.50 6.04 4.69 0.20 10.24 16.07 3.13 2.33 0.96 1.86 0.24

NM 355

For more recent analyses of East Eifel volcanic rocks see Duda and Schmincke (1978).

99.00

Rock name

41.59 2.67 11.71 5.82 5.40 0.21 10.08 14.18 3.18 3.31 0.63 0.25 0.04

M 49

West Eifel

Total

co2

H20+

p205

Fe0 MnO MgO C&O NazO KzO

Fez03

4203

SiOr TiOz

Sample no.

Representative chemical analyses of lavas from the East and West Eifel.

TABLE I

Dachsbiisch

basanite

99.39

42.81 2.55 13.72 7.49 3.37 0.18 9.69 11.97 3.03 3.30 0.55 0.32 0.41

NM 181

East Eifel

Rothenberg I --

tephrite

99.04

44.00 2.57 16.01 5.11 6.01 0.21 5.13 10.51 4.14 3.81 0.84 0.59 0.11

NM 387

207

Strombolian/Hawaiian eruptive mechanisms were dominant in most cinder cones and 1*30% of the eruptive centers in both fields have produced lava flows, rarely as much as 10 km in length. Plinian eruption columns must have developed in three dominantly phonolitic calderas in the East Eifel (Rieden, Wehr, Laacher See). Tephra from the largest and youngest of these eruptions (Laacher See) was distributed over much of Central Europe (Frechen, 1971).

mineralogical and chemical composition The chemical composition of rocks from both suites (Table I) corresponds to moderately potassic alkali basalt, with basanite being dominant in the East Eifel and nephelinites, melilite nephelinites and leucitites in the West Eifel where basanites appear to occur only in a southern sub-province (Bad Bertrich-Wartgesberg) (Duda and Schmincke, 1978; Schmincke et al., 1978b). The ratio K20/Na20 is about unity in the West Eifel but typically more than 1 in the slightly more potassic East Eifel magmas. Another major contrast between both fields is the relatively large number of eruptive centers of intermediate and highly differentiated composition in the East Eifel and their near absence in the West Eifel indicating -that crustal magma chambers were able to develop in the East Eifel but not generally in the West Eifel. This difference is thought to result from contrasting tectonic styles: extension in the East Eifel that lead to formation of the Pliocene-Quaternary Neuwied tectonic basin, which is downfaulted up to 300 m and to the development of magma chambers at different levels in the crust. Studies of crust-derived xenoliths show that the more differentiated lavas contain dominantly fragments from increasingly higher crustal levels and a rather speculative subdivision of reservoirs is attempted in Fig. 5. Granulitefacies xenoliths and schists from two localities were recently studied in detail by Okrusch et al., 1979. The tectonic style in the West Eifel appears largely basement-controlled with possibly right-lateral strike-slip faulting with only minor development of magma chambers (Schmincke et al., 1978a, b). This model is also suggested by a notable diflerence in type of nodule populations. Ultramafic nodules (lherzolites, dunites, harzburgites, wehrlites) are common and locally of large dimension in several West Eifel volcanoes, suggesting rapid and direct ascent of the magmas to the surface. The most famous locality is the Dreiser Weiher (e.g., Paul, 1971; Sachtleben et al., 1978; Stosch and Seek, in press). Amphibole-bearing lherzolite nodules from Dreiser Weiher are REE~n~ched thought to result from enrichment of a depleted lherzolite by a fluid or liquid phase (Stosch and Seek, in press). Peridot& nodules are mostly small and rare in East Eifel volcanoes (Heidekopf, Sat&&erg), except at Kempenich (Duda and Schmincke, 1978). Nodules dominated by clinopyroxene, phlogopite and ~phibole often with cumulate fabric occur in both East and West Eifel lavas. They are interpreted

EAST

EIFEL

MAGMA

NEPHELINITE ROCK

RESERVOIRS BASANITE

TEPHRITE BASANITE

PHONOLITE TEPHRITE

TYPES

GNEISS

-G?~A<TE MOHO

-

2 5 z

00

-----

ASTHENOSPHERE Fig. 5. Highly speculative section through crust and upper mantle beneath East EifeI volcanic field showing approximate location of magma reservoirs based on main type of crustal xenoliths. Bars to left of each volcanic-magmatic system show types of xenolith in erupted products. For simplicity the asthenosphere is drawn at 100 km depth and only one magma chamber (size not to scale) is show for each system. Crustal thickness is approximately as indicated by seismic data (Giese et al., 1976). Thickness of crustal rock types (phylhte to granuhte) is unknown except for Devonian which is estimated to be 4-5 km thick. Mantle-derived peridotite xenoliths encompass wehrlite, harzburgite and lherzolite, pargasite being a common phase. Bare granulite xenoliths in phonolitic pyro&sties may indicate that granulite-facies rocks may extend to high crustal levels.

