Mid-Ludfordian coeval carbon isotope, natural gamma ray and magnetic susceptibility excursions in the Mielnik IG-1 borehole (Eastern Poland)—Dustiness as a possible link between global climate and the Silurian carbon isotope record

Palaeogeography, Palaeoclimatology, Palaeoecology 339-341 (2012) 74–97

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Mid-Ludfordian coeval carbon isotope, natural gamma ray and magnetic susceptibility excursions in the Mielnik IG-1 borehole (Eastern Poland)—Dustiness as a possible link between global climate and the Silurian carbon isotope record Wojciech Kozłowski a,⁎, Katarzyna Sobień b a b

Faculty of Geology, University of Warsaw, Żwirki i Wigury Str., 93, 02–089 Warsaw, Poland Polish Geological Institute—National Research Institute, 4 Rakowiecka Street, 00–975 Warsaw, Poland

a r t i c l e

i n f o

Article history: Received 1 July 2011 Received in revised form 2 April 2012 Accepted 23 April 2012 Available online 2 May 2012 Keywords: Stable carbon isotope Spectral gamma ray Magnetic susceptibility Sequence stratigraphy Carbonate periplatform setting Silurian

a b s t r a c t The graptolite-bearing upper Silurian succession of the Mielnik IG-1 borehole (eastern Poland) represents a periplatform setting on a neritic carbonate platform located on the palaeocontinent of Baltica and records widely recorded positive mid-Ludfordian Carbon Isotope Excursion (CIE). The initial increase of δ13C values occurs in the uppermost part of the Neocucullograptus kozlowskii Zone whereas pronounced increase and maximal values (+ 6.74‰) are noted in the N. kozlowskii–Pseudomonoclimacis latilobus interzone (Pristiograptus dubius bloom interval). The thickness–time relation in the section indicates that the main, initial shift in carbon isotope ratio (of 2.7‰) occurred over a very short time, estimated to be around 4000 years, and is coeval with a rapid initiation of deposition of rock-forming, silt-sized detrital dolomite and quartz. The CIE decline is interpreted as steeper than in inner shelf sections, and is coincident with a flooding event in the Ps. latilobus–Slovinograptus balticus Zone; whereas the end of the excursion is coeval with the next flooding surface marked by the Ps. Latilobus–Uncinatograptus acer graptolite assemblage. Combined SGR (Spectral Gamma Ray), facies, and faunal records during the CIE suggest prolonged sea-level lowstand conditions. The abundant detrital material (‘dolomite siltstone facies’), occurring exclusively in the CIE interval, is noted in microfacies as well as in the geophysics record—as pronounced negative excursions in total natural gamma ray and magnetic susceptibility, both resulting from dilution by diamagnetic and low-radioactivity minerals. The perfectly sorted, vertically equally-sized silt material contains unique, constant and low diversified grain assemblage (euhedral dolomite, quartz, minute ooids)—suggesting its provenance exclusively from the shallower and/or emerged parts of the carbonate platform interior. The spot findings of graptolites and nautiloids along with delicate parallel lamination of sediments indicates relatively deep environment, with absence of traction processes. These features, along with the lack of dolomite detritus in accompanying redeposited carbonates, suggest its derivation by wind. With regard to the earlier noted presence of aeolian detritus in the middle Ludfordian, we suggest dustiness as a plausible driving force of the CIE. In the hypothesis proposed, bioavailable Fe-bearing dust inflow in concert with antiestuarine circulation on shelfs, caused nutrient (N:P ratio) disturbances, resulting in methanogenesis in surface layer of the seas. The surface layer methanogenesis should resulting in bathymetric gradient of isotope fractionation, what is consistent with the CIE record. Coincidence of sea-level lowstand, low-latitude aridity, and gustiness of the global climate, are in accordance with the hypothetical short-lived glaciation in mid-Ludfordian time. © 2012 Elsevier B.V. All rights reserved.

