Assessing tectonic subsidence from estimates of Holocene relative sea-level change: an example from the NW Mediterranean (Magra Plain, Italy)
Alessandro Chelli1, Marta Pappalardo2, Monica Bini2, Helmut Brückner3, Giorgio Neri4, Michele Neri4 and Giorgio Spada5
1 Dipartimento di Fisica e Scienze della Terra "M. Melloni", Parma University, Italy 2 Dipartimento di Scienze della Terra, Pisa University, Italy
3 Institute of Geography, University of Cologne, Germany 4 Ambiter s.r.l., Parma, Italy
5 Dipartimento di Scienze Pure e Applicate, Urbino University “Carlo Bo”, Italy and Istituto Nazionale di Geofisica e Vulcanologia (INGV), Bologna
Corresponding author:
Marta Pappalardo, Dipartimento di Scienze della Terra, Pisa University, Via S. Maria 53, 56126 Pisa, Italy
Email:[email protected]
Abstract
New sedimentological sea-level indicators are presented from the River Magra coastal plain, in NW Italy. Chronologically well constrained paralic peats and organic sediments which had been
deposited in a defined relationship with sea level, were recovered in four of the seven boreholes considered in this work. Their evolution scatters in the time span of the last 6,000 years. Since the cores are located within a single sedimentary basin, it was possible to correct the elevation of marker horizons for the effect of sediment compaction by means of both a field and a geotechnical method. Thus, seven reliable index points for the mid-late Holocene sea-level rise were obtained. The age/depth model derived from them was compared to that of sea-level predictions from two different Glacial Isostatic Adjustment (GIA) models available for the area. In both cases the modelled sea-level estimates overlie the index points, suggesting lower relative sea-level elevation than the one predicted considering the combined eustatic and hydro-isostatic components. Based on 1 2 3 4 5 7 8 9 10 11 12 13 14 15 16 17 18 19 20 21 22 23 24 25 26 27 28 29
the general tectonic setting of the area, this finding was interpreted as the effect of a tectonic subsidence of the basin, the rate of which can be quantified on average as 0.5 mm/yr since the mid-Holocene, with a sharp increase after 2,500 years BP. By providing a reliable estimate of the rate of tectonic subsidence in a coastal area of NW Italy, this research contributes to refining the
geodynamic model of this part of the Mediterranean basin.
Keywords: relative sea-level rise; tectonic subsidence, mid-late Holocene; sea-level index point; Mediterranean Sea; NW Italy
Introduction
Relative sea-level (RSL) estimates are available from many sites along the coastline of the Mediterranean Sea ( Lambeck and Bard, 2000; Lambeck et al., 2004; Lambeck and Purcell, 2005; Lambeck et al., 2011; Stocchi et al., 2005, Stocchi and Spada, 2007), in particular for the last 6,000 years. In some tracts of the basin intensive studies where performed, like in Greece (Pavlopoulos et al., 2007, 2013; Vött, 2007; Pirazzoli, 2005; Brückner et al., 2010; Vacchi et al., 2014; Ghilardi et al., 2014), Israel (Sivan et al., 2004), Lebanon (Morhange et al., 2006a), large parts of peninsular and insular Italy (Lambeck et al., 2004, 2011; Morhange, et al., 2006b; Antonioli et al., 2009; Furlani et al., 2011; Primavera et al., 2011; Spampinato et al., 2011; Amato et al., 2013) and the Gulf of Lion (Vella and Provansal, 2000; Morhange et al., 2001). For other areas data are scattered (e.g., Malta; Furlani et al., 2013); this is especially true for the southern side of the basin, like Tunisia (Morhange and Pirazzoli, 2005; Anzidei et al., 2011) and Lybia (Anzidei et al., 2011). Due to the geodynamic complexity of the Mediterranean region (Faccenna et al., 2014), Glacial Isostatic Adjustment (GIA) models alone can be unable to fully explain the relative sea-level observations along the coastlines, since they do not account for the tectonic component of deformation. In fact, the co-existence of near-coast mountain chains with different uplift rates, interspaced subsiding coastal plains, and scattered crustal displacement due to recent tectonic activity, constitute important sources of regional deformation that are superimposed to the long-wavelength effects of GIA (see e.g., Lambeck and Purcell, 2005; Antonioli et al., 2009). In 30 31 32 33 34 35 36 37 38 39 40 41 42 43 44 45 46 47 48 49 50 51 52 53 54 55 56 57
addition, GIA models usually do not include complexities such as the rheological variations suggested by geological evidence (see e.g., Milne and Mitrovica, 2008), and only seldom consider deglaciation history of the peri-mediterranean mountain chains (Stocchi et al., 2005). Accounting for these effects can be particularly difficult for a region of relatively limited extent like the Mediterranean region, since they demand advanced and high-resolution GIA modelling tools.