by Lloyd and Bailey (1975) as mantle rocks metasomatized by alkali transfer while Becker (1977) holds that they formed in crustal magma chambers at less than 5 kb pressure. The main phenocryst phase in both East and West Eifel “basaltic” lavas is titanaugite with volume percentages up to 20%. Olivine and lesser amounts of phlogopite are ubiquitous but generally make up less than 5 vol. % except in a few rocks with abundant olivine. In addition to these, intermediate rocks contain kaersutite in place of olivine, joined by ti~om~etite, sphene and apatite with increasing differentiation. The highly differentiated rocks contain plagioclase, alkali feldspar and hauyne or nosean as additional phenocryst phases. Leucite and/or nepheline occur in most rock types as groundmass phase only, with leucite predominantly in rocks with KzO/ NazO > 1. Fractionation calculations for a suite of East Eifel lavas show that the nephelinite magmas cannot be related to the basanite and tephrite magmas by crystal fractionation at low pressures (Duda and Schmincke, 1978). How-

209

ever, the major element chemistry of the tephrite magmss can be derived from the basanite magmas using the observed phenocryst assemblage. This computational result is strongly supported by the change in composition in several volcanoes from tephritic in the early stages to basanite in the last stages. The volume and spatial relationships between tephritic and basanitic rocks support this conclusion. Recent high-precision 8”86Sr-isotope ratios indicate that the mantle beneath the East Eifel from which the primitive basanite magmas were derived - or with which they last equilibrated - is slightly more radiogenic (0.7041-9.7047) than that beneath the West Eifel (0.7039-0.7044). Both these areas are more radiogenic than the mantle that fed the Tertiary alkali basaltic lavas (“Sr/s%r = 0.7034-0.7040). This might be expected judging from the slightly higher K- and Rb-contents of the East Eifel and much lower K- and Rb-contents in the Tertiary magmas. The Quails magmas must thus have come from different mantle areas than the Tertiary magmas. 143’144Nd-ratiosrange from 0.51277 in the Tertiary alkali basalt to 0.512600.51265 in the East Eifel Quatemary lavas (Schmincke et al., 1978b, Staudigel and Zindler, 1978). Cause, nature and timing of the enrichment or metasomatic event in the mantle beneath the Eifel is unknown as is the relationship between the enrichment process that led to a mantle yielding the potassic magmas (Duda and Schmincke, 1978) and that leading to widespread amphibole formation (Stosch and Seek, in press). 16

0 WEST EIFEL 0 EAST EIFEL BASANITE 0 EAST EIFEL NEPHELINITE A MIOCENE ALKALI BASALT

Fig. 6. CaO versus Alz03 for late Quaternary West Eifel (open circles) and East Eifel (filled circles) rocks and Miocene basalts (triangles). Group I: melilite nephelinites; Group II: nephelinites and leucitites; Group III: basanites in southern part of field; Group IV: northern basanite group; Group V: southern basanite group.

210

Primitive magmas of both East and West Eifel volcanic fields, while similar in many major and minor elements, show striking differences in some, parboth fields ticularly CaO and Alz03, with almost no overlap between (Fig. 6). This is tentatively interpreted as due to differences in the degree of partial melting with garnet remaining in the residue in the melting region beneath the West Eifel. REE-patterns of some West Eifel lavas show lower heavy REE-concentrations than most East Eifel lavas, in support of the above suggestion, but in general there is much overlap; more data are clearly needed.