1. Introduction The Silurian is a time of dynamic, global climate changes between the Hirnantian icehouse and Early–Middle Devonian greenhouse conditions. Carbonate rocks of this interval record a number of distinct, mostly positive carbon isotope excursions (CIE) noted worldwide, which have high potential for stratigraphic correlation ⁎ Corresponding author. E-mail address: [email protected] (W. Kozłowski). 0031-0182/$ – see front matter © 2012 Elsevier B.V. All rights reserved. doi:10.1016/j.palaeo.2012.04.024

(e.g. Kaljo and Martma, 2006; Cramer et al., 2011). Since the late 90s, numerous papers record, interpret and discuss the causes of the Silurian CIEs (reviewed in Munnecke et al., 2010). The coincidence of the Silurian CIEs with distinct sea-level falls, along with evidence of glacial events in high latitudes in early Silurian (e.g. Grahn and Caputo, 1992; Diaz-Martinez and Grahn, 2007) has provided a basis for a discussion of interaction between the evolution of climate, glaciations, sea-level changes, carbon cycle and the origin of carbon isotope record (see Loydell, 1998; Munnecke et al., 2003; Loydell, 2007, 2008; Cramer and Munnecke, 2008).

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Compared to earlier Silurian events, the mid-Ludfordian CIE reaching maximal δ 13C values around + 10‰ seems to be very enigmatic. The excursion is associated with sea-level fall (e.g. Eriksson and Calner, 2008; Kozłowski and Munnecke, 2010), and has been noted from low- to mid-palaeolatitude carbonate platforms of Baltica, Laurentia, East Gondwana (Australia), Peri-Gondwana (Bohemia, Carnic Alps) and Avalonia (see Munnecke et al., 2010 and references therein; Loydell and Frýda, 2011). The carbon isotope shift is approximately coeval with extinctions: termed the Lau event (with respect to shallow shelf biofacies—Jeppsson, 1990) and the kozlowskii event (for graptolite biofacies—Urbanek, 1993). The rock record of this interval from lower palaeolatitudes suggests the presence of a cooling event (Kaljo et al., 2003; Lehnert et al., 2007a; Eriksson and Calner, 2008), but confirmation of icehouse conditions by direct evidence in higher palaeolatitudes has not yet been found. Here, we present new data from the Mielnik IG-1 borehole, where Silurian sediments were deposited at low palaeolatitudes, on the shelf of Baltica in a carbonate-bearing periplatform setting (Fig. 1). The relatively uniform deep-water facies development, considerable thickness and completeness of the classical graptolite succession in this section (Urbanek, 1970, 1997), suggest continuous sedimentation with a relatively high and possibly nearly constant rate of deposition. The Mielnik IG-1 succession records the coincidence of a CIE with sea-level and climate changes recorded by: unusual facies, spectral gamma ray (SGR), magnetic susceptibility (MS) and petromagnetic characteristics. We have tried to compare this record with some examples of low-latitude offshore sedimentation during the last glacial, looking for analogies, which might support the hypothesis of a short-

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lived glaciation in the mid-Ludfordian (e.g. Kaljo et al., 2003; Lehnert et al., 2007a; Eriksson and Calner, 2008). The aims of the present study are: (1) correlation of the CIE with Urbanek's (1970, 1997) high-resolution graptolite biostratigraphical framework; (2) documentation and explanation of the natural gamma and magnetic susceptibility anomalies which are associated with the CIE; (3) revision of the sequence stratigraphy framework in the excursion interval; (4) providing evidencing for the important wind derived admixture during the CIE—noted in the periplatform setting and its influence on the natural gamma ray and petromagnetic record; (5) discussion of gustiness and eustacy as important triggers for the mid-Ludfordian CIE.

2. Material and methods The Mielnik IG-1 borehole, situated in eastern Poland (Fig. 1), was drilled in 1959, with assumed complete coring, as one of stratigraphical reference points in the Polish Lowland (Tomczyk, 1962, 1964). Unfortunately, the Ludfordian part of the succession contains a few intervals with incomplete core yield, including one of the key parts of the studied event (interval 823–836 m—representing the part of the CIE decline). The three-dimensional preservation of graptolites enabled very detailed palaeontological studies followed by high resolution biostratigraphical subdivision of the succession (Urbanek, 1970, 1997), which is used in the present study (Fig. 2). Urbanek's studies resulted in the Mielnik IG-1 borehole being one of the best graptolite-dated upper Silurian successions in the world, hence it is a very important section for global upper Silurian stratigraphy (see e.g. Sadler et al., 2009). This study, however, did not include information on the abundance of analyzed taxa, therefore, a re-examination of the core has been carried out by the author in order to include semi-quantitative graptolite information (Fig. 3E).