It is, therefore, important to improve and refine the amount of data available on Holocene shoreline displacements (Rovere et al., 2012) in order to provide new evidence of recent crustal dynamics and further constraints for isostatic models. Furthermore, there is still a certain amount of uncertainty about the eustatic sea-level rise along the coastline of the Mediterranean during the latest part of the Holocene; for example about the sea-level estimate during Roman times (Evelpidou et al., 2012; Morozova, 2013). This is determined by the lack of agreement on the true meaning and the mode of employment of certain sea-level indicators , and on the adopted models (Vacchi et al., 2014), and may create incorrect evaluation on the assessment of future sea-level changes and, indeed, with meaningful consequence on the hazard assessment and coastal management. This uncertainty mighty be, at least partially, overcome if the number of sea-level indictors is implemented in correspondence of the wide coastal areas with a scarce number of data as well as of those where data are completely lacking.
As stated by Pirazzoli (1991), peat from paralic swamps is considered one of the most reliable sea-level indicators because, among geologic indicators, coastal peat records provide a very good vertical resolution (Törnqvist and Hijma, 2012) especially where the tidal range is low, like along the Mediterranean coasts characterized by microtidal environments; therefore, they have been extensively used to reconstruct Holocene relative sea-level changes (e.g. Allen, 2000; Shennan and Horton, 2002; Engelhart and Horton, 2012; Marra et al., 2013; Giosan et al., 2006; Vacchi et al., 2014 ). Where peat layers are not available, wood fragments and plant remains can be used as sea-level indicators (Vött, 2007), provided that their primarary position with respect to palaeo sea sea-level can be inferred from the sedimentary facies of the layer they belong to (Table I). The vertical resolution of sedimentological sea-level indicators can be improved if contemporary analogues of the used deposits may be investigated in the studied area. Especially for peat, information on the 58 59 60 61 62 63 64 65 66 67 68 69 70 71 72 73 74 75 76 77 78 79 80 81 82 83 84 85
relationship between tide levels and contemporary vegetation, sediment and biostratigraphy helps to estimate the rate and direction of changes inferred from sediment sequences (Shennan, 1989).
Within the Mediterranean basin sedimentological sea-level indicators for Holocene sea-level reconstructions have been used by a number of papers. Apart from those mentioned in the
introduction, that we took as our main reference some case studies are provided by Marriner et al., 2006; Mauz and Bungenstock, 2007; Bernasconi and Stanley, 2011; Bini et al., 2012; Amato et al., 2013; Romano et al., 2013; Seeliger et al., 2013; Ghilardi et al., 2014. In all cases index points from sedimentological indicators proved to be good quality sea- level markers, compared to others ones that are often more difficult to relate to their contemporary sea-level elevation (e.g. archaeological markers) or more difficult to date (e.g. tidal notches). Furthermore, many sedimentary sequences in Mediterranean coastal plains are suitable for the reconstruction of Holocene RSL curves (e.g. Ghilardi et al., 2012; Amorosi et al., 2013a,b).
Subsidence is considered one of the factors controlling the formation of coastal plains, especially in connection of a delta (Morgan, 1970). Most authors that study the relationship between subsidence and delta formation differentiate tectonic subsidence from that dependant on other natural or anthropogenic factors. The former is due to the structural behaviour of the bedrock on top of which a delta systems develops. Overall subsidence, though, is driven also by factors such as sediment compaction, groundwater and natural gas over-pumping, dredging of canals,(see e.g. Törnqvist et al., 2008).
This paper addresses new sedimentological sea level indicators (SLIs) retrieved along the easternmost coast of Liguria region in NW Italy (Figure 1a). In the review on Holocene shorelines in Italy (Antonioli et al., 2009) sea-level index points are missing for the study area, except for the archaeological sea-level indicator, the seaward outlet of a sewage channel, on stable rocky coasts described by Chelli et al. (2005). Nevertheless, Liguria represents a rather interesting area for the study of Holocene shorelines, due to its extremely low uplift rate (lower than 0.2 mm/yr), at least since the last interglacial stage (Federici and Pappalardo, 2006; Rovere et al., 2011), but probably even for the last 300 ka (Federici and Pappalardo, 2001; Biagioni et al., 2007). As a result geomorphological sea-level markers are not preserved owing to the morphodynamic processes 86 87 88 89 90 91 92 93 94 95 96 97 98 99 100 101 102 103 104 105 106 107 108 109 110 111 112 113
typical of this region. By contrast, sedimentological index points for the Holocene may be present (Carobene et al., 2008; Bini et al., 2009a,b, 2010, 2012). In particular at the easternmost edge of the region, the River Magra coastal plain (Fig. 1) provides an excellent sedimentary coastal archive, which has been widely investigated in the framework of a geoarchaeological project focussed on the Roman city of Luna (Bini et al., 2014).
