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Late Miocene uplift in the Romagnan Apennines and the detachment of subducted lithosphere
Tectonophysics 315 (1999) 319–335 www.elsevier.com/locate/tecto
Late Miocene uplift in the Romagnan Apennines and the detachment of subducted lithosphere M.J. van der Meulen a, *, T.J. Kouwenhoven b, G.J. van der Zwaan b,c, J.E. Meulenkamp a,b, M.J.R. Wortel a a Utrecht University, Faculty of Earth Sciences, Vening Meinesz Research School of Geodynamics (VMSG), PO Box 80021, 3508 TA Utrecht, The Netherlands b Utrecht University, Faculty of Earth Sciences, Institute for Paleoenvironments and Paleoclimate Utrecht (IPPU), PO Box 80021, 3508 TA Utrecht, The Netherlands c University of Nijmegen, Faculty of Science, Department of Ecology/Biogeology, Toernooiveld 1, 6525 ED Nijmegen, The Netherlands
Abstract We report part of a test of the hypothesis that detachment of subducted lithosphere may be a process of lateral propagation of a horizontal tear [ Wortel and Spakman, Proc. Kon. Ned. Akad. Wetensch., 95 (1992) 325–347]. We have used the Apennines as a test area. The test procedure consists of the comparison of hypothetical vertical motions, predicted from the expected redistribution of slab pull forces, with observed vertical motions. We demonstrate that a Late Miocene depocentre migration from the Northern towards the Central Apennines is associated with uplift of (the fore-arc of ) the Northern Apennines. Such a combination of a depocentre shift and uplift is thought to be diagnostic for lateral migration of slab detachment. The depocentre migration was identified in earlier work [van der Meulen et al., Earth Planet. Sci. Lett., 154 (1998) 203–219]. This contribution focuses on uplift, which has primarily been identified through the geohistory analysis of the Monte del Casino Section (Romagnan Apennines, Northern Italy). Owing to methodological problems, the start and duration of the uplift phase could not be constrained, and only a minimum estimate of the total amount of uplift (483±180 m) is obtained. The data do allow for an estimate of the uplift rate: 163±61 cm/ky. A review of regional data results in better constraints on the timing of the above lateral reorganisation of the fore-arc, and on the spatial extent of the uplifted area. Depocentre development in the Central Apennines began between 8.6 and 8.3 Ma B.P. Uplift started between 9 and 8 Ma B.P., and affected the entire northernmost Apennines. © 1999 Elsevier Science B.V. All rights reserved. Keywords: Apennines; basin analysis; foredeep; Italy; Late Miocene; Messinian; slab detachment; Southern Europe; stratigraphy; Tortonian
* Corresponding author. Present address: Ministry of Transport, Public Works and Water Management, Road and Hydraulic Engineering Division (Raw Materials Section), PO Box 5044, 2600 GA Delft, The Netherlands. Fax: +31-30-2535030. E-mail addresses: [email protected] (M.J. van der Meulen), [email protected] (T.J. Kouwenhoven), [email protected] (G.J. van der Zwaan), [email protected] (J.E. Meulenkamp), [email protected] (M.J.R. Wortel ) 0040-1951/99/$ - see front matter © 1999 Elsevier Science B.V. All rights reserved. PII: S0 0 4 0- 1 9 51 ( 9 9 ) 0 02 8 2 -6
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1. Introduction The seismic velocity structure of the Central and Western Mediterranean shows evidence for slab detachment underneath the Dinarid–Hellenic and Apenninic Arcs (Spakman, 1991; see also Spakman et al., 1993). On the basis of this result, Wortel and Spakman (1992) formulated the hypothesis of lateral migration of slab detachment. This hypothesis offers two verifiable predictions with respect to surface effects of slab detachment (see Fig. 1). The redistribution of slab pull forces associated with the process of lateral migration of slab detachment is taken (1) to introduce a lateral
(i.e. along-arc) component to foredeep depocentre migrations, and (2) to cause dynamical rebound in depocentral areas after depocentre shifting. Van der Meulen et al. (1998) showed that a lateral component became superimposed on the internal–external migration of Apenninic foredeep depocentres from the Tortonian onwards, which is in agreement with the first of the predictions. Fig. 2 shows the position of Northern Apenninic Late Oligocene to Late Miocene foredeep depocentres (D1–D5). The first step in the mentioned lateral migration of foredeep depocentres (from D3 to D4) was established after a major reorganisation of the Northern and Central Apenninic foreland
Fig. 1. Graphical representation of the hypothesis of lateral migration of slab detachment ( Wortel and Spakman, 1992), and predicted surface effects. Reprinted from van der Meulen et al. (1998) with kind permission of Elsevier Science.
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Fig. 2. The onset of lateral migration of foredeep depocentres of the Northern Apennines in the Tortonian (modified from van der Meulen et al., 1998).
during a Tortonian deformation phase [see also Centamore et al. (1978, 1979)]. This deformation phase resulted in the compartmentation of the southern portion of the Miocene Northern Apenninic foredeep, the Marnoso Arenacea Basin (i.e. south of the Early to Middle Miocene foredeep depocentre D3), into a suite of smaller basins arranged along the Central Apenninic front, the so-called ‘Minor Basins’ sensu Centamore et al. (1978, 1979). One of these basins, the Lazian Tagliacozzo Basin, accommodated the Tortonian depocentre D4 [see van der Meulen et al. (1998)]. This paper addresses the second prediction, i.e. rebound. We have reconstructed Late Miocene vertical motions of the Apenninic foredeep (Marnoso Arenacea Basin), at the approximate along-arc position of the foredeep depocentre D3 (see Fig. 2). Implicitly, this study aims to validate the conclusion of van der Meulen et al. (1998), that the lateral depocentre shift may have been caused by lateral migration of slab detachment, by means of an independent test. This is done by verifying whether the vertical motions show the predicted rebound (as in Fig. 1, zone C ). Tortonian regressive sequences
from the Northern Apenninic foredeep (Ricci Lucchi, 1986; Patacca et al., 1990), as well as from piggyback basins (Ricci Lucchi, 1986), provide qualitative evidence for uplift of the Northern Apenninic foreland (van der Meulen et al., 1998). Palaeobathymetrical studies are, however, needed to quantify the uplift, and in order to be able to discriminate between tectonically induced regression and the effects of the onset of the Messinian salinity crisis. For this purpose, vertical motions have been reconstructed for the Monte del Casino section. Fig. 3 consists of a geological map showing the location of the Monte del Casino section, and a profile through the area. The general structure of the area is a monocline dipping towards the north– northeast. The Ligurian allochthon, in the western part of the map, is a gravitationally displaced unit that progressively invaded the Marnoso Arenacea Basin during the latest Early Miocene to Late Miocene (De Jager, 1979; Ricci Lucchi, 1986). The vertical motions of the Ligurian allochthon, as well as those of the aforementioned ‘Minor Basins’, will be reviewed in a discussion of the results for our study area in a regional context.
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Fig. 3. Simplified geological map showing the position of the Monte del Casino section, modified from Abbate et al. (1982). The profile is modified from Bruni (1973).
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2. An age model for the Monte del Casino section In order to reconstruct vertical motions, both the age and estimated palaeodepth of sampled levels have been determined. The Monte del Casino section shows a pronounced cyclic alternation of grey homogeneous marly clays and brown, organic-rich sapropelitic layers. In their turn, these cycles are grouped into clusters. This pattern is interpreted to have been astronomically induced ( Vai et al., 1993; Krijgsman et al., 1997) and, consequently, the age model for the section has potentially a 20 ky resolution, i.e. the periodicity of the precession cycle. A full cyclostratigraphic framework for the Late Miocene is given by Hilgen et al. (1995) (see also Krijgsman, 1996). The error in astronomical ages is difficult to asses. If the tuning of sedimentary cycles to the precession curves is correct, this error, arising from uncertainties in the astronomical solution, is estimated to be approximately 0.1% within the time span considered in this study [see Krijgsman et al. (1999) for a discussion]. A mistuning of n cycles obviously results in an error of n times 20 ky (i.e. 0.2 to 0.4% of the astronomical age of Upper Miocene cycles). Cycle mistuning, however, can be avoided because correlations are not only made on the level of individual cycles/beds, but also on that of cycle/bed clusters (see below). Fig. 4 shows bed-to-bed correlations of the sapropels of the Monte del Casino section to the astronomically dated sapropels of the Metochia section on the island of Gavdos, south of Crete (Hilgen et al., 1995). Individual sapropels have been numbered (C1–C56 for the Monte del Casino section, M1–M96 for the Metochia section). Sapropel clusters have been numbered II through IX, following Hilgen et al. (1995). In Fig. 4 only fully developed sapropels are shown; for more stratigraphic detail, the reader is referred to Fig. 5. The Monte del Casino section as presented by Krijgsman et al. (1997), who focused on the Tortonian–Messinian boundary, comprised of sapropels C15 to C42. We have added 14 cycles above (C43–C56) and 13 (C1–C13) below Krijgsman’s column. The pattern displayed by sapropels C3– C13, which can be subdivided in groups of three (C3–C5) and eight (C6–C13), perfectly matches
Fig. 4. Sapropel correlations of the Monte del Casino section with the Metochia section on the Gavdos Island, south of Greece (Hilgen et al., 1995; Krijgsman, 1996). For stratigraphic details of the Monte del Casino section: see Fig. 5.