2.1. Sampling procedure In the present study a detailed lithological profile of the core interval from the bottom of kozlowskii Zone to the top of spineus Zone (approximately the middle part of the standard formosus Zone) is analyzed. From each meter (box) of the core, samples representing calcareous-claystone facies have been collected from the lower and upper parts (in the excursion interval) and for the upper part (in intervals adjacent to the CIE) for the purposes of isotope and petromagnetic studies. Additionally, limestone beds have been sampled for microfacies studies and for comparison of the carbon isotope record between limestone facies and the marly background. Microscopic photographs were taken in the Scanning Electron Microscope and Microanalysis Laboratory of the Faculty of Geology, Warsaw University, on a Nikon ECLIPSE E600W POL with LUCIA software.

2.2. Carbon isotopes

Fig. 1. Paleogeographical map of the Baltic region in the time of the mid-Ludfordian regression (after Kozłowski and Munnecke, 2010 and references therein, partly changed); USB: Upper Silesian Block, TTL: Teisseyre-Tornquist Line.

For the purposes of stable isotope studies, 77 whole-rock samples of vein-free calcareous claystones and 23 samples of vein-free pelitic limestones were selected and crushed in a mortar. Carbon and oxygen isotope values were determined in the Stable Isotope Laboratory of the Polish Academy of Sciences in Warsaw. Sample powder was treated with phosphoric acid in a Kiel IV preparation system and analyzed in a conjunct Finnigan Delta+ mass spectrometer. Values are reported in the conventional delta notation with respect to the Vienna Pee Dee Belmnite (VPDB). Reproducibility for the isotopic analysis was better than ±0.1%.

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Fig. 2. Vertical ranges of graptolites in upper Silurian of the Mielnik IG-1 section, according to Urbanek (1970, 1997).

2.3. Magnetic susceptibility and petromagnetic parameters Magnetic susceptibility and other rock magnetic measurements are applied to characterize the magnetic mineralogy of the rocks studied. A total of 119 sliced rock samples from Mielnik IG-1 core were collected thought a 95.8 m thick interval (782.0 to 877.8 m). Samples were crushed, placed in 7 cm plastic boxes and weighed.

Rock magnetic analyses were carried out in the Palaeomagnetic Laboratory of the Polish Geological Institute—National Research Institute in Warsaw. Low field magnetic susceptibility was measured using a KLY-2 Kappabridge (Geofyzika, Brno) with on operating frequency of 920 Hz and sensitivity of 4 × 10− 8 SI. All of the obtained magnetic susceptibility values are normalized to the sample weight of 10 g and expressed in 10− 6 SI units. Rock magnetic parameters such

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as concentration-dependent anhysteretic (ARM) and isothermal (IRM) remanent magnetization, as well as ferromagnetic grain sizedependent S-ratio have been defined, in order to characterize magnetic mineralogy sensitive to environmental changes. ARM has been acquired directly in a 100 mT alternating magnetic field and 0.1 mT bias fields (AF Demagnetizer, Molspin Ltd, UK). IRM has been imparted directly in a field of 1.5 T for all samples using a MMP1 pulse magnetizer (Magnetic Measurements, UK). Stepwise acquisition of the IRM has been performed for 10 selected samples. S-ratio is calculated according to Bloemendal et al. (1992): S− 0.3T = [−(IRM− 0.3T/SIRM) + 1]/2, where SIRM is the IRM field measured after applying a pulsed field of 1.5 T, IRM− 0.3T was measured after imparting a backfield of 0.3 T along the same axis. An intensity of remanent magnetization has been measured with on Agico JR-6A spinner magnetometer (noise level 10− 5A × m − 1, Czech Republic). Additionally, to obtain more detailed curves, MS was measured directly in the core magazine at 0.2 m spaced intervals using a SM-30 Magnetic Susceptibility Meter by means of the temperature drift compensation method (ZH instruments, Czech Republic; operating frequency of 9 KHz, sensitivity 1 × 10 − 7 SI). 2.4. Spectral gamma-ray (SGR) The spectrum of the natural gamma radiation was analyzed directly in the core magazine using a portable Gamma Surveyor GMS/CN gamma ray spectrometer (GF Instruments, Czech Republic). For each 1 m core interval (box) three measurements were performed, each of 3-minute duration; eK, eTh and eU values (radiometrically determined abundances of potassium, thorium and uranium respectively) were subsequently averaged for each box. Natural gamma-ray total dose, measured in the core, shows good correlation (0.74) with the archival borehole log data (Deczkowski, 1960), and both curves show very similar local trends. Comparison of trends of relative dose in the borehole and in the core, along with diameter and electric resistance data, have been used locally also in reconstruction of the boxes position in a few incomplete core intervals. The general trends in absolute element abundance presented (Fig. 3J) have been recalculated from their mutual ratios measured in the core (natural gamma spectrum—Fig. 3K) and archival total gamma ray fluctuations measured in borehole (natural gamma intensity— Fig. 3H). The values recalculated according this procedure are in most cases very similar to this obtained directly from the core, however in the few intervals with strongly dilapidated core the recalculated values are slightly higher. Because of obvious imperfections of the whole procedure (the influence of possible shifts between the core and borehole data), we try to focus mainly on the proportion of each element in the total dose and ratios between eK, eTh and eU (natural gamma spectrum) as values not dependent of shape and size of the core. 3. Geological background The Mielnik IG-1 borehole is located in Eastern Poland in the Podlasie Depression. The undeformed, horizontally lying Silurian rocks have been documented between 531.8 m and 1138.4 m (606.6 m thick) and represent the Wenlock to lower Přídolí series (Tomczyk, 1970). The Wenlock rocks overlie with an erosive gap (unconformably) Ordovician [1138.4–1236 m (98 m thick)], Cambrian [1236–1576.3 (350 m)], and Ediacaran [1576.3–1728 m (152.7 m)] sedimentary rocks of platform cover. Between 1728 m and 1813 m the Proterozoic crystalline basement of the East European Craton has been found.