Discussion Duncan et al. (1972) suggested that the Quaternary volcanic Eifei province was situated above a hot spot (Eifel plume) that had previously fed Tertiary volcanic fields extending eastward for some 700 km to Silesia. They based their hypothesis on a supposed decrease in age from east to west. Abundant age data published in the last few years do not support this hypothesis (e.g. Cantarel and Lippolt, 1977). Duda and Schmincke (1978) emphasized that many Tertiary volcanic fields in Central Europe show striking similarities to self-sustained patterns of volcanic evolution which is well known from oceanic islands such as Hawaii and Gran Canaria: a main phase, followed, after erosional intervals, by one or more phases of much less voluminous, more primitive and undersaturated usually nodule-bearing lavas. The cycle of volcanic phases, so typical of some oceanic islands, might be interpreted as a self-perpetuating system in which the notable change in volume and composition might be due to such factors as downward migration of the zone of m.elting, enrichment of the previously partially melted domain in incompatible elements etc. during erosional intervals or by later fractures extending to greater depth. Volcanic loading and unloading have been suggested by some workers to trigger post-erosional volcanic phases, but regionally widespread causes such as epeirogenic vertical movements have also been invoked as resulting in second or third stage volcanism (Duda and Schmincke, 1978). Volcanic loading is unlikely to be important in the Eifel intraplate province because of the small volume of volcanics and the thick crust (Fig. 5). The most obvious recent tectonic events that affected the area and may contribute to magma generation are the formation of the Rhinegraben rift system and the subsequent regional uplift of the Rhenish shield. A presentday active seismotectonic belt extends from the northern termination of the Rhinegraben near Frankfurt across the uplifted block joining the Lower Rhine embayment in the Cologne area (Ahorner, 1975; Illies and Greiner, 1979). Especially interesting is a temporal correlation between the highest rate of uplift that started about 0.6 m.y. ago as evidenced by the main Rhine terraces in the area (Bibus and Semmel, 1977) and the phase of most abundant volcanic activity (see above). This suggests that the mantle processes

211

that lead to uplift also triggered melting and/or facilitated magma ascent to the surface, The active Fr~kf~--~olo~e tensional seismo~~tonic belt (see above) which crosses the East Eifel volcanic field combined with geochemical evidence suggesting slightly higher degrees of melting (shallower depth?) in the mantle might both be the result of asthenosphere rising along this area, possibly marking a future zone of rifting. Perhaps the Quaternary melting events were triggered by the lithosphere “sliding off” mantle material rising beneath this seismotectonic belt? In any case, the mantle composition from which the Quaternary magmas were derived was quite different from that of most Tertiary melting anomalies in Central Europe. Future work relating volcanism, tectonism, seismicity and crust and mantle structure should also attempt to interpret the relationships between Tertiary and Quaternary tectonic and volcanic events. For example: What is the relationship between tectonic movements along the Alpine front and the East-West belt of Tertiary volcanic fields to the north? Was the Tertiary volcanic activity of the Rhenish shield related to early stages of formation of the Rhine~aben system farther south representing an aborted attempt of the mantle system to extend the rift farther north? Is the striking near-absence of Quaternary eruptive centers in the Tertiary volcanic region between the East and West-Eifel due to “plugging” of potential fissures in the crust, “barren” areas at appropria~ mantle depth (perhaps as the result of the Tertiary melting event), or to inappropriate tectonic and/or thermal regimes in crust and/or mantle beneath this area? CRUSTAL STRUCTURE OF THE RHENISH SHIELD - A REINTERPRETATION OF SEISMIC REFRACTION DATA (C. Prodehi)

Up to now only a few detailed seismic investigations have been performed in the area of the Rhenish shield. Within two former priority programs of the German Research Society “The Deep Structure of Central Europe” (19581964) and “Upper Mantle Project” (1965-1974), explosion seismological investigations of Central Europe were based mainly upon quarry blast observations (Giese et al., 1976). A part of these programs covered also the Rheno-Hercynian zone of the Hercynian erogenic belt where numerous quarries firing large explosions are located (Fig. 7). In order to investigate the available data in detail most of the profiles in the Rhenish shield/RhenoHercynian zone have been jointly reinterpreted by Mooney and Prodehl (1978) using travel time and amplitude information. It can be noted that only a few of the profiles run entirely in the Rhenish shield proper; most partly cross Cenozoic volcanic areas within and east of the Rhenish shield. As the quarries with usable blasts are spread irregularly the normal requirements for seismic-refraction measurements of having reverse and overlapping profiles could only be fulfilled by one line. Also, the position of the profiles is not always optimum with regard to geologic and tectonic settings of the area.