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Fig. 5. Stratigraphic details and age model for the Monte del Casino section.
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the pattern of sapropels M58–M68 (cluster VI ) of the Metochia section. Sapropels C1 and C2 most probably represent the top of Metochia cluster V. The age of the sapropel C1 is estimated at 8.070 Ma. There is some uncertainty, because the lower part of our section is not entirely exposed. Sapropels C43–C56 cannot be dated astronomically, because at present there is no sound cyclostratigraphic framework available for the Messinian. Therefore, we inferred an age for these sapropels assuming that the duration of each cycle is 21.7 ky, the average periodicity of precession. This approach (see Krijgsman et al., 1994) yields an age of 6.307 Ma for the uppermost sapropel (C56). The resulting age model for the Monte del Casino section is shown in Fig. 5. Analysed samples (see Fig. 5 for stratigraphic position) have been dated assuming a constant sedimentation rate within each sedimentary cycle. According to the proposed age model, the time span recorded by the Monte del Casino section is 1.763 Ma.
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availability of nutrients throughout the water column is a function of both depth and of primary production [see Suess (1980) and Berger and Diester-Haas (1988)], the planktonic/epibenthic foraminiferal ratio is depth dependent when the availability of nutrients is a limiting factor for benthic life. Laminated, sapropelitic layers are deposited under anoxic or strongly dysoxic conditions. The relation between depth and planktonic/benthic foraminiferal ratios, as expressed in Eq. (1), does not hold when oxygen instead of the availability of nutrients becomes a limiting factor for benthic life. Therefore, sapropelitic beds are excluded from bathymetrical analysis. Allochthonous faunal elements, introduced in an association by re-sedimentation, may hamper palaeobathymetrical analysis. This problem can partly be avoided by screening samples for benthic species of which the estimated depth range clearly does not agree with the estimated palaeodepth. 3.2. Results: geohistory of the Monte del Casino section
3. Vertical motions 3.1. Methods: reconstructing palaeobathymetry Palaeobathymetrical estimates have been calculated from the ratio between planktonic and epibenthic (bottom dwelling) foraminifera. The share of epibenthic foraminifera in a total association is inversely correlated to depth (van der Zwaan et al., 1990). The relation between depth and the proportion of planktonic foraminifera can be expressed as: D=e[3.58718+(0.03534 · PP)]
(1)
where D is depositional depth (metres), and PP= [(P/P+B)×100] is the percentage of planktonic foraminifera, disregarding inbenthic species (mainly Bolivina, Bulimina and Uvigerina species) in total fauna. The implicit assumptions are that the number of planktonic foraminifera is proportional to the primary production, and that the number of epibenthic foraminifera is proportional to the availability of nutrients on the sea bottom. As the
Fig. 6 shows the results for the Monte del Casino section. The geohistory diagram [see van Hinte (1978) for more details concerning geohistory analysis] shows the estimated palaeodepths (triangles), and the vertical position of the reference level ( lowermost sampled level, circles), obtained by adding the amount of accumulated sediment to the palaeodepth. The subsidence/uplift curve displayed in the geohistory diagram is not corrected for sediment loading, compaction, or eustasy. When calculating the amount of uplift and the uplift velocity, appropriate corrections will, however, be applied (see below). A eustatic curve (modified from Haq et al., 1988) has been included in Fig. 6, to allow for the comparison of the order of magnitude of shallowing/deepening trends and estimated eustatic movements in the time span examined. The vertical scatter in the geohistory diagram is considerable. Therefore, we fitted the bathymetrical and subsidence/uplift curves using a threepoint moving average smoothing function. Four time slices and corresponding data populations
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can be distinguished (see Fig. 6). From 8.070 to 6.965 Ma B.P. (time slice A) a high average percentage of planktonic foraminifera (97%) yields palaeodepths of approximately 1100 m. Between 6.965 and 6.669 Ma B.P. (time slice B) the average palaeodepth decreases. The vertical scatter, however, increases, so no clear trend towards more shallow depths is shown by the data. The next time slice (C ), from 6.669 to 6.365 Ma B.P., is characterised by palaeodepths varying around 600 m. Again, the data scatter is considerable: the lowest value is 262 m and the highest value is 1010 m. Finally (time slice D, 6.365–6.294 Ma B.P.), the calculated palaeodepth increases again. Sedimentation rates between 8.070 and 7.973 Ma B.P. rapidly decrease from 124 to 20 cm/ky. Between 7.973 and 7.575 Ma B.P. there is a further, slower decrease from 20 to 8 cm/ky. After 7.575 Ma B.P. the sedimentation rates vary around 3 cm/ky.
4. Discussion of the vertical motions 4.1. Palaeobathymetry and the onset of the Messinian salinity crisis Although an uplift phase can be inferred from the geohistory diagram ( Fig. 6), it cannot straightforwardly be quantified. There are two main reasons for this. First, the palaeodepth inferred for time slice A is close to the lower detection limit for P/B-based depth estimates: 100% plankton yields a minimum depth of 1238 m. Bathymetrical variation below this depth is not recorded, so, in this case, the onset of uplift cannot be reconstructed. Secondly, although the analytical error of depth estimates decreases when depth decreases (van der Zwaan et al., 1990), the scatter of bathymetrical data clearly increases with decreasing depth (see Fig. 6). This suggests that P/B ratios are not
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exclusively determined by depth. In fact, the lithology of the upper part of the Monte del Casino section shows a general progressive deterioration of bottom water conditions, superimposed on the astronomically induced variation in bottom water conditions as expressed by the sedimentary cyclicity. Between approximately 123 m (C18) and the top of the section (7.167–6.294 Ma) the sedimentary cycles gradually change from duplets dominated by homogeneous clays, to duplets dominated by sapropelitic sediments (see Fig. 5). At ±138 m (C44) the homogeneous intervals become greenish grey and fetid. This suggests an increasing organic carbon content and, hence, a decreasing ventilation. Kouwenhoven et al. (1999) show that, starting at 7.16 Ma B.P., an open marine benthic foraminiferal assemblage was progressively replaced by a suite of increasingly stressed assemblages reflecting a gradually decreasing bottom water ventilation and an increasing salinity, prior to the Messinian salinity crisis. Nevertheless, we feel confident that, although the P/B ratios seem to be affected by stressed environmental conditions, this does not seriously hamper palaeodepth reconstructions. Moderately dysoxic conditions, as observed for the homogeneous intervals in the upper part of the section, do not affect the depth dependency of P/B ratios, because they mainly influence the size of the inbenthic association (van der Zwaan et al., 1990), and infaunal elements have been disregarded when calculating palaeodepths (see Section 3.1). Only for time slice D might the inferred deepening be an artifact. The corresponding stratigraphic interval is almost continuously laminated, the homogeneous intervals are thin and lithologically almost indistinguishable from the sapropels. This suggests persistent strongly dysoxic or anoxic conditions. The shallowing inferred from the P/B ratios can also be observed qualitatively from the benthic record [for a detailed account of the encountered species see Kouwenhoven et al. (1999)]. Among
Fig. 6. Geohistory and burial history diagrams for the Monte del Casino section. Four time slices are distinguished: (A) bathymetry below detection limits of P/B-based (i.e. plankton/benthos ratio) palaeodepth estimates; (B) shallowing/uplift; (C ) shallow; (D) apparent deepening due to the onset of the Messinian salinity crisis. A eustatic curve is shown to allow for a comparison of the order of magnitude of vertical and eustatic motions.
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the species present in the interval characterised by high P/B ratios are Cibicides wuellerstorfi, Siphonina reticulata, and Melonis spp., which are generally considered to have a deep (bathyal ) habitat [see e.g. see Jorissen (1988) and Murray (1991)]. In samples taken from the part of the sequence having lower P/B ratios, Astrononion spp., Bulimina subulata, Cibicides pseudoeugerianus, and Hanzawaia boueana have been encountered, of which the combination indicates a shelfal habitat [see e.g. Jorissen (1988) and Murray (1991)]. 4.2. A probabilistic estimate of the uplift rate The vertical scatter imposes a problem when attempting to infer the amount of uplift and the uplift rate directly from the geohistory diagram. Therefore, probabilistic estimates have been made of the amount of uplift and the net uplift rate within time slice B (see Fig. 7). The average bathymetry for time slice A is 1108±53 m. The average depth for time slice C is 607±172 m. The shallowing during time slice B consequently equals 501±180 m. The present thickness of sediment deposited during shallowing is 8 m. The burial depth of this
level was approximately 1.1 km (De Jager, 1979). Using the present thickness−initial thickness relations proposed by Perrier and Quiblier (1974), the restored thickness becomes 18 m. The net uplift is equal to the shallowing minus the amount of sediment deposited during shallowing, i.e. 483±180 m. The duration of time slice B is 296 ky, so the net uplift rate becomes 163±61 cm/ky. The calculated rate and (minimum) magnitude of the uplift exceed the estimated amplitude of eustatic movements by approximately three orders of magnitude (see Fig. 6). Also, the effect of sediment loading on the vertical motions can be neglected, since the sedimentation rates during the uplift phase were low (~6 cm/ky). A possible explanation for the low sedimentation rate will be offered in Section 5. 4.3. A comparison with uplift rates in similar geological settings Now that an uplift phase has been identified, a comparison of its rate with published vertical velocities in similar settings can be made. In two areas, uplift has been tentatively interpreted as rebound after slab detachment. Price and AudleyCharles (1987) present a scenario of events after
Fig. 7. The early Messinian uplift phase: a probabilistic approach. The data populations shown in the histogram are pertaining to time slice A (deep population) and time slice C (shallow population).