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The Silurian deposits of the Mielnik IG-1 core represent a succession deposited on the south-western shelf of the palaeocontinent of Baltica. In the late Silurian the borehole site was located in a transitional zone between open-shelf carbonate facies and deep-water graptolitic claystones (Fig. 1), and its facies position may be designed at a carbonate periplatform setting. The dominant Silurian facies in the core is light grey calcareous claystone with graptolites and nautiloids. In some intervals thin laminae and beds of pelitic, occasionally fine-grained limestones occur. The approximate rate of deposition (calculated without decompaction) for the Wenlock rocks (19 m/Ma on average) accelerated rapidly at the bottom of the Lobograptus scanicus Zone (mid-Gorstian, early Ludlow) up to 107.5 m/Ma on average, which reflects basin evolution from a passive continental edge to a foreland basin setting (Poprawa et al., 1999). This change is also reflected in the predominant facies. The Wenlock is represented by pelitic to organodetrital limestones with abundant benthic—nektic open shelf fossil assemblages (trilobites, brachiopods, bivalves, cephalopods); with few thicker intercalations of graptolitic calcareous claystones occurring in the Cyrtograptus murchisoni, Monograptus riccartonensis and C. lundgreni - Pristiograptus ludensis zones (Tomczyk, 1970). The Ludlow and Přídolí are represented mainly by graptolitic calcareous claystones with thin intercalations of pelitic limestones. The high thickness, along with the low-energy sedimentary environment and completeness of graptolite succession, suggest continuous and a nearly constant rate of sedimentation. For an estimation of the duration of observed events, an average equivalent of 9500 years for 1 m has been assumed for the studied interval.

4. Results 4.1. Facies record The studied interval of Mielnik IG-1 succession is composed mainly of calcareous claystones. Most variability in lithology in the studied interval of the Mielnik IG-1 succession results from differential proportions of claystones and carbonate intercalations (see also Fig. 4 in Urbanek, 1997). Limestone beds are generally thin—a few or few tens of cm thick; rarely up to 50 cm thick. The second significant differentiating feature is the different carbonate content in the claystones, which in a couple of intervals are almost carbonate-free. The microfacies spectrum of limestone intercalations is presented in Table 1. Based on macroscopic lithology characteristics and microfacies studies, the following succession of lithological complexes has been outlined (Fig. 3D): L1) 875.5–866 m (boxes: 375–367) grey calcareous claystones with thin (up to 10 cm thick) rare fine grained detrital wackestone intercalations; L2) 866–858.9 m (boxes: 366–360) grey slightly calcareous claystones distinctly laminated (“varve-like”—Fig. 4C), with alternation of blue-green laminae composed of finer clay, and grey—quartz pelite bearing; with rare, thin (up to 10 cm thick) coarser grained, sometimes neomorphosed—detrital wackestone intercalations; L3) 858.9–853.8 m (boxes: 359–355) grey slightly calcareous mudstones, laminated (“varve-like” – Fig. 4A); locally crosslaminated (Fig. 5A), with, thin, often partly neomorphosed, lenticular-cross laminated detrital packstone to graded-grained grainstone intercalations;