In our study area sedimentological sea-level indicators, if correctly interpreted, provide reliable index points of mid- to late-Holocene sea-level changes. In fact, the coastal area is characterized by extremely small tidal ranges, a Holocene stratigraphy represented mostly by minerogenic sediments, and small residual marshes, which are modern analogues of similar Holocene features of the coastal plain stratigraphic sequence. These circumstances are favourable for identifying with sufficient accuracy sedimentological sea-level indicators. Thus, there is the chance for unravelling the tectonic component in the sea-level equation in order to investigate the crustal mobility in the area.
Materials and methods
In the framework of the Luna geoarchaeological project (Bini et al., 2014) seven new sedimentological sea-level indicators (SLIs) were collected from several cores drilled in the River Magra coastal plain in NW Italy. In this area the city of Luna was built by the Romans in 177 BC. The settlement provided anchorage facilities until the first centuries AD when its harbour underwent an unstoppable process of siltation (Gervasini, 2007). The cores LUNI4, 5, 16 and 18 were collected with a hand-operated mechanical corer (Cobra mk1 of Atlas Copco Co.), ORTO06 with a drilling machine. Three of these cores have already been described in previous works: ORTO06 in Bini et al. (2006; 2012); LUNI4 and 5 in Bini et al. 2009a. LUNI16 and 18 are unpublished. Sedimentological, palaeontological (micro- and macrofaunal) and geochemical analyses were carried out on the retrieved material in order to reconstruct the succession of sedimentary facies through time. By integrating the additional cores BH6 and 38 it was possible to generate a NE-SW cross-section through the lagoon deposit at its maximum extension. These logs are derived from a database of cores (Bini et al., 2006) built-up on cores retrieved in the past for 114 115 116 117 118 119 120 121 122 123 124 125 126 127 128 129 130 131 132 133 134 135 136 137 138 139 140 141
technical purposes (e.g., construction of roads and buildings). They are also interpreted in terms of sedimentary facies, highlighting the major environmental changes that had occurred in the Holocene evolution of the basin. Unfortunately, the cores BH6, BH38 and LUNI18 lack age constraints. Nevertheless, since they belong to the same basin, it was possible to correlate the changes of deposits (facies).
The altitude of the ground surface at drilling points was determined using a differential GPS “Leica GS09”, referenced to the Italian Ordnance Datum. From this the elevation of SLIs of the cores was inferred; these elevations were corrected for the effect of sediment compaction.
The relevance of compaction in Holocene coastal sedimentary sequences has been extensively recognized in coastal wetlands, where peat is the main sediment component (Morozova and Smith, 2003; Long et al., 2006; Edwards, 2006; Van Asselen, 2011; Engelhart and Horton, 2012), but also in minerogenic, mainly organic-rich sediments from coastal alluvial plains and deltas (Törnqvist et al., 2008; van Asselen et al., 2009; Woodroffe, 2009; Brain et al., 2011; 2015; Horton et al., 2013).
Concerning the assessment of sediment compaction as contribution to the total subsidence, a double approach has been proposed in the scientific literature (Tovey and Paul, 2002; van Asselen, 2011; Marra et al., 2013): (i) the field approach based on the presence of dated peat samples on top of the incompressible basement, and (ii) a geotechnical method which is aimed at decompacting the core sediments. They were both applied in our work, and the results were cross-checked for validation. Finally, a compaction rate for the sedimentary layers on top of the incompressible basement was assessed for each core, and the lowering of each SLI was computed with respect to its true height at the time of deposition.
Seven radiocarbon ages (Table II), carried out on peat and organic matter in order to overcome reservoir effect corrections, provided the necessary chronologic framework. Conventional ages were calibrated with the CALIB5 software and the calibration curves of Reimer et al. (2009). The 14C ages are reported with a double standard deviation.
The palaeo-level band, i.e. the interval in altitude limiting the position of the palaeo sea-level pointed out by an indicator (e.g., peat, wood and plant remains), was stated (Table III), for 142 143 144 145 146 147 148 149 150 151 152 153 154 155 156 157 158 159 160 161 162 163 164 165 166 167 168 169
each kind of indicator employed. Its amplitude was assessed referring to the review of Vött (2007) on the relative sea-level changes in NW Greece, following the methodology outlined in some of the IGCP projects (starting fom IGCP 61, Shennan, 1989) and applied in different types of coastal sedimentary environments. Further information about palaeo-sea-level bands applied to peats in a Mediterranean context was derived from Vella and Provansal (2000; Table I, p. 31). In the Rhône delta (Gulf of Fos, west of Marseilles, France), these authors applied a sea-level band of about 70-75 cm around the depths at which the peat levels were found, based on observations of the modern relationship between peat development and sea level in the same area. In our work the vertical resolution of the sedimentological sea-level indicators was checked against the relationship of modern analogues, i.e. the mean elevation of the bottom of present-day small coastal peat bogs ,far about 2 km from the coastline, has been measured using the GPS “Leica GS09” and resulted 0.02 m asl.
