GEOID ANOMALIES

Geoid and gravity data provide information on the lateral variation of subsurface density (see also Chapter 4) and thus yield essential information on the local and regional state of isostasy, and the extent to which crustal and lithospheric masses may be isostatically compensated or maintained by deep geodynamic processes.

A number of geoid maps exist for the European area. The one shown here (Figure 5-10) is from Marquart and Lelgemann (1992) and is based on the global geoid deduced from SEASAT and gravity data expanded in spherical harmonics to degree 180 (Rapp 1981). In order to emphasise components related to regional crustal and lithospheric structures, a low pass filter with spherical harmonic coefficients between degree 10 and 180 (wavelengths between about 200 and 3800 km) has been applied. We see a strong negative anomaly

«-10 m) in northem Europe centred over Fennoscandia. Central and southern Europe show a positive anomaly of 5 to 7 m amplitude. A number of minor regional highs correlate with high topography (e.g. southern Norway, Western Alps, Italy, and Yugoslavia), and negative anomalies may be associated with sedimentary basins (e.g. North Sea and Po basin).

Marquart and Lelgemann calculated the geoid effect of isostatically compensated to-pography and crustal thickness variations for the area shown in Figure 5-10 and modelIed geoid variations along the main EGT seismic profiles, considering also the shorter wavelength variations. They concluded that the main features of the residual geoid field can generally be explained by topography and crustal thickness variations that are isostatically compen-sated at depth. In modelling regional as weIl as local variations, the combined effect of the crust and the lithosphere, and in some areas lateral density variations within the crust such

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400N~~~----~~~--~~~--~~~~~~~~~~~

10⁰W

Figure 5-10. Geoid height anomaly map of Europe, low passfiltered between spherical coefficients 10 and 180, after Marquart and Lelgemann (1992). Contour va lues in m.

as those associated with sedimentary basins, were emphasised.

The particular features of the Fennoscandian and the Alpine areas of significant but very different recent activity are discussed separately in the following seetions. The complex structures and recent activity along the southemmost parts ofthe EGT and adjacent areas are treated in subsequent seetions ofthis chapter in relation to recent volcanic activity (Section 5.4) and transient heat flow (Seetion 5.5) and again in Chapter 6.6 and 6.7.

IEUROPE'S LITHOSPHERE - RECENT ACTIVITY 127

2'E 8' 14' 20' 26' 32' 2'E 8' 14' 20' 26' 32'

(a)

8'E 14' 20' 26' 8'E 14' 20' 26'

Figure 5-11. (a) Present rates ofvertical crustal movement in Fennoscandia obtainedfrom repeated precise levelling surveys and tide gauge records, after Ba/ling (1980). Dots show locations of tide gauge stations. Dotted fines indicate areas lacking levelling profiles or sufficient tide gauge data. Rates of movement in mma'/ are relative to mean sea level.

(b) Regional Free Air gravity an oma lies over Fennoscandia correctedfor effects oft opogra-phy, after Balling (1980).

5.3.2 FENNOSCANDIAN UPLIFf

Less than 20000 years ago great parts of northem Europe were covered with ice sheets probably reaching thicknesses in central parts of about 3000 m. The glacial maximum seems to have occurred between 20000 and 25000 years aga and deglaciation between 18000 and 10000 years aga (Denton and Huges 1981, Fjeldskaar and Cathles 1991). Through the glaciation and deglaciation cycles, nature is here providing a fascinating naturallaboratory in which to study the response of the Earth to 10ading and unloading. Due to gravitational disequilibrium during and after deglaciation (unloading), masses within the Earth are redistributed, resulting in rapid vertical crustal movements and changes in sea level. The response of the Earth reflects a complex interaction between ice loads, water loads and deep masses, and is controlled by its deep rheological propel1ies and the loading history . Observation al data constrain mantle viscosity and may provide information on the extent to which the mantle density profile is near adiabatic (see Peltier 1982, 1989 for recent reviews).