8”

Common-depth pwlt profile

A

A’

Locot~on of cros5 sectmn of fig 9

Location ,BB ’ sectnn

of cross of ftg 10

Fig. 7. Location of seismic-refraction profiles on a generalized geologic map of the Rhenish shield and adjacent areas. AC = Aachen; DO = Dortmund; LE = Liege; KS = K = K&r; GI = Gies+n; SB = Kassel; L = Luxembourg; F = Frankfurt; BR = Brussels; Saarbriicken; HD = Heidelberg; NA = Namur; MS = Miinster; MZ = Mainz; WU = Wiirzburg. A = Ardennes; E = Eifel; H = Hunsriick; HA = Harz; HB = Hessische Senke; MS = Miinsterllnder Bucht; N = Neuwied basin; NB = Niederrheinische Bucht (Lower Rhine embayment); 0 = Odenwald; RG = Rhinegraben; S = Sauerland; SI = Siebengebirge; SN = Saar-Nahe trough; SP = Spessart; T = Taunus; V = Vogelsberg; W = Westerwald.

In spite of these complications the record sections of the numerous profiles exhibit a similar pattern of travel time branches which permit a cl~sifi~tion of the observed record sections in relation to the geologic situation. Though for some profiles the data clearly show that lateral heterogeneities play a major role, the variable quality of the available data and the infrequency of reversing and overlapping profiles does not justify the use of methods more sophisticated than, e.g., flat-layer approximations as shown for all profiles by Mooney and Prodehl(l978). The general pattern of observed phases can be divided into four types, each type corresponding to a distinct type of velocity structure (Fig. 8):

213

m

strong reflectron of 28-30km

ES

weak reflection of 28 - 30 km l

shotpoints

!zzI

strong reflectton from o depth between IO and 22 km

EZI

weak reflectlon from o depth between IO and 22 km

from o depth from o depth

at quorries

o shotpoints

of wide-angle

profile

Fig. 8. Map of the area of the Rheno-Hercynian zone showing main features of crustal structure. Shotpoints: 02 = Hilders; 03 = Gersfeld; 05 = Birkenau; 06 = Adelebsen; 08 = Kir~hheimbolanden; 12 = Romsthal; 13 = Dorheim; I4 = Mehrberg; 15 = Biidingen; 16 = Taben Rodt; I7 = Bransrode; 20 = Birresborn; 21 = Bermel; 22 = Dorndorf.

(1) Throughout the central Rhenish shield and the adjacent “Hessische Senke” (depression) a strong P-phase reflection from the crust-mantle boundary is recorded in regions where the lines of observations do not cross any major volcanic features. The average crustal thickness is 28-29 km, the average crustal velocity, excepting sediments, is 6.2-6.3 km/see, and the crust does not show strong inhomogeneities. This structure is referred to as the Rheno-Hercynian crustal model by Mooney and Prodehl(l978). (2) Beneath the southern part of the Rhenish shield and two areas in the northeast and southeast some structure within the crust is evident. Both an intracrustal and the Moho discontinuity are evidenced by strong reflected phases, the Moho reflection being the stronger one. (3) Along profiles crossing major volcanic features such as Vogel&erg and

central Westerwald composite volcanoes, but not beneath the East Eifel, the producing only weak Mdiscontinuity is heavily disrupted or “smeared” reflections and an intermediate intracrustal boundary at about 20 km depth forms the main reflector for seismic waves. Beneath this boundary the velocity increases gradually from about 7 km/set to upper mantle velocities. (4) For profiles crossing the northern Rhinegraben area as well as for a line from the Siebengebirge through the Rhenish shield to the north, east of the Lower Rhine embayment, the only observed reflection indicates a major seismic boundary at a depth of about 23 km where the velocity increases rapidly to 7.3 km/see. Below this boundary the velocity increases gradually with depth reaching 8 km/see at 27-28 km. The P,-phase, (mantle head wave) is recorded with variable success throughout the area of investigation, It disappears completefy on a line passing the East Eifel volcanoes, but is clear on the lines through Vogelsberg and central Westerwald. Figures 9 and 10 show an attempt to compile the main features of crustal structure along two selected lines. The first line (Fig. 9) crosses the Rhenish shield in an E-W direction (AA’ in Fig. 7). The eastern part is located in the fault trough of the “Hessische Senke”, showing a structure very similar to