Two different GIA models were employed to compare our observational data. From the GIA model used for the coast of Italy by Lambeck et al. (2011), based on the global ice model K33, the curve for the Versilia Plain was extracted. Moreover a specific envelope was worked out for the Magra Plain based on the GIA model ICE-5G (VM2) (Peltier, 2004; Spada and Stocchi, 2007) and run through the SELEN program for solving the sea-level equation (Spada and Melini, 2013). The envelope is obtained using as input data a 90 km thickness of the lithosphere and a mantle viscosity ranging from 2.4 to 3.1021 Pa.s (lower mantle) and from 0.4 to 0.6.1021 Pa.s (upper mantle).
General setting of the area
The investigated coastal plain along the NW coast of Italy is at the mouth of River Magra (Fig. 1), which discharges into the Ligurian Sea, one of the sub-basins of the Mediterranean Sea. The area is characterised by a microtidal environment with a tidal range of ca. 40 cm and a maximum storm surge of ca. 30 cm (data from the Italian National Tide Gauge Network, tide gauge: La Spezia data time interval: 2009–2012)a. The plain is developed at the foothills of the so called “Apuan Alps”, a sector of the Northern Apennines (Argnani et al., 2003; Molli, 2008). The inner portion of the mountain chain is characterized by an extensional tectonic regime that caused the development of 170 171 172 173 174 175 176 177 178 179 180 181 182 183 184 185 186 187 188 189 190 191 192 193 194 195 196 197
intermontane basins (grabens) characterized by NW-SE striking normal faults and intervening horsts since the Late Miocene. The lower Magra Valley (Figure 1a) represents a part of one of these graben systems: the Magra - Vara graben system (Bernini et al., 1996; Bartolini, 2003), developed since the Middle Miocene in response to the opening of the Provençal and of the Tyrrhenian Sea (Molli, 2008 and references therein); the Magra Plain represents its seaward termination (Figure 1a). The age and distribution of the deposits filling-up the graben, roughly accounts for its development from SW to NE and from S to N (Bernini, 1991; Bernini et al., 1996; Bartolini, 2003). Federici (1973) accounts for the tilting (40°–50° towards SSW) of the Pliocene deposits recognized in the underground lignite mines. Based on observations carried out on the alluvial terraces bounding the valley, this author infers that the movement continued throughout the Pleistocene. Borghini et al. (2000) suggest that faults bordering the NE side of the graben would have been active during the Pleistocene and even the Holocene. The crustal movements revealed by a campaign of GPS measurements across the northern Apennines belt conducted between 1996 and 2010 confirm that in the western side of the Northern Apennines, where the strain rate field is well resolved, the deformation field is still active and is dominated by extensional strain (Bennet et al., 2012). Diffuse, small magnitude earthquakes (M<6.5) are frequent in the area (Tertulliani and Maramai, 1998), but the largest ones seldom exceed magnitude M5.5. An earthquake is reported to have probably occurred around the end of 4th and the beginning of 5th century AD (Ward Perkins, 1978; Rossignani, 1989; Fazzini and Maffei, 2000; Durante, 2001), which is considered among the natural causes of the decay of the Roman city of Luna (Durante, 2001). Both aseismic and coseismic displacement along fault planes are possible mechanisms which may be considered responsible for the subsidence of the plain.
The Magra graben was filled in the Pliocene and early Pleistocene, firstly by clay-rich lacustrine and fluvial deposits (Federici, 1973), then by middle-upper Pleistocene gravel-rich alluvial fans. The Luna project investigated the top of the Holocene valley fill in the seaward portion of the plain, where sediments accounting for marine to paralic environments dominate the topmost 10-15 m of the depositional sequence; these deposits started to develop after the peak of the Holocene transgression (Bini et al., 2009a; 2012).
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The continental shelf in front of the analysed coastal tract corresponds to the northern rim of the Viareggio basin, a half-graben characterized by NNW-SSE and NS dip-slip normal faults and NE-SW prevailing strike-slip faults or fault belts. Even if the Viareggio basin has been characterized by a subsiding regime since the upper Miocene (Fanucci, 1978), since the mid Pleistocene its tectonic inactivity was accounted for (Carboni et al., 2010).