Fig. 3. Lithological, geochemical and geophysical data of Mielnik IG-1 section and its sequence stratigraphic interpretation. A) depth in metres; B) core interval/box; C) lithological complexes; D) lithological log; E) macrofossils occurrences (current examination of the core); F) graptolite ranges (biostratigraphy)—after Urbanek, 1970, 1997; G) carbon isotope data; H) total, natural gamma ray log (after Deczkowski, 1960); I) magnetic susceptibility (laboratory measurements—squares, on core measurements—background line); J) SGR data: eK, eTh, eU values; K) SGR data: contribution of eK, eTh, eU in total dose; L) eK*eTh and (eK*eTh)/eU proxies (relative facies “proximity”, relative rate of deposition); M–N) ARM and IRM values; O) S-ratio; Interpretation: P) Sequence stratigraphic framework; R) relative sea-level changes; horizontal lines: blue—flooding surfaces, orange—maximal progradation surfaces, green—kozlowski extinction.

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L4) 853.9–850.5 m (boxes: 354–352) grey calcareous silty mudstones massive or locally distinctly laminated (“varvelike”—Fig. 4B); L5) 850.5–843.2 m (boxes: 352–347) light grey, dolomitic siltstone massive or locally finely laminated (Figs. 4C, 6); L6) 843.2–841.5 m (boxes: 346–345) light grey dolomitic siltstone (Fig. 6) with laminae of ostracode wackestone and blue-green fine clay material (“varve-like” texture), (Fig. 6B); L7) 841.5–838.2 m (boxes: 344–342) grey calcareous claystone intercalated with limestones (ostracode wackestone, detrital packstone or peloidal graded grained grainstone facies—Fig. 5E), with mass occurrence of Pristiograptus dubius at 840 m (box: 343); L8) 838.2–836.6 m (box: 341) grey slightly calcareous claystones with fine muscovite flakes; L9) 836.6–834.8 m (boxes: 340–339) grey marls interbedded with ostracode wackestone (Fig. 5D) with abundant trilobites, crinoids and ostracodes; L10) 834.8–830.0 m (boxes: 338–337) grey calcareous claystones; L11) 830.0–822.2 m (boxes: 336–334) grey slightly calcareous claystones interbedded with wackestone with biomorpha and graded grained grainstone (Fig. 5B,C); L12) 822.2–817.2 m (boxes: 333–329) grey calcareous claystones interbedded with detrital wackestone (part of the beds are completely neomorphosed); L13) 817.2–807.2 m (boxes: 328–323) grey calcareous claystones with scarce detrital wackestone interbeds; L14) 807.2–794.5 m (boxes: 322–314) grey, slightly calcareous claystones interbedded with detrital wackestones (Fig. 5F); L15) 794.5–782.2 m (boxes: 313–301) grey calcareous claystones with single interbeds of detrital wackestone and wackestone with biomorpha /box 307/; A prominent facies anomaly occurs between 850.5 m and 835 m. The lower part of the interval consists of very fine-grained dolomitic siltstone recording maximal CIE values (see Section 4.2). The “Dolomitic siltstone” (Fig. 6) form texturally massive, to delicate laminated, 8-m-thick continuous single bed, with wackestone to packstone texture (Fig. 6A). Dolomite grains are Fe-rich (according to preliminary chemical analysis) and represented by euhedral crystals, rounded grains and fragments or whole spheroids (Fig. 6C–D). The grains are very fine (from 0.01 to 0.05 mm) and equal-sized in any single horizon. Dolomite grains across the profile vary in abundance and size (Fig. 6A–B), locally forming finely-spaced delicate lamination (Fig. 6B). The grains in laminae often show gradation, with partly neomorphosed coarser (upper) part of laminae (Fig. 6F). Quartz grains, forming up to 20% of rock volume (Fig. 6C), occur as very finely-grained pelite (around 0.01 mm in diameter); or single larger grains up to 0.1 mm in diameter. Locally, in quartz-rich levels, abundant fine mica flakes occur. Dolomite-rich horizons are enriched in relatively larger quartz, and vertically the positive correlation in grain size and in abundance between the dolomite and quartz can be observed. The coarser horizons contain also numerous crushed and/or abraded carbonate spheroids up to 0.1 mm in diameter (Fig. 6C) and minute (ca 0.05 mm) ooids (Fig. 6E). Scarce bioclasts are of similar size and do not show any signs of recrystallization (dolomitization) (Fig. 6G). Numerous flat aggregates (up to 0.3 mm wide) of amorphous organic matter (AOM) occur preferentially in coarse-grained parts of the rock;