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215

the Rheno-Hercyni~ type: a rather homogeneous crust separated from the upper mantle by a well-defined crust-mantle transition zone. To the west, beneath the Vogelsberg, the crustal structure gets more complicated. The main features are a seismic discontinuity at a depth of 20 km and the replacement of a clear crust-mantle boundary by a wide transition zone. Upon entering the Rhenish shield the Rheno-Hercynian structure is again encountered. The area of the central Westerwald, covered by Tertiary basalts, however, shows an anomalous structure similar to that evident beneath the Vogelsberg. In contrast, the area further to the west in the Neuwied basin and adjacent East Eifel, penetrated by alkali-basaltic to phonolitic extrusions, while apparently showing an undisturbed Rheno-Hercynian crustal structure, shows an anomalous mantle: the clear arrivals reflected from the Moho reach an apparent velocity of only 7.3 km/set and P,-arrivals are not at all visible. The second line (Fig. 10) crosses the Rhenish shield in a NW-SE direction (BB’ in Fig. 7). Although it is not evident from the geologic setting, the Rhenish shield east of Cologne shows an unexpectedly anomalous crustal structure: a seismic discontinuity is apparent at about 21 km depth under-

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lain by a smooth and continuous transition to the upper mantle. No data arc available between here and the central Westerwald where line BZ3’ crosses line AA’ (Fig. 9). South of the central Westerwald composite volcano RhenoHercynian crust is encountered. Upon entering the Taunus mountains, a densely recorded common-depth point profile (Fig. 7) revealed a more complex crustal structure with pronounced diseontinuities at about 16 km and 29 km depth (Meissner et al., 1976a; Mooney and Prodehl, 1978). The cross section terminates in the northern end of the Rhinegraben. Here a low-velocity upper crust reaching to a discontinuity at 23-24 km is followed by a high-velocity lower crust in which gradually upper-mantle velocity values are reached with increasing depth. The results obtained by the reinterpretation may differ in detail but in general strengthen and confirm the results published by various authors at earlier stages, as is dicussed in detail by Mooney and Prodehl (1978). Here only two features shall be noted. In the Siegerland area, north of the central Westerwald Tertiary volcanics, Diirbaum et al. (1971) reported strong reflections with 5-7 set two-way travel time. This may correspond to the strong reflector at about 20 km depth found east of Cologne and in the central Westerwald. The particular shape of the velocitydepth distribution for the area of the northern end of the Rhinegraben was already stated by Meissner et al. f1976b). It should be noted that also beneath the southern Rhinegraben proper the crust--mantle transition zone shows a similar structure (Edel et al., 1975; Prodehl et al., 1976). As has been also observed by Giese (1976), the influence of volcanic events on crustal structure is significant (Fig. 9). The crustal structure beneath the East Eifel and the adjacent Neuwied basin, penetrated by alkalibasaltic to phonolitic extrusions, is very different from the structure beneath the central Wes~rwald and the Vogelsberg where Tertiary basalts are exposed. The petrologic differences seem to cause the differing seismic wave propagation. The tholeiitic and basaltic magmas of the Vogelsberg and the central Westerwald may originate within the upper mantle at depths greater than about 50 km. During their eruption through the crust the crust-mantle boundary has been heavily disrupted and to a certain extent been replaced by a boundary at intermediate depths of about 20 km. The alkalic volcanism of the East Eifel has evidently influenced the upper mantle beneath the Moho and so attenuates the propagation of P,-waves penetrating into the uppermost parts of the mantle. There remains, however, a sharp crustmantle boundary thus generating the strong reflection of seismic energy. As is evident from Fig. 8 and from the cross sections in Figs. 9 and 10 in areas where weak or no Moho reflection is found, the intracrustal reflector at about a 20-22 km depth produces strong reflected seismic energy. For the southern Rhinegraben Edel et al. (1975) stated the existence of a highly anomalous crust---mantle transition between 20 and 25 km which may be caused by mass intrusion from the upper mantle and/or by phase transformations. A similar conclusion may also be drawn for the northern end of the Rhine-