Towards SE the Magra Plain merges into the wider Versilia Plain, which was considered as the type site of the mid-Holocene transgression (so-called “Versilian”, Blanc, 1942; Federici, 1993; Antonioli et al., 2000). The geoarchaeological survey of the lower Magra Plain (Bini et al., 2012 and references therein) revealed the existence and the evolution of a wide wetland, ca. 1 km in diameter, locally called “Seccagna” (Figure 1b), that started as a lagoon and evolved throughout the mid- and upper Holocene overlapping the distal portion of one of the alluvial fans which borders the foothills of the Apuan Alps (the Parmignola Stream alluvial fan, Figure 1b). The Seccagna lagoon was separated from the open sea by a sandbar created at the mouth of the River Magra, and it became more and more closed due to the progradation of the plain. Its sedimentary sequence testifies the innermost penetration of the Holocene transgression and the following regressive phase, chronologically constrained thanks to a minimum age (5840±50 14C years BP) provided by the oldest basal lagoonal sediments which directly overlie the alluvial fan; this facies was retrieved in core LUNI5 at a depth of 7.30 m bsl. (Bini et al., 2012). Unfortunately, for the Magra Plain, the basal littoral sediments are still undated. The lagoon in the Seccagna basin disappeared around 2500 years BP (Bini et al., 2012), possibly in connection with the first Neoglacial cooling episode, when a widespread development of swamps occurred in the area. The complete siltation of the basin took place after the Middle Ages. All the SLIs used in this work were retrieved from the Seccagna deposit.
Results
The geologic cross-section across the Magra Plain (Figure 2) enables us to reconstruct the depositional sequence of the former Seccagna waterbody.
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Even if the main environmental transitions are recognizable in all the cores, the BH6 core, the deepest one used in this work, illustrates the entire depositional sequence over the sandy gravel deposit, likely the top of the Parmignola Stream alluvial fan. The sequence starts at about 28 m bsl with a 6 m thick silty clay, rich in organic material, attributable to a lagoon deposit. It is followed by an about 12 m thick deposit of sand to silty sand with gravel lenses and shell fragments, probably representing a sand bar in the littoral environment. It is overlapped by a second lagoonal environment, as testified by the clay to silty clay deposit, rich in organic matter. The subsequent evolution of the basin is represented by its progressive siltation and the transformation into a swamp, progressively filled-in by a floodplain deposit, and lastly by land reclamation activities that started already during Roman times. In Table IV the main features of the sedimentary facies
retrieved in the basin sedimentary sequence are described, upon which their interpretation was based.
At its NE margin, the Seccagna lagoon (LUNI5, Figure 2) directly developed on top of the Parmignola stream alluvial fan. In the most peripheral part (LUNI4 and ORTO06, Figure 2), the lagoonal facies is absent; the waterbody deposit is only represented by a swamp deposit out of clayey silt or fine-sand with organic matter, both diffused and concentrated in thin layers. The top of the Parmignola Stream alluvial fan deepens from NE to SW as shown in the BH6 core (Figure 2). Here the Seccagna lagoon deposit lies on a sandy littoral deposit, containing also pebbles and shell fragments. Similar littoral deposits were found in the LUNI18 and BH38 cores, beneath the lagoon and swamp deposits, respectively. The depositional sequence was chronologically constrained by radiocarbon dating carried out on botanical remains (Table III).
The features of the seven sedimentological SLIs are reported in Table III. In two cases (LUNI5/21 and ORTO06) the marker is represented by basal peat, which provides a very good accuracy in sea-level estimation (Behre, 2004; van Asselen, 2010). In all other cases the marker was represented by partly decomposed plant or wood remains, wherefore a broader palaeo-sea-level band was assigned.
Vertical uncertainties in the paleo-sea-level value assessment were evaluated. The limits of palaeo-sea-level bands for the SLIs employed in this work are presented in Table II, according to 253 254 255 256 257 258 259 260 261 262 263 264 265 266 267 268 269 270 271 272 273 274 275 276 277 278 279 280
literature on the subject. Every dated sample was also evaluated against its sedimentological and geomorphological background, considering also the relationship between former sea levels and the elevation of coastal water bodies, as revealed by the facies of the deposits and the geomorphological setting of the study area. The total vertical error was calculated as the sum of the uncertainties due to the environmental/sedimentological context and the tidal range (cf. Pavlopoulos et al., 2012).
Sediment compaction was assumed to be negligible for the alluvial fan deposits, represented by sandy gravels that are found at the bottom of LUNI4, ORTO06, LUNI5 and BH6 cores. In fact, the alluvial fan may be considered as incompressible as compared to the overlaying alluvial-lagoonal-littoral sequence. The displacement for the SLIs collected from the LUNI16 core was calculated assuming that the stratigraphic sequence below the bottom of the core continues quite similar to that of BH6, which the reconstructed geological model seems to highlight.
The peat layers in ORTO06 and LUNI5 directly ontop of the alluvial fan (incompressible) can be considered as basal peat, thus not subsiding due to sediment compaction. Moreover, the peat layer at the bottom of LUNI5 may be considered as basal peat, i.e., the lowermost basal peat formed during the Holocene transgression (sensu Lange and Menke, 1967; see also Behre, 2004).