Fig. 4. Succession of rock textures around the positive carbonate isotope excursion interval; from the bottom (A) to the top (C): A) “Varve-like” lamination in the L3 complex, note preferential occurrence of amorphous organic matter flakes near or inside the quartz-rich laminae; B) mudstone from top of the L4 complex (depth 851 m) containing relatively higher admixture of quartz pelite, occurring in both thin laminae and thicker interlayers (probably partly reworked); C) Dolomitic siltstone in the same scale for comparison, note relatively coarser quartz pelite, without reworking interlayers and more abundant amorphous organic matter aggregates.

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Table 1 Characteristics of limestone microfacies noted in studied interval. Facies

Texture

Allochems

Matrix

Other remarks

Interpretation

WB wackestone with biomorpha

Wackestone

Whole, or partly fragmented parts of nautiloids, graptolites, ostracods and gastropods

Dark micrite

Beds are often cut by erosive surfaces on the top

Pelagic sedimentary condensation levels

OW ostracod wackestone

Wackestone sometimes with laminas or lenses of packstone to grainstone

Numerous bivalved ostracods, crinoid fragments; in the grainy bodies larger dark peloids, shell detritus and silt-sized quartz grains are concentrated

Mainly dark micrite

Ostracod shells are in most cases filled by pyrite and sparite which form geopetal structures

Distal hemipelagic sedimentation with carbonate influx from platform

DW–DP detrital wackestone

Fine grained wackestone to coarser grained packstone with discontinuous laminae or lenses of packstone to grainstone

Small dark peloids and silt-sized quartz grains concentrated in the grainy bodies; graptolite rhabdosomes, nautiloid and ostracode fragments

Light grey micrite to microsparite, in places neomorphosed

DW–DP facies form the general background of the carbonate sedimentation in the studied interval

Peri-platform sedimentation with considerable influx of platform derived fine material

G graded grained grainstone

Graded grained grainstone passing upward to packstone

Silt-sized various bioclasts and dark peloids

Sparite, locally neomorphosed along with grains

Up to 1 cm thick beds starting with erosive surfaces; δ13C values in pelitic limestone are often 0–1‰ elevated with respect to the general trend

Finely grained, distal calcareous turbidite; periplatform setting

DS dolomitic siltstone

Wackestone with very fine, very well sorted grains; locally with very thin packstone laminae (with gradual base and top sharp transitions)

Dark, clay bearing Equal sized, minute, euhedral, to partly or entirely rounded dolomite micrite crystals and spheroids; silt-sized quartz grains; minute ooids; small bioclasts; amorphous flat-shaped clusters of organic matter

The content and size of dolomite, quartz grains and content of organic matter clusters are positively correlated; bioclasts and ooids are not recrystallized

Periplatform low energy environment with high aeolian-derived admixture

Rare layers (up to 5 mm thick) of radiaxial fibrous cements; δ13C values lowered by 1–2‰ in respect to the marly background

Carbonate sediment neomorphosed due to sulphate reduction in anoxic conditions

NL neomorphosed Neomorphosed limestone

Amorphous organic matter partly replaced by bunch-like aggregates of (?framboidal) pyrite, relicts of fossils