217

graben and for the areas of Tertiary volcanism and the area east of Cologne. No crustal information is available for the Lower Rhine embayment proper. Concerning the Rhenish shield proper, the map showing the main features of crustal structure (Fig. 8) also indicates a division of the Rhenish shield parallel to the WSW-ENE directed strike of the Hercynian fold axes. While the southern part - Hunsriick and Taunus - shows some structure within the crust, the central part - Westerwald and Eifel- is characterized by a more homogeneous crust. This is concordant with the lithological composition of the southern part of the Rhenish shield that is mainly composed of low-grade metamorphic quartzites (Ahrendt et al., 1978). Bonjer (1977) has reported that areas such as Hunsriick and Taunus and also the “Hohes Venn” (northern Eifel) are connected with increased micro-earthquake activity. There remain many open questions which cannot be solved with the presently available data. Only little can be stated about the pre-Hercynian basement beneath the Rheno-Hercynian fold unit. A special structure of the crust and the mantle beneath the Tertiary volcanic areas is clearly indicated by the existing data. However, for a detailed solution more data are required. In this context it has to be asked why the “Hessische Senke” which exhibits a thick Mesozoic to Lower Tertiary sediment cover did not take part in the uplifting process of the Rhenish shield, though for both the RhenoHercynian crustal structure has been determined. The results of this investigation are regarded as the basis for future work in order to fill gaps where necessary and to plan carefully towards a detailed investigation of the crustal structure of selected areas and towards the extension of our knowledge into the lower lithosphere. Especially the knowledge of the velocity structure at the base of the lithosphere may be a key for the understanding of the causes of the uplifting processes in the Rhenish shield. MECHANISM

OF UPLIFT

-

GEOLOGICAL

IMPLICATIONS

(H. Illies)

Rhenish shield and the Rhinegraben An active rift belt crosses Central Europe parallel with the course of the Rhine river. Physiographically, but not seismotectonically, the Rhenish shield interrupts this belt (Fig. 1). The strain relief, as observed in the different segments of this belt, by creep, seismic activity, and perhaps heat production, is in response to a regional stress regime, whose magnitudes generally increase towards a stress generator in the Central Alps (Illies and Greiner, 1978). Between the stress-generating Alpine mountain body and the uplifted plateau of the Rhenish shield there is the rift segment of the Rhinegraben, its sinistral shear motion guides the stress and strain transmission towards the more northward adjacent unit of the foreland. In terms of stress magnitudes and strain release (Illies and Greiner, 1979) the Rhinegraben is considered to be the active partner within the match of relative block motions round the Rhenish shield. The graben has a two stage

tectonic history (Illies, 1978). From mid-Eocene to Early Miocene (about 48-18 m.y.) rifting evolved as an extensional process. Graben subsidence and shoulder upwarping at that time were controlled by a regional stress field, its horizontal component of maximum compression was roughly parallel to the average axial trend of the rift. The paleo-stress pattern of that time is inferred from the trend of long basaltic dikes of Paleogene age as well as by the average orientation of the axes of horizontal stylolites (Illies, 1975). Rifting and crustal spreading controlled by this stress regime persisted actively up to about 18 m.y. ago. Subsequently, rifting became extinct, the graben lagoon dried up, and wide parts of its primary sediment fill were eroded by rivers. Graben tectonics were strongly reactivated during the Pliocene, about 4 m.y, ago. After that, rifting was controlled by regional stress conditions similar to those presently inferred from in situ stress determinations (Illies and Greiner, 1978) and fault-plane solutions of earthquakes (Ahorner, 1975). The general orientation of the horizontal component of maximum compression of the active stress field is directed about SE-NW, oblique to the general graben axis. By the rotation of the regional stress field the pre-existent weakness zone of the graben became about parallel oriented to the sinistral shear component of the regional stress field (Fig. 12A). Consequently, normal faults of the first generation of rifting were reactivated into sinistral strike-slip faults in the second stage; vertical slickensides were overprinted by horizontal ones. A series of inner northsouth-trending graben faults (Fig. 1) were reactivated as en echelon faults, i.e. axial shear strain became released by extensional cracks along 2nd order shear planes. Subsidence was also reactivated in those graben segments where the general strike-slip motion produced additional spa&e for lateral extension. In the central segment the sinistral shear motion operated against the grain of the crooked course of the rift valley (lilies and Greiner, 1979). Here, the graben fill was subjected to compressional shear and subsequent uplift (Fig. 1lA). At the northern end of the Rhinegraben near Frankfurt a doglegged rift/ rift offset connects the primary extensional rift segment with its northern continuation, the Hessen depression, 20 km to the east (Fig, 11s). This former %ansform” element was lithologically controlled by a belt of rigid quart&es that forms the southern rim of the Rhenish shield. The rift/rift offset makes the progression of axial sinistral shear infeasible at that place, as it blocks the strike-slip motion (Illies and Greiner, 1978). The Rhenish shield reacted by two different kinematic responses: (1) By a yielding as a whole by means of a slight anticlockwise rotation of the unit. Such motion is indicated by extensional faulting at and parallel with the southern rim of the shield. At its southeastern margin related dextral shear displacements are observed (Illies and Greiner, 1979). (2) By internal brittle fracturing and deformation caused by a wedging of the apex of the block east of the ~hine~a~n into the southern rim of the shield.