For each core, compaction was calculated based on a double approach (Tab V). The field method was inspired by the work of van Asselen (2011), in which isochrones derived from basal peat are employed to “decompact” sediments, i.e. to relocate them in their original position with respect to present-day sea level. In this paper we have dated samples of basal peat in order to trace isochrones across the lagoon deposit; we used the two 14C ages of ORTO06 and LUNI5, determined just on top of the incompressible basement (“base of basal” according to the terminology of
Engelhart and Horton, 2012), and calculated a “sedimentation rate above the incompressible basement” based on age and vertical distance between the two basal samples. We thus used this sedimentation rate to decompact SLIs intercalated in other cores. This operation yields a neglectable uncertainty because the basal peat sample in LUNI5 is the oldest of the dataset. The procedure and the results are outlined in Table V.
In the geotechnical method the computation of lowering was approached using Terzaghi’s law of effective stress (Terzaghi, 1947) and applying a one-dimensional model which describes 281 282 283 284 285 286 287 288 289 290 291 292 293 294 295 296 297 298 299 300 301 302 303 304 305 306 307 308
primary and secondary compression (Van Asselen, 2010, p. 45, 2.7). The geotechnical model was driven by the geologic one. Considering the lithology and grain size of sediments, the share of compaction of each layer which contributes to the stratigraphic sequence underlying each SLI was calculated (see supplemental material). Since the site-specific values for the necessary geotechnical parameters are not available, the minimum and maximum vertical settlement was computed for each layer, using both the minimum and maximum values recommended for each grain size. These values were cautiously selected considering the physical properties of the water basin. The
calculated settlement refers only to the period after the deposition of single samples considered, wherefore they do not represent the overall basin subsidence.
Cross-checking the results obtained with the two different methods (Table V) suggests that they are consistent, except for LUNI5/15. Therefore, the values from the field method were used to relocate the SLIs in order to correct for the compaction effect. These corrections are rather different from sample to sample, ranging from a minimum of 0.14 m for LUNI5/15, the oldest among
intercalated SLIs, located close to the incompressible basement, to a maximum of 2.15 m for LUNI16/13H, located in the central part of the core retrieved from the middle of the basin;
consistently the more recent SLI from core LUNI16 displays a high amount of compaction despite its recent age.
The corrections remarkably change the position of the index points fixed by the selected SLIs, reported in Figure 3 provided with a vertical error bar scaled according to their accuracy. In the last 6500 years, the indicators considered in this study suggest a mean rate of sea-level rise of about 1.1 mm/yr.
Discussion
The elevation of the RSL index points worked out for the study area was interpreted as the effect of eustatic, glacio-hydro-isostatic and tectonic components. In fact the effect of sediment compaction, which proves to be critical using sedimentological sea-level markers (Giosan et al., 2006; Long et al., 2006; Törnquvist, 2008; Horton and Shennan, 2009; Brain et al., 2011; van Asselen et al., 2011; Horton et al., 2013), was efficiently corrected. The tectonic component was 309 310 311 312 313 314 315 316 317 318 319 320 321 322 323 324 325 326 327 328 329 330 331 332 333 334 335 336
evaluated subtracting the elevation of the available index points, corrected for compaction, from that of the corresponding points inferred from available GIA models at the same age (cf. Figure 3). The displacement between observations and models is justified assuming that the incompressible basement was subjected to subsidence, which is consistent with the general tectonic model for the area.
Compared to those extracted from GIA models (Figure 3), our points attain a position which is lower than that of the curve from the Versilia model by Lambeck et al. (2011) and much lower than the envelope from the Magra Plain model (this work), although the general trend is similar to both.
Rates of tectonic subsidence based on the gap between models and observations were calculated for three different time frames, considering the difference in elevation between sea-level estimates from our data and from each of the models employed (Table VI). They range from 0.2 to 0.5 mm/yr according to the reference model. In both cases during the last seven millennia (6630 BP is the age of SLI LUNI5/21), tectonic subsidence is lower in the 7000 – 2500 years BP time span and much higher in the 2500 to present day time span. The values of tectonic subsidence calculated with respect to the Versilia Plain model by Lambeck et al. (2011) though are, for each time span, about 50 % of those yielded by the Magra Plain model (this work). The origin of the difference in the predictions obtained using the models developed by Kurt Lambeck and co-authors and ICE-5G(VM2) have been addressed in section 5 of Antonioli et al. (2009). They arise because of the different assumptions about the rheological profile of the mantle and the history of melting of the late-Pleistocene ice sheets.