however, single aggregates, of identical size, have been noted also in fine-grained parts of the rock (Fig. 6A, C). Very similar lamination and distribution of flat organic aggregates is visible in the “varve-like” mudstones occurring immediately below and above the dolomitic siltstone interval (Fig. 4A). The rock is composed of clay and siltbearing laminae. Clay laminae contain very small quartz pelite and single graptolites. Silt-bearing laminae are composed of coarser, equally sized (around 0.02 mm), densely packed quartz grains. Inside and/or on top of the silt laminae numerous aggregates of amorphous organic matter occur (Fig. 4A,B). 4.2. Stable carbon isotope record in relation to the graptolite biostratigraphy Whole-rock data from the Mielnik IG-1 section records the globally observed mid-Ludfordian CIE, with maximal δ13C values of +6.74‰ PDB (Table 2, Fig. 3G). Samples of calcareous claystone and marls show well-defined trends, but some of pelitic limestones show anomalous values—lower or higher in comparison with the marly background. Microfacies studies show that positive δ13C point anomalies appear in pelitic to fine-grained limestone associated with intercalations of the graded grained grainstone facies (redeposited sediments), generally in a limited interval—i.e. in and immediately above the CIE maximum. Anomalous lighter values occur in sedimentary rock representing partly neomorphosed detrital wackestone to packstone, characterized by abundant bunch-like aggregates of (?framboidal) pyrite, which have replaced amorphous organic matter (Fig. 7). The CIE observed in marly sedimentary rock of the Mielnik IG-1 section (Fig. 3G) arises from moderately stable, low δ13C values of +0.6 to +1.0‰. The pre-excursion gradual increase in carbon isotope, from +0.6‰ to +2‰, is noted in the upper half of the Neocucullograptus kozlowskii range (upper part of the L2 and L3 complexes), with

Equigranular often with gradually changing crystal sizes, xenotopic, in places porphyrotopic

anomalous values of +0.28 and −0.35‰ noted in partially neomorphosed detrital wackestone to packstone intercalations. At level of disappearance of the N. kozlowskii the δ13C values are +1.87‰ and subsequently increase then decrease slightly in the lowermost part of the N. kozlowskii–Pseudomonoclimacis latilobus interzone (L4 complex), parallel with the increase of carbonate content. The onset of the main excursion is very rapid and occurs between depths of 850.5 m and 850.1 m in the lower part of the N. kozlowskii– Ps. latilobus interzone. The abrupt shift from +1.53 to +4.22‰ occurs between the middle and the top parts of box no. 352, over a distance of the 0.4 m; this is coeval with the first appearance of the dolomitic siltstone facies (which form the L5 and L6 complexes). In the following interval, represented by the dolomitic siltstone facies, δ 13C values rapidly increase up to +5.7‰ and stabilize at this level in its upper part. It is important to note that: (1) samples with maximal dolomite content (e.g. sample no. 351′) are isotopically lighter in respect to other samples from adjacent horizons; (2) recurrent facies change (L6–L7)—from dolomitic siltstone to calcareous claystones, observed immediately above, record the next positive shift in δ 13C values (up to +6.5‰). Both these observations suggest that dolomite grains are not the main carrier of carbon heavy isotope signatures, despite their synchronous appearance with the main isotope shift. The upper part of the maximal CIE interval is coincident with the Pristiograptus dubius fauna bloom in the N. kozlowskii–Ps. latilobus interzone (complex L7), occurring in calcareous claystones with ostracode wackestone interbeds. Maximal δ 13C values (of +6.74) have been noted in the intercalation of pelitic limestones containing the interbeds of reworked (allochthonous) peloidal grainstone (depth 839.9 m). The decrease in the δ13C values is stepwise. In the carbonate-poor intervals (L8, L11) rapid decreases are observed, while in intervals represented by calcareous facies (L9–L10) the decrease in values is

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Fig. 5. Micrographs of facies from Mielnik IG-1 section: A) Cross laminated calcareous mudstone at the base of the regressive facies interval (L3, depth 856 m); B–C) fine, graded grainstone (GGG facies) intercalation (B) co-occurring with condensed wackestone with biomorpha (WB facies) marking post-CIE transgressive interval (L11 complex); D) Ostracode wackestone (OW) facies with crinoids from benthic fauna occurrence interval (L9 complex); E) Ostracode wackestone (OW) facies with interlayer of peloidal grainstone with crinoids. The sample records maximal CIE values in the Ludfordian of the Mielnik IG-1 succession (depth 840 m); F) Typical, background detrital packstone–wackestone facies (DW–DP), depth 803.5 m.

slower. Some limestone intercalations (adjacent to the graded grained grainstone facies) in this interval show heavier values compared to the marly background. At the level of the first appearance of the Pseudomonoclimacis latilobus (L12 complex) values decrease to a level below that preceding the excursion (around 0‰). The last occurrence of

relatively heavy δ13C values (fluctuating around +1‰) is observed in the uppermost part of the Ps. latilobus–Slovinograptus balticus range zone and lower part of the S. balticus–Uncinatograptus acer interzone—in the calcareous complex L13, after which they decrease to around +0.5‰ and remain at this level in the Uncinatograptus acer range zone