Rhenish

shield

graben ftll actually under erosion

@%$?j

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200-250

m

250-300

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300-350

m

’ ‘/

/

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of the graben floor

Fig. 11. The seismotectonic activity of the Rhinegraben is that of a sinistral shear zone. A. Quaternary vertical motions of the graben floor are shear controlled. The present-day rates of vertical motions (0) are about 10 times higher than the average rates throughout the Pleistocene as calculated by the thickness distribution (B). Vertical oscillations of a late Pleistocene river terrace (C) indicate that the higher rates of deformation probably started after the deposition of that terrace. The sinistral shear parallel to the graben axis is blocked near Frankfurt by a rift/rift offset of the primary extensional rift belt.

The sinistral shear motion parallel with the Rhinegraben axis may be explained by the rapid Holocene uplift and the related lateral widening of the Alpine mountain body, the latter in a direction oblique to the graben (Illies and Greiner, 1978). If the observed rates of vertical displacements at, the graben floor are controlled by a horizontal shear motion parallel with the graben axis, the shear rates have accelerated during the recent geological past. In the northernmost segment of the graben, the maximum thickness of Pliocene sediments attains 760 m, that of the Pleistocene about 380 m (Doebl, 1967). In the same segment, the geodetic elevation changes (Fig. 11D) show actual rates of subsidence of about 1 mm/yr and locally even more (Schwarz, 1976). This is about 10 times higher than the average rates during Pliocene and Pleistocene. In the central segment, Pleistocene erosional landforms and related river terraces were uplifted up to 80 m above the base level of erosion at the nearby Rhine river. Positive elevation changes in the same area range between 0.4 and 1 mm/yr (Nivellementnetze, 1977). A discrepancy of the same order is observed at the southernmost graben segment. Here, the Pleistocene infill of local basins is up to 243 m thick (Fig. 11B) (Bartz, 1974). At the same place, the actual rates of susidence are about 1 mm/yr (Malzer and Schlemmer, 1975). Of course, geodetic elevation changes may be sometimes influenced by atectonic factors like alterations of the ground-water table, gas-fields in exploitation, or compaction of sediments (Prinz, 1978). But there is a geological reference level that confirms a considerable increase of the rates of vertical motions during the last 20,000 years. It is given by the late Wiirmian terrace of the Rhine river that covers wide areas of the graben floor. Near the end of this glacial period, about 20,000 years ago, the river started to erode and formed its present flood plain. The difference in elevation between the Holocene flood plain and the late Wiirmian terrace alternates between 0 and 22 m in height within relatively short distances (Wittmann, 1955). In general, the oscillations in the height of the terrace (Fig. 11C) are about parallel with the measured geodetic elevation changes (Fig. 11D). This coincidence supports a predominantly tectonic origin of the elevation changes. Rhenish

shield and Lower Rhine embayment

Near Frankfurt, the northern apex of the triangular Alpine foreland block of Bavaria wedges into the southern rim of the Rhenish shield (I&es and Greiner, 1978). The observed seismotectonic shear rates along Rhinegraben (Ahorner, 1975) plus the proportionally higher rates of creep reveal roughly the present-day rates of horizontal shift. Within the shield, parallel with the northwest direction of shift, there is the Neuwied basin, which was under Pleistocene subsidence and is today an area of tensional seismotectonic activity (Ahorner, 1975). Near Bonn, in the northwestward prolongation of the Neuwied basin the rift segment of the Lower Rhine embayment begins (Fig. 1). Its pattern of Pleistocene growth faulting reveals a fan-shaped crustal