The present rate of uplift in Fennoscandia and its uplift history are known in considerable detail (see Mömer 1980, 1990, for reviews). Classical studies include those of De Geer (1888/90), Haskell (1935) and Liden (1938). An essentially eJliptical region of emergence about 1800 km in length and 1200 km in width (Figure 5-11) is weIl established from repeated precise levelling over large areas of Sweden and Finland and for profiles in Norway and Denmark, supplemented for the whole area by tide gauge station records. Generally tide gauge and interpolated levelling results agree within ±0.5 mma-^I. Uplift rates shown in

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Figure 5-11a are relative to mean sea level for the period of observations, mainly the past 50 to 100 years. If regional sea level is rising by 1-2 mma-^I which seems a reasonable estimate for 'global eustasy' (Warrik and Oerlermans 1990), these figures must be added to obtain values of 'absolute' uplift relative to the Earth's centre. In addition, the uplift induces a regionally varying geoid (sea level) rise due to subsurface mass movements, which in the central area of uplift is estimated to be between 0.5 mma-¹ (Ekman 1991) and 0.8 mma-¹ (Sjöberg 1989). Absolute uplift values are thus reaching a maximum of 10-11 mma-¹ in central Fennoscandia, with peripheral areas of submergence of 1-2 mma-^I. The highest observed present shoreline (Ängermanland, northem Sweden) dated to about 9250 years BP is observed at 281 m above present sea level. By extrapolating shore line curves into the area of maximum uplift, Mömer (1980) estimated the total uplift in the centre of the area to be between 800 and 850m.

Although the glacial nature of the uplift is generally accepted there has been much debate over the years on the character of the deep processes and the isostatic state. Thus Jeffreys (1970, 1975) found the Fennoscandian uplift to be inconsistent with the hypothesis that it is due to viscous flow, a conclusion he based mainly on gravity observations. His data showed a positive Free Air gravity anomaly (about 10 mGal) in Scandinavia, which should indicate sinking rather than uplift. Mömer (1990) argued that the 'true glacial isostatic factor' died out some 4500 years ago, and that the present uplift may be driven by phase boundary displacements, and/or readjustments within the lower lithosphere. Much of the debate is related to the problem of gravity and isostasy and to what extent a mass deficiency exists.

Balling (1980) demonstrated that there is a regional gravity low over Fennoscandia (Figure 5-11 b). Free Air gravity anomalies corrected for a regional positive correlation with topography due to the combined effects of masses above sea level and compensating masses at depth (long term regional isostatic compensation of topography along the Scandinavian peninsula) show a longer wavelength negative anomaly of 15-20 mGal covering the region of land uplift. Due to a least squares definition of residual gravity, there is in Figure 5-11 a a 'balance' between positive and negative anomaly areas, and the zero contour does not represent isostatic equilibrium.

The geoid also shows a residuallow over Fennoscandia (Bjerhammar 1980, Marquart 1989, Peltier 1989, Marquart and Lelgemann 1992, see Figure 5-10). The regional geoid shows a marked NW gradient in this area, and the geoid residuallow is found after filtering.

The low covering the area of uplift seems to be between 5 and 10 m. The centre of the low and to some extent its amplitude, depends upon the 'harmonic window' applied (cf.

B jerhammar 1980). Peltier (1989) showed geoid lows over the northemmost part ofthe Earth including Fennoscandia and Laurentia. He applied harmonic coefficients in the range 10-22 which, compared with the map in Figure 5-10, contain less effect from the crust. His map shows a geoid low of about 6 m covering that part ofFennoscandia which is in astate of uplift.

Mömer (1990) argued for the possibility that major parts of the gravity and geoid lows are related to crustal and lithospheric thickness variations and not to a residual glacial isostatic uplift. Marquart (1989) found a clear correlation between the geoid low and crustal thickness and only small geoid variations are left after she had corrected for isostatically compensated crustal thickness variations and topography. Marquart used existing regional Moho maps, which indicated a close correlation. However, recent compilations of crustal thickness variations in Fennoscandia including new results from seismic refraction and reflection investigations (see Figures 3-3, 3-4 and 3-17) show a broad region of deep crust not only beneath the central part of Fennoscandia but also in central and eastem Finland with Moho depths to 50-60 km. Thus we observe no clear general correlation between the regional

I

Figure 5 -12. Present rate of uplijt, and observed and calculated secular change of gravity across Fennoscandia along a fine at about 63 oN, after Sjöberg (1989). Fi/led circles are ob-se/Ted values, with standard error bars; open circles are calculated values. Observations are from an 18 year period. A mantle f10w model is used for the calculations.