In order to support or otherwise our interpretation of the misfit between observed and modelled palaeo-sea levels in the study area as the effect of tectonic subsidence, we compared the subsidence rates obtained in this work with those of a number of case studies available from coastal plains and delta environments worldwide (Table VII) . Particularly significant are the other Mediterranean case studies where coastal subsidence is not limited to the mouth of large rivers, but affects also minor coastal plains, due to the dynamic nature of its coastal profile. Especially in the Aegean Sea, where uplift is the dominant vertical motion type, subsidence is accounted for along 337 338 339 340 341 342 343 344 345 346 347 348 349 350 351 352 353 354 355 356 357 358 359 360 361 362 363 364
specific coastal tracts. Apart from the already mentioned work of Vött (2007), an exhaustive compilation of vertical displacements assessed comparing RSL index points to GIA modelled sea-level elevations throughout the Holocene is provided by Pavlopoulos et al. (2011). Trends of negative displacements are typical of the central Aegean region, affected by extensional tectonics, along tracts of coastal Asia Minor and in the Thessaloniki Plain; calculated rates, on average around 1 mm/yr, reach up to 6.1 mm/yr along the western coast of Crete. In the paper by Pavlopoulos et al. (2011), though, only overall subsidence is calculated, without disentangling the tectonic component from other ones. In particular estimates of compaction are currently unavailable in coastal plains, such as, e.g., the Thessaloniki Plain (Stiros, 2001; Psimoulis et al., 2007) where this factor is likely to affect to a large proportion overall subsidence rates.
Although aseismic vertical displacement due to motion along fault planes as well as to large-scale crustal movements are probably the driving mechanism of tectonic subsidence, in some cases the occurrence of coseismic ruptures in moderately subsiding areas was demonstrated (e.g. Papanikolau et al., 2015: 6 cm in a single event). Cundy et al. (2000) reported that subsidence occurred in the coastal area of the Gulf of Atalanti (Central Greece) as consequence of two moderate-magnitude earthquakes (two events in 7 days: the first one M6.2, the second one M6.9) was due to localised slumping, liquefaction and broader tectonic downwarping. Bertrand et al. (2011) and Kızıldağ et al. (2012) highlighted evidence of co-seismic subsidence in coastal plains of Asia Minor.
In the Po delta, the coastal edge of the largest plain of Italy, subsidence rates have been estimated 1 to 2 mm/yr throughout the Holocene (Carminati et al., 2003). . According to Teatini et al. (2011), the present measured subsidence rate (around 4 mm/yr) should be mainly ascribed to compaction of sediments, whereas Cenni et al. (2013) account for the existence of a tectonic component of the measured subsidence rate.
Further evidence of the late Holocene subsidence of the Magra Plain derives from the comparison of the elevation of our SLIs at ca. 2,000 years BP (i.e., ORTO06 2200 years BP) and that of the index point at Varignano Cove (Figure 1a) where a Roman maritime Villa provides a reliable geoarchaeological marker. This latter testifies the sea-level height during Roman times 365 366 367 368 369 370 371 372 373 374 375 376 377 378 379 380 381 382 383 384 385 386 387 388 389 390 391 392
recorded along a tract of the stable, rocky coast 20 km W of Luna (Figure 1a). The rocky coast of Liguria is generally considered stable or affected by a very moderate uplift (Federici and
Pappalardo, 2006; Rovere et al., 2011), due to the position of the MIS 5.5 shoreline.Measuring and interpreting that marker in terms of its functional height, Chelli et al. (2005) stated that sea level at 90–80 BC was not higher than 70 cm (0.715±0.001 m, referenced to the Italian Ordnance Datum) below present sea level. This indicator was enclosed in the review of Antonioli et al. (2009), and the sea-level elevation inferred was -0.85±0.40 m asl. Estimates of sea-level elevation during the Roman period inferred for this site are intermediate between those provided for the northern Tyrrhenian Sea by Evelpidou et al. (2012) and Lambeck et al. (2011). The index point represented by the Varignano Cove geoarchaeological indicator provides, for the 1st century BC, a RSL estimate intermediate between those yielded by the two GIA models considered in this work. This enables us to confidently interpret the reason for the misfit of our RSL observations with both curves reported in Figure 3 as the effect of local subsidence.
Our data suggest that subsidence rate was not constant throughout the analysed time frame. In fact the mean subsidence rate calculated for the time frame 2500 years BP – present day is higher than the mean value calculated from 6000 to 2500 years BP as well as for the one calculated for the whole investigated period (Table VI). This means that within the period 2500 years BP – present day an event occurred that produced an exceptional displacement along one of the graben faults and caused a down-throw of the Plain (coseismic subsidence) or that within this period, starting from an unknown date, the subsidence rate increased (aseismic subsidence). In the first case the only event accounted for in that time frame occurred between the 4th and 5th century AD
(Rossignani, 1989). . The earthquake occurrence has been deduced by archaeological evidence: e.g., collapsed house walls, columns in the Great Temple fallen down following an unique direction, several crushed dishes and furniture testifying also the sudden interruption of domestic activities (Ward Perkins, 1978; Durante, 2001).