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an inferred early Late Pliocene (3.5 Ma B.P.) detachment of the Australian plate below the southern Banda arc. This includes an uplift in the collision zone, supposed to be caused by the elastic unflexing of the downbending plate after rupture. Estimates of Pliocene uplift velocities along the southern Banda arc vary between 50 and 500 cm/ky (Fortuin and De Smet, 1991; see also Audley-Charles, 1986). The uncertainty in the estimates is large and it remains unclear if all of the uplift can be interpreted as rebound, so a comparison with the Tortonian Northern Apenninic uplift rate seems unjustified. Westaway (1993) shows that Calabria and the Southern Basilicata have been uplifted with a rate of approximately 100±10 cm/ky after 0.7 Ma B.P. The uplift is interpreted to be rebound after slab detachment (see also Hippolyte et al., 1994), which is in fairly good accordance with the conclusions of van der Meulen et al. (1998). The vertical velocity has the same order of magnitude as that of northern Italy in the Miocene. Estimated uplift velocities for episutural basins along retreating boundaries that became involved in accretionary processes vary between 27 and 125 cm/ky for central Crete during the Early Pliocene (Meulenkamp et al., 1994) and between 500 and 1000 cm/ky for Timor during the Late Pleistocene ( Fortuin and De Smet, 1991, and references cited therein). The rates suggested for rebound after slab detachment fall completely within this range. In general, caution should be taken when comparing vertical velocities from different studies. In many cases, poorly constrained palaeodepth reconstructions introduce large uncertainties in estimates of the vertical velocities. Also, owing to the episodic nature of uplift in tectonically active areas (Fortuin and De Smet, 1991; Pirazzoli et al., 1993, 1994), the vertical velocities obtained are strongly dependent on age resolution.
5. Discussion of the results in a regional context 5.1. Late Miocene uplift Now that uplift has been reconstructed for an area where uplift is expected on the basis of a
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depocentre analysis (see above), i.e. that of the Miocene Northern Apenninic depocentre (D3 in Fig. 2), its spatial extent will be discussed. For this purpose, published data are reviewed of (1) the area northwest and west of the Monte del Casino area, occupied by the allochthonous Ligurian sheet, and (2) the area between the Monte del Casino area and the Tortonian depocentre, i.e. the southern Romagnan Apennines, the Montefeltro Colata, and the Umbro–Marchean ‘Minor Basins’ (see Figs. 2 and 8). The second objective of the following discussion is to reconsider the timing of depocentre development in the Central Apennines in the light of the outcomes of the study of Northern Apenninic uplift. 5.1.1. The external Ligurids Some general inferences on vertical motions of the external Ligurian sheet can be made on the basis of the stratigraphy of late Tortonian– Messinian epi-Ligurian sequences, i.e. successions deposited in piggyback basins on top of the Ligurian sheet (see Fig. 8). The relationships between foredeep and epi-Ligurian piggyback basin sequences have been discussed by Ricci Lucchi (1986). This author demonstrated that each of the Northern Apenninic foredeeps has a counterpart on top of the Ligurian sheet. The epiLigurian equivalent of the late Tortonian clays of the upper Romagnan Marnoso Arenacea Formation, is the Termina Formation. The sedimentary facies of the marly Termina Formation is similar to that of the upper Romagnan Marnoso Arenacea. Like its counterpart in the foredeep, the Termina Formation overlies a coarser-grained turbiditic unit [Bismantova Formation; see e.g. Amarosi et al. (1993) and Fioroni and Panini (1987)], and is overlain by evaporites at places where the record is complete. According to Ricci Lucchi (1986), the transition between 9 and 8 Ma B.P. from the Bismantova to the Termina Formation, as well as the similar transition from the Marnoso Arenacea turbidites to the upper Marnoso Arenacea clays in the Romagnan Apennines (see Figs. 5 and 6), is associated with uplift and erosion of the Ligurian sheet. The decreasing sedimentation rates and fining upward trend may seem in conflict with the inter-
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Fig. 8. Map of the northernmost Apennines, showing the position of the Ligurian sheet and its piggyback basins. The Alpine units on the left side of the figure mark the western boundary of the Northern Apennines [modified from Muttoni et al. (1998)].
preted uplift. However, Ricci Lucchi (1986) explained this by the trapping of sediment with Alpine provenance within the evolving Apenninic topography. Apart from that, sedimentation rates in the latest Miocene should be discussed in the wider context of the changing environments prior to the Messinian salinity crisis. A full assessment of the latter aspect, however, is beyond the scope of this paper. In our study area, the transition from coarseto fine-grained sediments was gradual, and completed between 8.1 and 8.0 Ma B.P. (see Fig. 6). Because of the gradual nature of the transition it is not possible to date the fining event more accurately than the age of 9–8 Ma proposed by Ricci Lucchi (1986). A shallowing trend within the open marine Termina marls, which is evident from the presence of either an erosive surface, or paralic Messinian evaporites (where the record is complete) on top of this unit, has hitherto only been interpreted in the context of the Messinian salinity crisis (e.g. Vai and Ricci Lucchi, 1981; Ricci Lucchi, 1986). Only in a high-resolution study, such as the present one, is it possible to distinguish a tectonically
induced shallowing from an evaporative sea-level drop. Such studies are not available for the Termina marls. This is partly due to the fact that the dating potential of the Monte del Casino section, because of its combination of excellent outcropping and low degree of deformation, is quite unique. However, some inferences on palaeotopography can be made on the basis of the distribution of syn- and post-evaporitic Messinian deposits. On top of the Ligurian sheet such sediments are largely absent (Ricci Lucchi, 1986). Only on top of the southeasternmost part of the unit, approximately between its eastern limit and Bologna (see Fig. 8), are evaporites present. From this observation it can be inferred that, during the Messinian, the Ligurian sheet west of Bologna was elevated above base level. Such a topography can also be inferred from the distribution of continental, post-evaporitic Messinian sediments, which are absent on top of the Ligurian sheet, thin and discontinuous in the area of the Monte del Casino section, but rapidly increasing in thickness towards the southeast. In summary, there is a NW–SE increase of the
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thickness of upper Tortonian–Messinian sediments that allows for the identification of the external Ligurian sheet as an area that was uplifted together with the Romagnan Apenninic foredeep from 9– 8 Ma onwards. It is most significant that during the late Tortonian and the Messinian no indications exist for active displacement of the Ligurian sheet (see e.g., Ricci Lucchi, 1986). Therefore, vertical motions are most probably not related to the deformation (thrusting) of this unit. 5.1.2. The internal Ligurids The internal Ligurids are found on the Tyrrhenian side of the northernmost Apennines (see Fig. 8). Fission-track data from the internal Ligurian sheet suggest that the uplift, which has resulted in its present elevation, started between 9 and 8 Ma B.P. (Balestrieri et al., 1996; see also Fig. 9). This corroborates fully with the inferences on the timing of uplift based on the analysis of epi-Ligurian sequences. The uplift was recognised as far west as along the western limit of the Northern Apennines (Ghibaudo et al., 1985) (see
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Fig. 8). Along the Tyrrhenian coast, the area around La Spezia (see Fig. 8) seems to be a transitional area between the internal Ligurids, which were uplifted from the Late Miocene onwards, and the Apuan Alps, which were uplifted from the Early Pliocene onwards (Abbate et al., 1994). As for the internal Ligurids, fission-track ages of about 8 Ma have been reconstructed, but the length distribution suggests a relatively long residence in the partial annealing zone, followed — as for the Apuan Alps — by uplift during the Pliocene. Contemporaneously with uplift of the northernmost Apennines, extensional tectonics resulted in the formation of the Tyrrhenian Basin and its currently onshore equivalent in Tuscany, south and southeast with respect to the uplifted area. 5.1.3. The foredeep between the study area and the Tortonian Central Apenninic depocentre The area between the Monte del Casino section in the northern Romagnan foredeep and the Tortonian foredeep depocentre (the Tagliacozzo
Fig. 9. Miocene to Recent vertical motions of the internal Ligurids [modified from Balestrieri et al. (1996)].