129

gravity and geoid lows and crustal thickness variations, in particular not for the central and eastern part of the area. A certain contribution from the combined effect of crustal and lithospheric thickness variations should not be excluded, but modelling is difficult.

Due to the observations of generally small free air gravity anomalies we may conclude that crustal thickness variations are largely isostatically compensated, but the depth and nature of compensation are not known. Modelling a residual gravity, corrected for the effect of the crust and the compensating masses, including lithospheric thickness variations, involves the calculation of differences between very large numbers and is associated with significant uncertainty. Anomalies of the order of 10-20 mGal are not easily resolved.

Repeated precise measurements of gravity along lines crossing the Fennoscandian region are being carried out. Aperiod of about 20 years has been covered and observations are beginning to yield significant results on the secular change in gravity (Mäkinen et al. 1986, Sjöberg 1989). Sjöberg's analysis for a line crossing Fennoscandia at about 63°N (the best line) shows in Figure 5-12 a reasonable agreement between observations and theoretical values using a viscous flow model. He obtained theoretical values of secular gravity variation typically about -0.18 mGalmm-¹ (absolute uplift) to comparewith -0.16±0.04 mGalmm-^I calculated from observations. We recall that the theoretical Free Air gravity gradient (no mass flow) is -0.31 mGalmm-^1• Estimates of the maximum remaining uplift range from 40-50 m (Fjeldskaar and Cathles 1991, Ekman 1991) to about 130 m (Balling 1980).

In arecent study of the Fennoscandian uplift, Fjeldskaar and Cathles (1991) investigated a number of viscosity models including the two 'extremes' , a uniform viscosity mantle (deep flow) and a rigid mantle overlain by a low-viscosity astheoosphere (channel flow). They fouod that the best fit to the observed present uplift pattern came from a model with mantle

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Molasse Basin Molasse Basin

Profile

Figure 5-13. (a) Profiles across the Alps and surrounding regions showing average elevations and observed and calculated Bouguer gravity anomalies_ Observed Bouguer anomaly values, indicated by triangles (with terrain corrections applied) or by dots (with no terrain corrections), are compared with Bouguer anomalies calculated assuming local isostatic equilibrium, rafter Lyon-Caen and Molnar (1989), eproducedwith permission from the Royal Astronomical Society.

(b) Location map ofthefour profiles. Basin areas are shown in dark shading.

I

ofviscosity 1.2x10²¹ Pas overlain by a 75 km thick asthenosphere ofviscosity 2.0x 10¹⁹ Pas (both channel and deep flow). They emphasise the need for a low viscosity asthenosphere to explain the present uplift data. There are some trade-offs between asthenosphere viscosity and thickness and mantle viscosity and ice thickness that were not specifically addressed.

They noted that changes ofthe lithospheric rigidity (within the range of 1-1 OOx 1 0²³ Nm) result in only minorchanges in modelling the presentrate ofuplift. In another study, Peltier (1989) applyied viscoelastic theory for glacial isostatic adjustment to model Free Air gravity and geoid anomalies over Fennoscandia and Laurentia consistent in general with observations using viscosities of 10²¹ and 4.5x 10²¹ Pas, respectively, for the upper and lower mantle.

Lambeck et al. (1990) obtained an upper mantle viscosity of3-5x 1 0²⁰ Pas and lower mantle viscosity of 2-7x 1 0²¹ Pas from inversion of the observations of the postglacial sea level changes in northwestern Europe. Thus there is a general agreement about the order of magnitude of viscosity in the Earth' s mantle as a whole, a figure that has changed little since the c1assical study of Haskell (1935). However, new information and ideas continue to appear and further observations and integrated modelling are still needed. We consider these observations, and their geodynamic significance, in Chapter 7.2.