221

spreading (Ahorner, 1962) relative to a pole of rotation about 25 km south of Bonn. As the Lower Rhine embayment is flanked by the separated halves of the northern part of the Rhenish shield, we presume an extensional widening of the northwestward frame of the shield. This implies plastic deformations of wide parts of the shield unit in accordance to the amount of angular rifting. Apparently, it was the lithologically controlled incompetent tectonic behavior of the shield which enabled this unit to such a specificly RhenoHercynian reaction, The Rhenish shield intervenes kinematicly between the active shear along Rhinegraben and the active spreading of the Lower Rhine embayment. The most spectacular about 150 m uplift of the shield had started after the deposition of the about 600,000 years old main terrace of the Rhine river (Bibus and Semmel, 1977). At the same time in the Rhinegraben a transition evolved from a broader scattered shear pattern to a more discrete strike-slip motion and the subsequent accentuation of shear controlled domains of subsidence (Illies and Greiner, 19’79). This caused in the Lower Rhine embayment simultaneous tensional tilt block rotations and related growth faulting (Boenigk, 1978). In accordance with the rapid increase of the rates of tectonic motions in the Rhinegraben during the last 15,000 years, the Lower Rhine embayment geodetic elevation changes on the down-dropped flanks of local tilt blocks attain 2.7 mm/yr, several times higher than the average rates for the Pleistocene period (Quitzow and Vahlensieck, 1955). An epeirogenic response to the inte~ening link of the Rhenish shield is given by its presentday rates of geodetically determined vertical movements (Hein, 1978). Volcanic extrusions that spread out over the shield on the left bank of the Rhine (Fig. 1) began immediately after both tectonic stages, about 600,000 and 15,000 years ago, presumably indicating magmatic decoupling along layers of tectonic remobilization.

The present-day Rhinegraben is under sinistral shear whereas the Lower Rhine embayment acts as an extensional rift valley. The lack of rifting at the intervening Rhenish shield is mainly controlled by the kinematically incompetent behavior of the slaty Rhenohercynian rocks of the unit which impeded ruptural yielding. The basement under the framing rift segments exhibits a more competent behavior which favored rift type faulting. tutus kinematic interrelations are noted between rifting and the intervening plateau uplift of the Rhenish shield. This is especially clear, as the stages of plateau uplift and volcanic activity on the plateau are coincident with the stages of rifting observed in the Rhinegraben and the Lower Rhine embayment. The rift-triggered plateau uplift of the Rhenish shield may be interpreted by a ductile crustal elongation which had caused crustal attenuation and the

subsequent processes of mantle rise and isostatic rebound, The Rhenish shield, in spite of its widely distributed Cenozoic volcanism, persisting up to 8000 years ago, is characterized by neutral to slightly negative heat flow anomalies ranging between about 1.3 and 1.7 HFU (Haenel, 1976). The activation of a magma source in a specific layer of the upper mantle which is demonstrated by the Quaternary volcanic activity (Duda and Schmincke, 1978), did not cause a measurable heat transfer up to the earth’s surface. Subsequently triggered mechanisms like transformations of mineral phases and volcanic eruptions consumed the heat before it had been conducted to the upper crust. Hidden rift progression evolved in the deeper crust of an unit which surficially marks the missing link within the Western European rift belt. Mantle rise and crustal doming had substituted here the physio~aphic phenomenon of rift valley formation. A ruptural incompetent behavior of the involved rock masses of the upper crust impeded the development of rift type morphostructures. To investigate in detail the timing, mode, and mechanism of the observed plateau uplift of the Rhenish shield and to test the models explaining the phenomena, a [i-years multidisciplinary program has been set up by the Deutsche Fors~hungsgemeinschaft. It comprises high precision levellings, a microseismic array, in situ stress determinations, geomorphic and neotectonic studies, heat flow measurements, investigations of the Cenozoic volcanics and their mantle xenolithes, and geochemical studies of the carbon dioxide jets. To investigate the crust and upper mantle, a program of magneto-telluric sounding is incorporated. As the most expensive experiment a long-range explosion seismic profile is in preparation crossing the shield between northern France and the Hessen depression. To synthesize these and other data and to find out the best fitting explanation of the plateau uplift, some programs of finite element modelling are in preparation. ACKNOWLEDGEMENTS

This work gives an interim report on a priority research program of the Deutsche Forschungsgemeinschaft, “The driving mechanism of plateau uplift in the Rhenish shield”. Other parts of this research have been supported by the geothermal programs of the European Community and the Bundesministerium fiir Forschung und Teehnologie. We thank P. v.d. Bogaard, K. Bonjer, G. Brey, G. van Kooten, T. Schmitt, and G. WSrner for discussions and critical comments, and E. Bibus, G. Greiner, H. Mertes and W.D. Mooney for release of unpublished data. The final draft of the manuscript has benefited from our discussion with Celal Sengijr. REFERENCES Ahorner, L., 1962. Untersuchungen zur quart&en Bucht. Eiszeitalter Ggw., 13: 24-106.

Bruchtektonik

der Niederrheinischen

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