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Basin) includes: the southeastern Romagnan foredeep, the Montefeltro Colata, and the Umbro– Marchean ‘Minor Basins’ (see Fig. 2). Of these three areas, only the Montefeltro Colata requires some introduction. The Montefeltro Colata is a Ligurian unit, which, at present, is geographically separated from the main Ligurian sheet. Along the strike of the Apennines, it marks the transition between the Romagnan Apenninic foredeep, and the area that is occupied by the Umbro–Marchean ‘Minor Basins’. As for the epi-Ligurian sequences, no inferences on vertical motions can be made on the basis of published data on upper Tortonian–lower Messinian autochthonous sequences. Within the Romagnan foredeep, no clear lateral or vertical facial differentiation is present. Within the ‘Minor Basins’, the upper Tortonian has not been systematically distinguished, and the scattered outcrops of their sequences do not allow for a detailed palaeobathymetrical study. However, on top of the Montefeltro Colata, a submarine high during the Late Miocene, sedimentary sequences reveal a different pattern of vertical motions with respect to those of the Ligurian sheet (see De Feyter, 1991). The upper Tortonian–lower Messinian equivalents of the upper Marnoso Arenacea clays and the Termina Formation show a deepening
trend, which would exclude the Montefeltro Colata from the uplifted area. This would suggest that the southern Romagnan foredeep (i.e. between the Monte del Casino area and the Montefeltro Colata) marked the transition from an uplifted external arc to a subsiding external arc in the late Miocene (see Fig. 10). The distribution and thickness of Messinian syn- and post-evaporitic sediments corroborates with such a configuration. These deposits show a southeastward increasing thickness within the Romagnan Apennines. In the area of the Umbro– Marchean ‘Minor Basins’ these sediments have variable thicknesses (100–1000 m), but they are invariably present (e.g. Boccaletti et al., 1986) (see Fig. 10). 5.2. Late Miocene subsidence: the timing of Central Apenninic depocentre development The present study on Late Miocene uplift was performed in the context of a Tortonian lateral foredeep depocentre shift (van der Meulen et al., 1998). The combination of such a lateral reorganisation and rebound is thought to be diagnostic of slab detachment sensu Wortel and Spakman (1992). For the sake of a uniform approach in the analysis of the heterogeneous stratigraphic datasets
Fig. 10. Panel showing a subdivision of the Late Miocene Northern Apenninic fore-arc into an uplifted area, a transitional area and an area of maximum subsidence. LIG: (main) Ligurian sheet; RF: Romagnan foredeep; MC: Montefeltro Colata; UMMB: Umbro– Marchean ‘Minor Basins’; CAD: Central Apenninic foredeep depocentre. Vertical black arrows qualitatively indicate vertical motions.
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used for the depocentre analysis of van der Meulen et al. (1998), the Oligocene and Miocene stages were not subdivided. Now that more accurate constraints have been put on the timing and areal extent of the uplift, the timing of the Late Miocene lateral depocentre shift along the Northern Apennines has to be re-assessed. Cipollari and Cosentino (1995) attribute the fill of the Lazian ‘Minor Basins’, including the Tagliacozzo Basin [see van der Meulen et al. (1998)] to the Late Tortonian NN11a nannoplankton zone (8.6–7.2 Ma B.P.; Berggren et al., 1995). Micropalaeontological analysis of our own samples puts the base of this sequence in the M13a planktonic foraminiferal zone [10.9–8.3 Ma B.P.; (Berggren et al., 1995); ~N16 (Blow, 1969), ~Globorotalia acostaensis (=Neogloboquadrina acostaensis) zone (Iaccarino, 1985)], characterised by the presence of Neogloboquadrina acostaensis and the absence of Globigerinoides obliquus extremus marker species. Combining the two biostratigraphic datings gives an age of 8.6–8.3 Ma to the beginning of sedimentation in the Central Apenninic depocentre. This falls completely within the time range for the onset of uplift in the northernmost Apennines.
6. Conclusions In this paper we report part of a test of the hypothesis of lateral migration of slab detachment ( Wortel and Spakman, 1992), for which the Apennines are used as a test area. In earlier work (van der Meulen et al., 1998) we found a pattern of lateral shifts of foredeep depocentres to be in support of this hypothesis. In this contribution we have demonstrated that the onset of the lateral, southeastward depocentre shifts was associated with uplift of the northernmost Apennines. The combination of these two observations is in good agreement with the hypothesis under scrutiny. The Late Miocene uplift phase was initially identified through the analysis of the Monte del Casino section in the northern Romagnan Apennines. Although the age model for this section is very good, the uplift phase cannot be considered completely constrained. The onset of uplift has
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most probably not been detected, because the basin started its uplift having a depth of more than 1.1 km, and only when a basin becomes less deep than approximately 1.1 km, can vertical motions be reconstructed quantitatively from micropalaeontological evidence (P/B ratios). Furthermore, the uplift may have continued throughout the Messinian, but the effects of deteriorating bottom-water conditions, and of sea-level lowering associated with the Messinian salinity crisis, preclude reliable depth estimates. Although the amount of uplift remains unresolved in our primary study area, the uplift rate is reconstructed to have been 163±61 cm/ky. An overview of the Late Miocene vertical motions of the Northern Apennines, combined with the results for the study area, gives a full insight in a lateral reorganisation of the Apenninic arc. Between 9 and 8 Ma B.P., a pattern of differential motions along the Northern Apennines came into being. The northernmost Apennines, the realm occupied by the Ligurian sheet, and including the northern Romagnan foredeep, became an area of uplift. The Central Apenninic foredeep became an area of maximum subsidence between 8.6 and 8.3 Ma B.P. A transitional domain was present between these two areas. The Central Apenninic foredeep remained an area of ( local ) maximum subsidence during the late Tortonian to Early Pliocene [see van der Meulen et al. (1998)]. It cannot be determined exactly how long the uplift of the northernmost Apennines lasted. Uplift was recorded up to the middle Messinian. At the beginning of the Pliocene there is a new phase of pronounced subsidence in the Northern Apenninic foredeep [see van der Meulen et al. (1998)], which suggests that the uplift may have ceased in the course of the middle to late Messinian. The presence of Lower Pliocene sequences on top of the Ligurian sheet (Ricci Lucchi et al., 1981; Ricci Lucchi, 1986) corroborates with this inference. It cannot be determined whether or not a time lag exists between depocentre development and uplift, because of uncertainties in dating. Nevertheless, given the regional pattern, the uplift may be interpreted as rebound after slab detachment. This interpretation is corroborated by the
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observation that, during uplift, no displacement of the Ligurian sheet occurred. The uplift is, therefore, most probably not related to local deformation. It must be emphasised that our conclusions are primarily formulated in terms of consistency with the hypothesis of lateral migration of slab detachment. Hence, whereas the good agreement between predictions and observations is in strong support of this hypothesis, we do not intend to imply that no alternative mechanisms are possible. If alternative mechanisms would appear to explain our observations equally well, further tests will have to be developed to discriminate between them.
Acknowledgements F.J. Hilgen, E. Carminati, L.J. Lourens, and A.J. van der Meulen are acknowledged for discussion and suggestions. A. Okx and J. de Wit are thanked for sample picking. G.J. van’t Veld and G.C. Ittmann prepared the samples. We thank the reviewers F. Roure and A. Fortuin for helpful suggestions. This work was conducted at the Faculty of Earth Sciences, Utrecht University, under the programme of the Vening Meinesz Research School of Geodynamics ( VMSG). M.J. van der Meulen acknowledges financial support from the Geosciences Foundation (GOA, currently ALW ) of the Netherlands Organization for Scientific Research (NWO).