Conclusions 393 394 395 396 397 398 399 400 401 402 403 404 405 406 407 408 409 410 411 412 413 414 415 416 417 418 419
Although in a great number of works the tectonic component of the sea-level equation was obtained by subtracting elevation data of a GIA model from those of a reconstructed RSL curve, it is quite unusual to get, with the same method, estimates of the tectonic component responsible for negative land motion from overall subsidence estimates. This is only possible if the contribution of those factors responsible for subsidence that are not linked to the dynamic behaviour of the
uncompressible bedrock, can be independently quantified. In our case study the tectonic component of the Magra Plain was evaluated subtracting the elevation of the available sea-level index points, corrected for compaction, from that of the corresponding points inferred from available GIA models. The good quality of the index points employed and the cross-checking of two independent methods for calculating sediment compaction, yields a reliable estimate of a tectonic subsidence rate of 0.5 mm/yr. This is consistent with the general tectonic model available for the area that accounts for an active deformation field dominated by extensional strain. Accommodation of the motion is achieved through both aseismic and, probably, coseismic mechanisms, as testified by the sharp increase in subsidence rate highlighted by our data after 2500 years BP in connection with a documented palaeo-earthquake. Our results are also enforced by the cross-checking of the sea-level curve extracted from the Magra Plain GIA model with a geoarchaeological marker available close to the study area, located on a stable, rocky coast.
Acknowledgements
This work has been funded by research grants of the Dipartimento di Fisica e Scienze della Terra "M. Melloni", Parma University (A. Chelli personal funds), Dipartimento di Scienze della Terra, Pisa University (M. Pappalardo, personal funds) and of the, Dipartimento di Scienze Pure e Applicate (DiSPeA), Urbino University “Carlo Bo” (CUP H32I160000000005 and
H32I15000160001). The University of Cologne (Institute of Geography) provided the drilling equipment and some of the laboratory facilities. Authors kindly acknowledge the archaeological authority in charge for the site of Luni (Soprintendenza per I Beni Archeologici della Liguria), in particular Dr. Lucia Gervasini and Dr. Marcella Mancusi. Authors are particularly indebted to all those people that were very helpful during fieldwork, in particular Nick Marriner, Nicole Klasen, 420 421 422 423 424 425 426 427 428 429 430 431 432 433 434 435 436 437 438 439 440 441 442 443 444 445 446 447
Simone Da Prato and to a number of students from Marburg, Cologne and Pisa Universities that took part to field and laboratory activities (Sandra Noß, Melanie Bartz, Martin Seeliger, Marc Bormann, Rilana Rahut, Helen Kehmeier, among others).
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Captions:
Figure 1 - (a) Location map of the study area (base map from ESRI 816 ArcMap® database). The main faults are plotted (solid line= certain fault; dotted line = uncertain fault); (b)
Geomorphological sketch map of the lower Magra Valley (modified after Bini et al., 2012). (1) bedrock; (2) alluvial plain; (3) alluvial fan; (4) wetland; (5) beach ridge; (6) present-day beach; (7) main stream and/or channel; (8) fault. The profile of Figure 2 is plotted.
Figure 2 - Cores from Luna plain used as source of sea-level indicators accounted for in this work. The logs are representative of the sedimentary facies detected through sedimentological,
geochemical and microfaunal analyses.
Figure 3 - Relative sea-level changes in the Magra Plain during the last 8 ka. Index points from sedimentological markers used in this work are plotted with their error bars. The archaeological indicator (Varignano) taken as a reference for the neighboring rocky coast is plotted at the
maximum sea-level elevation consistent with its functional height (according to Chelli et al., 2005). The curve for the Versilia Plain is redrawn after Lambeck et al. (2011), whereas the envelope derives from different runs of a GIA predictive model worked out for the Magra Plain (this work), based on different mantle viscosity values (see text for details).
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Table II- Palaeo sea-level bands for the different kinds of sedimentologic sea-level indicators (SLIs) used in this work.
Table III - Amount of compaction estimated for not “base of basal” sea-level indicators (SLIs). Table IV- Description of the main features of the sedimentary facies recognized in the cores of Figure 2, upon which palaeoenvironmental interpretation is based (data from Bini et al, 2009a, Bini et al 2012 and Noß, 2012).
Table V - Correction in the elevation of sea-level indicators (SLIs) after sediment decompaction. Table VI- Rates of tectonic subsidence based on the gap between models and observed data
calculated, for three different time frames, considering the difference in elevation between sea-level estimates from our data and from each of the GIA models employed (see text).
Table VII- Holocene subsidence rates calculated for different coastal areas or deltas worldwide. 782 783 784 785 786 787 788 789 790 791 792 793 794 795