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M.J. van der Meulen et al. / Tectonophysics 315 (1999) 319–335 mentologia ed evoluzione tettonico–sedimentaria dei depositi margine sud-orientale del bacino terziaro lugurepiemontese (valli Borbera, Scriva e Lemme). Boll. Soc. Geol. Ital. 104, 349–397. Haq, B.U., Hardenbol, J., Vail, P.R., 1988. Mesozoic and Cenozoic chronostratigraphy and cycles of sea-level change. In: SEPM Spec. Publ. 42, 71–108. Hilgen, F.J., Krijgsman, W., Langereis, C.G., Lourens, L.J., Santarelli, A., Zachariasse, W.J., 1995. Extending the astronomical (polarity) timescale into the Miocene. Earth Planet. Sci. Lett. 136, 495–510. Hippolyte, J.C., Angelier, J., Roure, F., 1994. A major geodynamic change revealed by Quaternary stress patterns in the Southern Apennines (Italy). Tectonophysics 230, 199–210. Iaccarino, S., 1985. Mediterranean Miocene and Pliocene planktic foraminifera. In: Bolli, H.M., Saunders, J.B., Perch-Nielsen, K. ( Eds.), Plankton Stratigraphy. Cambridge University Press, Cambridge, UK, pp. 283–314. Jorissen, F.J., 1988. Benthic foraminifera from the Adriatic sea, principles of phenotypic variation. Utrecht Micropal. Bull. 37. Kouwenhoven, T.J., Seidenkrantz, M.S., Van der Zwaan, G.J., 1999. Deep-water changes: the near-synchronous disappearance of a group of benthic foraminifera from the late Miocene Mediterranean. Palaeogeogr. Palaeoclim. Palaeoecol. 152, 259–281. Krijgsman, W., 1996. Miocene magnetostratigraphy and cyclostratigraphy in the Mediterranean: extension of the astronomical polarity time scale. Ph.D. Thesis, Utrecht University, Geologica Ultraiectina, 141, 207 pp. Krijgsman, W., Hilgen, F.J., Langereis, C.G., Zachariasse, W.J., 1994. The age of the Tortonian/Messinian boundary. Earth Planet. Sci. Lett. 121, 533–547. Krijgsman, W., Hilgen, F.J., Negri, A., Wijbrans, J.R., Zachariasse, W.J., 1997. The Monte del Casino section (Northern Apennines, Italy): a potential Tortonian/Messinian boundary stratotype? Palaeogeogr. Palaeoclim. Palaeoecol. 133, 27–47. Krijgsman, W., Hilgen, F.J., Marabini, S., Vai, G.B., 1999. New paleomagnetic and cyclostratigraphic age constraints on the Messinian of the Northern Apennines ( Vena del Gesso Basin, Italy). Mem. Soc. Geol. Ital. 54, 25–33. Meulenkamp, J.E., van der Zwaan, G.J., van Wamel, W.A., 1994. On Late Miocene to Recent vertical motions in the Cretan segment of the Hellenic arc. Tectonophysics 234, 53–72. Murray, J.W., 1991. Ecology and Palaeoecology of Benthic Foraminifera. Longman, Harlow, UK. Patacca, E., Sartori, R., Scandone, P., 1990. Tyrrhenian Basin and Apenninic arcs: kinematic relations since Late Tortonian times. Mem. Soc. Geol. Ital. 45, 425–451. Perrier, R., Quiblier, J., 1974. Thickness changes in sedimentary
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Late Miocene uplift in the Romagnan Apennines and the detachment of subducted lithosphere M.J. van der Meulen a, *, T.J. Kouwenhoven b, G.J. van der Zwaan b,c, J.E. Meulenkamp a,b, M.J.R. Wortel a a Utrecht University, Faculty of Earth Sciences, Vening Meinesz Research School of Geodynamics (VMSG), PO Box 80021, 3508 TA Utrecht, The Netherlands b Utrecht University, Faculty of Earth Sciences, Institute for Paleoenvironments and Paleoclimate Utrecht (IPPU), PO Box 80021, 3508 TA Utrecht, The Netherlands c University of Nijmegen, Faculty of Science, Department of Ecology/Biogeology, Toernooiveld 1, 6525 ED Nijmegen, The Netherlands
Abstract We report part of a test of the hypothesis that detachment of subducted lithosphere may be a process of lateral propagation of a horizontal tear [ Wortel and Spakman, Proc. Kon. Ned. Akad. Wetensch., 95 (1992) 325–347]. We have used the Apennines as a test area. The test procedure consists of the comparison of hypothetical vertical motions, predicted from the expected redistribution of slab pull forces, with observed vertical motions. We demonstrate that a Late Miocene depocentre migration from the Northern towards the Central Apennines is associated with uplift of (the fore-arc of ) the Northern Apennines. Such a combination of a depocentre shift and uplift is thought to be diagnostic for lateral migration of slab detachment. The depocentre migration was identified in earlier work [van der Meulen et al., Earth Planet. Sci. Lett., 154 (1998) 203–219]. This contribution focuses on uplift, which has primarily been identified through the geohistory analysis of the Monte del Casino Section (Romagnan Apennines, Northern Italy). Owing to methodological problems, the start and duration of the uplift phase could not be constrained, and only a minimum estimate of the total amount of uplift (483±180 m) is obtained. The data do allow for an estimate of the uplift rate: 163±61 cm/ky. A review of regional data results in better constraints on the timing of the above lateral reorganisation of the fore-arc, and on the spatial extent of the uplifted area. Depocentre development in the Central Apennines began between 8.6 and 8.3 Ma B.P. Uplift started between 9 and 8 Ma B.P., and affected the entire northernmost Apennines. © 1999 Elsevier Science B.V. All rights reserved. Keywords: Apennines; basin analysis; foredeep; Italy; Late Miocene; Messinian; slab detachment; Southern Europe; stratigraphy; Tortonian
* Corresponding author. Present address: Ministry of Transport, Public Works and Water Management, Road and Hydraulic Engineering Division (Raw Materials Section), PO Box 5044, 2600 GA Delft, The Netherlands. Fax: +31-30-2535030. E-mail addresses: [email protected] (M.J. van der Meulen), [email protected] (T.J. Kouwenhoven), [email protected] (G.J. van der Zwaan), [email protected] (J.E. Meulenkamp), [email protected] (M.J.R. Wortel ) 0040-1951/99/$ - see front matter © 1999 Elsevier Science B.V. All rights reserved. PII: S0 0 4 0- 1 9 51 ( 9 9 ) 0 02 8 2 -6
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1. Introduction The seismic velocity structure of the Central and Western Mediterranean shows evidence for slab detachment underneath the Dinarid–Hellenic and Apenninic Arcs (Spakman, 1991; see also Spakman et al., 1993). On the basis of this result, Wortel and Spakman (1992) formulated the hypothesis of lateral migration of slab detachment. This hypothesis offers two verifiable predictions with respect to surface effects of slab detachment (see Fig. 1). The redistribution of slab pull forces associated with the process of lateral migration of slab detachment is taken (1) to introduce a lateral
(i.e. along-arc) component to foredeep depocentre migrations, and (2) to cause dynamical rebound in depocentral areas after depocentre shifting. Van der Meulen et al. (1998) showed that a lateral component became superimposed on the internal–external migration of Apenninic foredeep depocentres from the Tortonian onwards, which is in agreement with the first of the predictions. Fig. 2 shows the position of Northern Apenninic Late Oligocene to Late Miocene foredeep depocentres (D1–D5). The first step in the mentioned lateral migration of foredeep depocentres (from D3 to D4) was established after a major reorganisation of the Northern and Central Apenninic foreland
Fig. 1. Graphical representation of the hypothesis of lateral migration of slab detachment ( Wortel and Spakman, 1992), and predicted surface effects. Reprinted from van der Meulen et al. (1998) with kind permission of Elsevier Science.
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Fig. 2. The onset of lateral migration of foredeep depocentres of the Northern Apennines in the Tortonian (modified from van der Meulen et al., 1998).
during a Tortonian deformation phase [see also Centamore et al. (1978, 1979)]. This deformation phase resulted in the compartmentation of the southern portion of the Miocene Northern Apenninic foredeep, the Marnoso Arenacea Basin (i.e. south of the Early to Middle Miocene foredeep depocentre D3), into a suite of smaller basins arranged along the Central Apenninic front, the so-called ‘Minor Basins’ sensu Centamore et al. (1978, 1979). One of these basins, the Lazian Tagliacozzo Basin, accommodated the Tortonian depocentre D4 [see van der Meulen et al. (1998)]. This paper addresses the second prediction, i.e. rebound. We have reconstructed Late Miocene vertical motions of the Apenninic foredeep (Marnoso Arenacea Basin), at the approximate along-arc position of the foredeep depocentre D3 (see Fig. 2). Implicitly, this study aims to validate the conclusion of van der Meulen et al. (1998), that the lateral depocentre shift may have been caused by lateral migration of slab detachment, by means of an independent test. This is done by verifying whether the vertical motions show the predicted rebound (as in Fig. 1, zone C ). Tortonian regressive sequences
from the Northern Apenninic foredeep (Ricci Lucchi, 1986; Patacca et al., 1990), as well as from piggyback basins (Ricci Lucchi, 1986), provide qualitative evidence for uplift of the Northern Apenninic foreland (van der Meulen et al., 1998). Palaeobathymetrical studies are, however, needed to quantify the uplift, and in order to be able to discriminate between tectonically induced regression and the effects of the onset of the Messinian salinity crisis. For this purpose, vertical motions have been reconstructed for the Monte del Casino section. Fig. 3 consists of a geological map showing the location of the Monte del Casino section, and a profile through the area. The general structure of the area is a monocline dipping towards the north– northeast. The Ligurian allochthon, in the western part of the map, is a gravitationally displaced unit that progressively invaded the Marnoso Arenacea Basin during the latest Early Miocene to Late Miocene (De Jager, 1979; Ricci Lucchi, 1986). The vertical motions of the Ligurian allochthon, as well as those of the aforementioned ‘Minor Basins’, will be reviewed in a discussion of the results for our study area in a regional context.
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Fig. 3. Simplified geological map showing the position of the Monte del Casino section, modified from Abbate et al. (1982). The profile is modified from Bruni (1973).
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2. An age model for the Monte del Casino section In order to reconstruct vertical motions, both the age and estimated palaeodepth of sampled levels have been determined. The Monte del Casino section shows a pronounced cyclic alternation of grey homogeneous marly clays and brown, organic-rich sapropelitic layers. In their turn, these cycles are grouped into clusters. This pattern is interpreted to have been astronomically induced ( Vai et al., 1993; Krijgsman et al., 1997) and, consequently, the age model for the section has potentially a 20 ky resolution, i.e. the periodicity of the precession cycle. A full cyclostratigraphic framework for the Late Miocene is given by Hilgen et al. (1995) (see also Krijgsman, 1996). The error in astronomical ages is difficult to asses. If the tuning of sedimentary cycles to the precession curves is correct, this error, arising from uncertainties in the astronomical solution, is estimated to be approximately 0.1% within the time span considered in this study [see Krijgsman et al. (1999) for a discussion]. A mistuning of n cycles obviously results in an error of n times 20 ky (i.e. 0.2 to 0.4% of the astronomical age of Upper Miocene cycles). Cycle mistuning, however, can be avoided because correlations are not only made on the level of individual cycles/beds, but also on that of cycle/bed clusters (see below). Fig. 4 shows bed-to-bed correlations of the sapropels of the Monte del Casino section to the astronomically dated sapropels of the Metochia section on the island of Gavdos, south of Crete (Hilgen et al., 1995). Individual sapropels have been numbered (C1–C56 for the Monte del Casino section, M1–M96 for the Metochia section). Sapropel clusters have been numbered II through IX, following Hilgen et al. (1995). In Fig. 4 only fully developed sapropels are shown; for more stratigraphic detail, the reader is referred to Fig. 5. The Monte del Casino section as presented by Krijgsman et al. (1997), who focused on the Tortonian–Messinian boundary, comprised of sapropels C15 to C42. We have added 14 cycles above (C43–C56) and 13 (C1–C13) below Krijgsman’s column. The pattern displayed by sapropels C3– C13, which can be subdivided in groups of three (C3–C5) and eight (C6–C13), perfectly matches
Fig. 4. Sapropel correlations of the Monte del Casino section with the Metochia section on the Gavdos Island, south of Greece (Hilgen et al., 1995; Krijgsman, 1996). For stratigraphic details of the Monte del Casino section: see Fig. 5.
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Fig. 5. Stratigraphic details and age model for the Monte del Casino section.
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the pattern of sapropels M58–M68 (cluster VI ) of the Metochia section. Sapropels C1 and C2 most probably represent the top of Metochia cluster V. The age of the sapropel C1 is estimated at 8.070 Ma. There is some uncertainty, because the lower part of our section is not entirely exposed. Sapropels C43–C56 cannot be dated astronomically, because at present there is no sound cyclostratigraphic framework available for the Messinian. Therefore, we inferred an age for these sapropels assuming that the duration of each cycle is 21.7 ky, the average periodicity of precession. This approach (see Krijgsman et al., 1994) yields an age of 6.307 Ma for the uppermost sapropel (C56). The resulting age model for the Monte del Casino section is shown in Fig. 5. Analysed samples (see Fig. 5 for stratigraphic position) have been dated assuming a constant sedimentation rate within each sedimentary cycle. According to the proposed age model, the time span recorded by the Monte del Casino section is 1.763 Ma.
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availability of nutrients throughout the water column is a function of both depth and of primary production [see Suess (1980) and Berger and Diester-Haas (1988)], the planktonic/epibenthic foraminiferal ratio is depth dependent when the availability of nutrients is a limiting factor for benthic life. Laminated, sapropelitic layers are deposited under anoxic or strongly dysoxic conditions. The relation between depth and planktonic/benthic foraminiferal ratios, as expressed in Eq. (1), does not hold when oxygen instead of the availability of nutrients becomes a limiting factor for benthic life. Therefore, sapropelitic beds are excluded from bathymetrical analysis. Allochthonous faunal elements, introduced in an association by re-sedimentation, may hamper palaeobathymetrical analysis. This problem can partly be avoided by screening samples for benthic species of which the estimated depth range clearly does not agree with the estimated palaeodepth. 3.2. Results: geohistory of the Monte del Casino section
3. Vertical motions 3.1. Methods: reconstructing palaeobathymetry Palaeobathymetrical estimates have been calculated from the ratio between planktonic and epibenthic (bottom dwelling) foraminifera. The share of epibenthic foraminifera in a total association is inversely correlated to depth (van der Zwaan et al., 1990). The relation between depth and the proportion of planktonic foraminifera can be expressed as: D=e[3.58718+(0.03534 · PP)]
(1)
where D is depositional depth (metres), and PP= [(P/P+B)×100] is the percentage of planktonic foraminifera, disregarding inbenthic species (mainly Bolivina, Bulimina and Uvigerina species) in total fauna. The implicit assumptions are that the number of planktonic foraminifera is proportional to the primary production, and that the number of epibenthic foraminifera is proportional to the availability of nutrients on the sea bottom. As the
Fig. 6 shows the results for the Monte del Casino section. The geohistory diagram [see van Hinte (1978) for more details concerning geohistory analysis] shows the estimated palaeodepths (triangles), and the vertical position of the reference level ( lowermost sampled level, circles), obtained by adding the amount of accumulated sediment to the palaeodepth. The subsidence/uplift curve displayed in the geohistory diagram is not corrected for sediment loading, compaction, or eustasy. When calculating the amount of uplift and the uplift velocity, appropriate corrections will, however, be applied (see below). A eustatic curve (modified from Haq et al., 1988) has been included in Fig. 6, to allow for the comparison of the order of magnitude of shallowing/deepening trends and estimated eustatic movements in the time span examined. The vertical scatter in the geohistory diagram is considerable. Therefore, we fitted the bathymetrical and subsidence/uplift curves using a threepoint moving average smoothing function. Four time slices and corresponding data populations
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can be distinguished (see Fig. 6). From 8.070 to 6.965 Ma B.P. (time slice A) a high average percentage of planktonic foraminifera (97%) yields palaeodepths of approximately 1100 m. Between 6.965 and 6.669 Ma B.P. (time slice B) the average palaeodepth decreases. The vertical scatter, however, increases, so no clear trend towards more shallow depths is shown by the data. The next time slice (C ), from 6.669 to 6.365 Ma B.P., is characterised by palaeodepths varying around 600 m. Again, the data scatter is considerable: the lowest value is 262 m and the highest value is 1010 m. Finally (time slice D, 6.365–6.294 Ma B.P.), the calculated palaeodepth increases again. Sedimentation rates between 8.070 and 7.973 Ma B.P. rapidly decrease from 124 to 20 cm/ky. Between 7.973 and 7.575 Ma B.P. there is a further, slower decrease from 20 to 8 cm/ky. After 7.575 Ma B.P. the sedimentation rates vary around 3 cm/ky.
4. Discussion of the vertical motions 4.1. Palaeobathymetry and the onset of the Messinian salinity crisis Although an uplift phase can be inferred from the geohistory diagram ( Fig. 6), it cannot straightforwardly be quantified. There are two main reasons for this. First, the palaeodepth inferred for time slice A is close to the lower detection limit for P/B-based depth estimates: 100% plankton yields a minimum depth of 1238 m. Bathymetrical variation below this depth is not recorded, so, in this case, the onset of uplift cannot be reconstructed. Secondly, although the analytical error of depth estimates decreases when depth decreases (van der Zwaan et al., 1990), the scatter of bathymetrical data clearly increases with decreasing depth (see Fig. 6). This suggests that P/B ratios are not
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exclusively determined by depth. In fact, the lithology of the upper part of the Monte del Casino section shows a general progressive deterioration of bottom water conditions, superimposed on the astronomically induced variation in bottom water conditions as expressed by the sedimentary cyclicity. Between approximately 123 m (C18) and the top of the section (7.167–6.294 Ma) the sedimentary cycles gradually change from duplets dominated by homogeneous clays, to duplets dominated by sapropelitic sediments (see Fig. 5). At ±138 m (C44) the homogeneous intervals become greenish grey and fetid. This suggests an increasing organic carbon content and, hence, a decreasing ventilation. Kouwenhoven et al. (1999) show that, starting at 7.16 Ma B.P., an open marine benthic foraminiferal assemblage was progressively replaced by a suite of increasingly stressed assemblages reflecting a gradually decreasing bottom water ventilation and an increasing salinity, prior to the Messinian salinity crisis. Nevertheless, we feel confident that, although the P/B ratios seem to be affected by stressed environmental conditions, this does not seriously hamper palaeodepth reconstructions. Moderately dysoxic conditions, as observed for the homogeneous intervals in the upper part of the section, do not affect the depth dependency of P/B ratios, because they mainly influence the size of the inbenthic association (van der Zwaan et al., 1990), and infaunal elements have been disregarded when calculating palaeodepths (see Section 3.1). Only for time slice D might the inferred deepening be an artifact. The corresponding stratigraphic interval is almost continuously laminated, the homogeneous intervals are thin and lithologically almost indistinguishable from the sapropels. This suggests persistent strongly dysoxic or anoxic conditions. The shallowing inferred from the P/B ratios can also be observed qualitatively from the benthic record [for a detailed account of the encountered species see Kouwenhoven et al. (1999)]. Among
Fig. 6. Geohistory and burial history diagrams for the Monte del Casino section. Four time slices are distinguished: (A) bathymetry below detection limits of P/B-based (i.e. plankton/benthos ratio) palaeodepth estimates; (B) shallowing/uplift; (C ) shallow; (D) apparent deepening due to the onset of the Messinian salinity crisis. A eustatic curve is shown to allow for a comparison of the order of magnitude of vertical and eustatic motions.
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the species present in the interval characterised by high P/B ratios are Cibicides wuellerstorfi, Siphonina reticulata, and Melonis spp., which are generally considered to have a deep (bathyal ) habitat [see e.g. see Jorissen (1988) and Murray (1991)]. In samples taken from the part of the sequence having lower P/B ratios, Astrononion spp., Bulimina subulata, Cibicides pseudoeugerianus, and Hanzawaia boueana have been encountered, of which the combination indicates a shelfal habitat [see e.g. Jorissen (1988) and Murray (1991)]. 4.2. A probabilistic estimate of the uplift rate The vertical scatter imposes a problem when attempting to infer the amount of uplift and the uplift rate directly from the geohistory diagram. Therefore, probabilistic estimates have been made of the amount of uplift and the net uplift rate within time slice B (see Fig. 7). The average bathymetry for time slice A is 1108±53 m. The average depth for time slice C is 607±172 m. The shallowing during time slice B consequently equals 501±180 m. The present thickness of sediment deposited during shallowing is 8 m. The burial depth of this
level was approximately 1.1 km (De Jager, 1979). Using the present thickness−initial thickness relations proposed by Perrier and Quiblier (1974), the restored thickness becomes 18 m. The net uplift is equal to the shallowing minus the amount of sediment deposited during shallowing, i.e. 483±180 m. The duration of time slice B is 296 ky, so the net uplift rate becomes 163±61 cm/ky. The calculated rate and (minimum) magnitude of the uplift exceed the estimated amplitude of eustatic movements by approximately three orders of magnitude (see Fig. 6). Also, the effect of sediment loading on the vertical motions can be neglected, since the sedimentation rates during the uplift phase were low (~6 cm/ky). A possible explanation for the low sedimentation rate will be offered in Section 5. 4.3. A comparison with uplift rates in similar geological settings Now that an uplift phase has been identified, a comparison of its rate with published vertical velocities in similar settings can be made. In two areas, uplift has been tentatively interpreted as rebound after slab detachment. Price and AudleyCharles (1987) present a scenario of events after
Fig. 7. The early Messinian uplift phase: a probabilistic approach. The data populations shown in the histogram are pertaining to time slice A (deep population) and time slice C (shallow population).
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an inferred early Late Pliocene (3.5 Ma B.P.) detachment of the Australian plate below the southern Banda arc. This includes an uplift in the collision zone, supposed to be caused by the elastic unflexing of the downbending plate after rupture. Estimates of Pliocene uplift velocities along the southern Banda arc vary between 50 and 500 cm/ky (Fortuin and De Smet, 1991; see also Audley-Charles, 1986). The uncertainty in the estimates is large and it remains unclear if all of the uplift can be interpreted as rebound, so a comparison with the Tortonian Northern Apenninic uplift rate seems unjustified. Westaway (1993) shows that Calabria and the Southern Basilicata have been uplifted with a rate of approximately 100±10 cm/ky after 0.7 Ma B.P. The uplift is interpreted to be rebound after slab detachment (see also Hippolyte et al., 1994), which is in fairly good accordance with the conclusions of van der Meulen et al. (1998). The vertical velocity has the same order of magnitude as that of northern Italy in the Miocene. Estimated uplift velocities for episutural basins along retreating boundaries that became involved in accretionary processes vary between 27 and 125 cm/ky for central Crete during the Early Pliocene (Meulenkamp et al., 1994) and between 500 and 1000 cm/ky for Timor during the Late Pleistocene ( Fortuin and De Smet, 1991, and references cited therein). The rates suggested for rebound after slab detachment fall completely within this range. In general, caution should be taken when comparing vertical velocities from different studies. In many cases, poorly constrained palaeodepth reconstructions introduce large uncertainties in estimates of the vertical velocities. Also, owing to the episodic nature of uplift in tectonically active areas (Fortuin and De Smet, 1991; Pirazzoli et al., 1993, 1994), the vertical velocities obtained are strongly dependent on age resolution.
5. Discussion of the results in a regional context 5.1. Late Miocene uplift Now that uplift has been reconstructed for an area where uplift is expected on the basis of a
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depocentre analysis (see above), i.e. that of the Miocene Northern Apenninic depocentre (D3 in Fig. 2), its spatial extent will be discussed. For this purpose, published data are reviewed of (1) the area northwest and west of the Monte del Casino area, occupied by the allochthonous Ligurian sheet, and (2) the area between the Monte del Casino area and the Tortonian depocentre, i.e. the southern Romagnan Apennines, the Montefeltro Colata, and the Umbro–Marchean ‘Minor Basins’ (see Figs. 2 and 8). The second objective of the following discussion is to reconsider the timing of depocentre development in the Central Apennines in the light of the outcomes of the study of Northern Apenninic uplift. 5.1.1. The external Ligurids Some general inferences on vertical motions of the external Ligurian sheet can be made on the basis of the stratigraphy of late Tortonian– Messinian epi-Ligurian sequences, i.e. successions deposited in piggyback basins on top of the Ligurian sheet (see Fig. 8). The relationships between foredeep and epi-Ligurian piggyback basin sequences have been discussed by Ricci Lucchi (1986). This author demonstrated that each of the Northern Apenninic foredeeps has a counterpart on top of the Ligurian sheet. The epiLigurian equivalent of the late Tortonian clays of the upper Romagnan Marnoso Arenacea Formation, is the Termina Formation. The sedimentary facies of the marly Termina Formation is similar to that of the upper Romagnan Marnoso Arenacea. Like its counterpart in the foredeep, the Termina Formation overlies a coarser-grained turbiditic unit [Bismantova Formation; see e.g. Amarosi et al. (1993) and Fioroni and Panini (1987)], and is overlain by evaporites at places where the record is complete. According to Ricci Lucchi (1986), the transition between 9 and 8 Ma B.P. from the Bismantova to the Termina Formation, as well as the similar transition from the Marnoso Arenacea turbidites to the upper Marnoso Arenacea clays in the Romagnan Apennines (see Figs. 5 and 6), is associated with uplift and erosion of the Ligurian sheet. The decreasing sedimentation rates and fining upward trend may seem in conflict with the inter-
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Fig. 8. Map of the northernmost Apennines, showing the position of the Ligurian sheet and its piggyback basins. The Alpine units on the left side of the figure mark the western boundary of the Northern Apennines [modified from Muttoni et al. (1998)].
preted uplift. However, Ricci Lucchi (1986) explained this by the trapping of sediment with Alpine provenance within the evolving Apenninic topography. Apart from that, sedimentation rates in the latest Miocene should be discussed in the wider context of the changing environments prior to the Messinian salinity crisis. A full assessment of the latter aspect, however, is beyond the scope of this paper. In our study area, the transition from coarseto fine-grained sediments was gradual, and completed between 8.1 and 8.0 Ma B.P. (see Fig. 6). Because of the gradual nature of the transition it is not possible to date the fining event more accurately than the age of 9–8 Ma proposed by Ricci Lucchi (1986). A shallowing trend within the open marine Termina marls, which is evident from the presence of either an erosive surface, or paralic Messinian evaporites (where the record is complete) on top of this unit, has hitherto only been interpreted in the context of the Messinian salinity crisis (e.g. Vai and Ricci Lucchi, 1981; Ricci Lucchi, 1986). Only in a high-resolution study, such as the present one, is it possible to distinguish a tectonically
induced shallowing from an evaporative sea-level drop. Such studies are not available for the Termina marls. This is partly due to the fact that the dating potential of the Monte del Casino section, because of its combination of excellent outcropping and low degree of deformation, is quite unique. However, some inferences on palaeotopography can be made on the basis of the distribution of syn- and post-evaporitic Messinian deposits. On top of the Ligurian sheet such sediments are largely absent (Ricci Lucchi, 1986). Only on top of the southeasternmost part of the unit, approximately between its eastern limit and Bologna (see Fig. 8), are evaporites present. From this observation it can be inferred that, during the Messinian, the Ligurian sheet west of Bologna was elevated above base level. Such a topography can also be inferred from the distribution of continental, post-evaporitic Messinian sediments, which are absent on top of the Ligurian sheet, thin and discontinuous in the area of the Monte del Casino section, but rapidly increasing in thickness towards the southeast. In summary, there is a NW–SE increase of the
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thickness of upper Tortonian–Messinian sediments that allows for the identification of the external Ligurian sheet as an area that was uplifted together with the Romagnan Apenninic foredeep from 9– 8 Ma onwards. It is most significant that during the late Tortonian and the Messinian no indications exist for active displacement of the Ligurian sheet (see e.g., Ricci Lucchi, 1986). Therefore, vertical motions are most probably not related to the deformation (thrusting) of this unit. 5.1.2. The internal Ligurids The internal Ligurids are found on the Tyrrhenian side of the northernmost Apennines (see Fig. 8). Fission-track data from the internal Ligurian sheet suggest that the uplift, which has resulted in its present elevation, started between 9 and 8 Ma B.P. (Balestrieri et al., 1996; see also Fig. 9). This corroborates fully with the inferences on the timing of uplift based on the analysis of epi-Ligurian sequences. The uplift was recognised as far west as along the western limit of the Northern Apennines (Ghibaudo et al., 1985) (see
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Fig. 8). Along the Tyrrhenian coast, the area around La Spezia (see Fig. 8) seems to be a transitional area between the internal Ligurids, which were uplifted from the Late Miocene onwards, and the Apuan Alps, which were uplifted from the Early Pliocene onwards (Abbate et al., 1994). As for the internal Ligurids, fission-track ages of about 8 Ma have been reconstructed, but the length distribution suggests a relatively long residence in the partial annealing zone, followed — as for the Apuan Alps — by uplift during the Pliocene. Contemporaneously with uplift of the northernmost Apennines, extensional tectonics resulted in the formation of the Tyrrhenian Basin and its currently onshore equivalent in Tuscany, south and southeast with respect to the uplifted area. 5.1.3. The foredeep between the study area and the Tortonian Central Apenninic depocentre The area between the Monte del Casino section in the northern Romagnan foredeep and the Tortonian foredeep depocentre (the Tagliacozzo
Fig. 9. Miocene to Recent vertical motions of the internal Ligurids [modified from Balestrieri et al. (1996)].
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Basin) includes: the southeastern Romagnan foredeep, the Montefeltro Colata, and the Umbro– Marchean ‘Minor Basins’ (see Fig. 2). Of these three areas, only the Montefeltro Colata requires some introduction. The Montefeltro Colata is a Ligurian unit, which, at present, is geographically separated from the main Ligurian sheet. Along the strike of the Apennines, it marks the transition between the Romagnan Apenninic foredeep, and the area that is occupied by the Umbro–Marchean ‘Minor Basins’. As for the epi-Ligurian sequences, no inferences on vertical motions can be made on the basis of published data on upper Tortonian–lower Messinian autochthonous sequences. Within the Romagnan foredeep, no clear lateral or vertical facial differentiation is present. Within the ‘Minor Basins’, the upper Tortonian has not been systematically distinguished, and the scattered outcrops of their sequences do not allow for a detailed palaeobathymetrical study. However, on top of the Montefeltro Colata, a submarine high during the Late Miocene, sedimentary sequences reveal a different pattern of vertical motions with respect to those of the Ligurian sheet (see De Feyter, 1991). The upper Tortonian–lower Messinian equivalents of the upper Marnoso Arenacea clays and the Termina Formation show a deepening
trend, which would exclude the Montefeltro Colata from the uplifted area. This would suggest that the southern Romagnan foredeep (i.e. between the Monte del Casino area and the Montefeltro Colata) marked the transition from an uplifted external arc to a subsiding external arc in the late Miocene (see Fig. 10). The distribution and thickness of Messinian syn- and post-evaporitic sediments corroborates with such a configuration. These deposits show a southeastward increasing thickness within the Romagnan Apennines. In the area of the Umbro– Marchean ‘Minor Basins’ these sediments have variable thicknesses (100–1000 m), but they are invariably present (e.g. Boccaletti et al., 1986) (see Fig. 10). 5.2. Late Miocene subsidence: the timing of Central Apenninic depocentre development The present study on Late Miocene uplift was performed in the context of a Tortonian lateral foredeep depocentre shift (van der Meulen et al., 1998). The combination of such a lateral reorganisation and rebound is thought to be diagnostic of slab detachment sensu Wortel and Spakman (1992). For the sake of a uniform approach in the analysis of the heterogeneous stratigraphic datasets
Fig. 10. Panel showing a subdivision of the Late Miocene Northern Apenninic fore-arc into an uplifted area, a transitional area and an area of maximum subsidence. LIG: (main) Ligurian sheet; RF: Romagnan foredeep; MC: Montefeltro Colata; UMMB: Umbro– Marchean ‘Minor Basins’; CAD: Central Apenninic foredeep depocentre. Vertical black arrows qualitatively indicate vertical motions.
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used for the depocentre analysis of van der Meulen et al. (1998), the Oligocene and Miocene stages were not subdivided. Now that more accurate constraints have been put on the timing and areal extent of the uplift, the timing of the Late Miocene lateral depocentre shift along the Northern Apennines has to be re-assessed. Cipollari and Cosentino (1995) attribute the fill of the Lazian ‘Minor Basins’, including the Tagliacozzo Basin [see van der Meulen et al. (1998)] to the Late Tortonian NN11a nannoplankton zone (8.6–7.2 Ma B.P.; Berggren et al., 1995). Micropalaeontological analysis of our own samples puts the base of this sequence in the M13a planktonic foraminiferal zone [10.9–8.3 Ma B.P.; (Berggren et al., 1995); ~N16 (Blow, 1969), ~Globorotalia acostaensis (=Neogloboquadrina acostaensis) zone (Iaccarino, 1985)], characterised by the presence of Neogloboquadrina acostaensis and the absence of Globigerinoides obliquus extremus marker species. Combining the two biostratigraphic datings gives an age of 8.6–8.3 Ma to the beginning of sedimentation in the Central Apenninic depocentre. This falls completely within the time range for the onset of uplift in the northernmost Apennines.
6. Conclusions In this paper we report part of a test of the hypothesis of lateral migration of slab detachment ( Wortel and Spakman, 1992), for which the Apennines are used as a test area. In earlier work (van der Meulen et al., 1998) we found a pattern of lateral shifts of foredeep depocentres to be in support of this hypothesis. In this contribution we have demonstrated that the onset of the lateral, southeastward depocentre shifts was associated with uplift of the northernmost Apennines. The combination of these two observations is in good agreement with the hypothesis under scrutiny. The Late Miocene uplift phase was initially identified through the analysis of the Monte del Casino section in the northern Romagnan Apennines. Although the age model for this section is very good, the uplift phase cannot be considered completely constrained. The onset of uplift has
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most probably not been detected, because the basin started its uplift having a depth of more than 1.1 km, and only when a basin becomes less deep than approximately 1.1 km, can vertical motions be reconstructed quantitatively from micropalaeontological evidence (P/B ratios). Furthermore, the uplift may have continued throughout the Messinian, but the effects of deteriorating bottom-water conditions, and of sea-level lowering associated with the Messinian salinity crisis, preclude reliable depth estimates. Although the amount of uplift remains unresolved in our primary study area, the uplift rate is reconstructed to have been 163±61 cm/ky. An overview of the Late Miocene vertical motions of the Northern Apennines, combined with the results for the study area, gives a full insight in a lateral reorganisation of the Apenninic arc. Between 9 and 8 Ma B.P., a pattern of differential motions along the Northern Apennines came into being. The northernmost Apennines, the realm occupied by the Ligurian sheet, and including the northern Romagnan foredeep, became an area of uplift. The Central Apenninic foredeep became an area of maximum subsidence between 8.6 and 8.3 Ma B.P. A transitional domain was present between these two areas. The Central Apenninic foredeep remained an area of ( local ) maximum subsidence during the late Tortonian to Early Pliocene [see van der Meulen et al. (1998)]. It cannot be determined exactly how long the uplift of the northernmost Apennines lasted. Uplift was recorded up to the middle Messinian. At the beginning of the Pliocene there is a new phase of pronounced subsidence in the Northern Apenninic foredeep [see van der Meulen et al. (1998)], which suggests that the uplift may have ceased in the course of the middle to late Messinian. The presence of Lower Pliocene sequences on top of the Ligurian sheet (Ricci Lucchi et al., 1981; Ricci Lucchi, 1986) corroborates with this inference. It cannot be determined whether or not a time lag exists between depocentre development and uplift, because of uncertainties in dating. Nevertheless, given the regional pattern, the uplift may be interpreted as rebound after slab detachment. This interpretation is corroborated by the
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observation that, during uplift, no displacement of the Ligurian sheet occurred. The uplift is, therefore, most probably not related to local deformation. It must be emphasised that our conclusions are primarily formulated in terms of consistency with the hypothesis of lateral migration of slab detachment. Hence, whereas the good agreement between predictions and observations is in strong support of this hypothesis, we do not intend to imply that no alternative mechanisms are possible. If alternative mechanisms would appear to explain our observations equally well, further tests will have to be developed to discriminate between them.
Acknowledgements F.J. Hilgen, E. Carminati, L.J. Lourens, and A.J. van der Meulen are acknowledged for discussion and suggestions. A. Okx and J. de Wit are thanked for sample picking. G.J. van’t Veld and G.C. Ittmann prepared the samples. We thank the reviewers F. Roure and A. Fortuin for helpful suggestions. This work was conducted at the Faculty of Earth Sciences, Utrecht University, under the programme of the Vening Meinesz Research School of Geodynamics ( VMSG). M.J. van der Meulen acknowledges financial support from the Geosciences Foundation (GOA, currently ALW ) of the Netherlands Organization for Scientific Research (NWO).
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