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Fig. 6. Micrographs of dolomitic siltstone facies from the Mielnik IG-1 section: A) Vertical change in average grain size, coeval with increase of abundances of flat aggregates of amorphous organic matter; B,F) Lamination in dolomitic siltstone with reverse gradation of grains in silty laminae; C–E) Grain assemblage of the DS facies, quartz grains are highlighted in blue(C); note equal euhedral (D) and abraded dolomite grains and spheroids, and minute well preserved ooids (E); G) Un-neomorphosed bioclasts in silty dolomite matrix.

(where also Psudomonoclimacis latilobus reappears) and in subsequent intervals. In the partly neomorphosed, pyrite-bearing detrital wackestone– packstone intercalations of upper part of the studied interval (L12L15) values around and below − 1‰ have been recorded.

4.3. Spectral Gamma Ray The use of the SGR-method was inspired by earlier observations of the regional distribution of two distinct negative, natural gamma anomalies in the Silurian of the Polish Lowland basins—occuring in

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Table 2 Carbon and oxygen isotope ratio values (vs VPDB) from the Mielnik IG-1 section. 13

18

Box

Depth

Sample

δ C VPDB

δ O VPDB

Lithology

277 282 286 290 292 294 294 296 298 300 302 305 307 309 311 311 313 315 317 318 319 320 321 323 324 324 325 326 327 327 328 329 329 329 330 331 332 332 333 333 334 334 335 335 335 336 336 336 337 337 338 338 339 339 340 340 340 341 341 341 342 342 343 344 345 346 347 348 349 350 351 351 351 352 352 353

758.1 763 767 771.5 773.5 776 776 778 780 781.9 783.6 786 788 790 792 792 794 797 799 799.6 803.5 804.1 805.9 807.9 809.4 809.4 810.4 814.4 815.8 815.8 816.8 817.8 817.8 817.8 818.8 819.8 820.8 820.8 821.8 821.8 822.8 822.8 828 828 828 829 829 829 830.4 830.4 833.5 833.5 834.7 834.7 835.4 835.8 836.2 837.5 837.1 837.5 838.5 838.9 839.9 840.9 841.9 842.9 843.9 844.9 846.1 847.3 848.5 848.5 848.1 850.1 850.5 851.8

277 282 286 290 292 294 294 lm 296 298 300 302 305 307 309 311 31 lm 313 315 317 318u 319 320u 321 323 324 324 lm 325 326u 327 327 lm 328 m 329 m 329A 329B 330 331 332 332 lm 333 333 lm 334 334 lm 335 335A 335B 336 336 lm 336 337 337 338 338 339 339 340u 340 lm 340d 341u 341 m 341 342u 342 m 343 344 345 346 347 348 349 350 351u 351′ 351 sr 352u 352 m 353

0.37 0.53 0.8 0.6 0.18 0.31 − 0.52 0.1 0.79 0.34 0.59 0.55 0.36 0.31 0.17 − 1.66 0.46 0.25 0.53 0.64 0.28 0.53 0.92 1.2 0.9 − 0.21 1.11 0.96 0.44 − 0.56 0.19 0.62 1.3 1.44 0.1 0.44 0.31 0.56 0.97 − 0.26 1.06 1.26 1.68 1.61 1.82 2.54 2.77 − 2.25 3.33 4.14 3.68 3.76 3.78 5.04 3.98 2.86 4.21 5.39 3.87 6.63 5.6 5.53 6.74 6.24 6.53 6.55 5.75 5.71 5.76 5.63 4.8 3.7 5.09 4.22 1.53 1.72

− 6.88 − 6.11 − 6.31 − 5.98 − 6.6 − 5.1 − 5.12 − 6.01 − 6.45 − 6.59 − 5.94 − 6.28 − 5.93 − 5.43 − 6.85 − 4.74 − 6.47 − 6.36 − 5.68 − 6.69 − 5.04 − 6.77 − 5.82 − 6.14 − 6.12 − 4.88 − 6.24 − 6.57 − 6.69 − 5.64 −6.67 − 6.67 − 11.38 − 6.37 − 6.08 − 6.04 − 6.92 − 6.43 − 6.12 − 2.81 − 4.97 − 4.31 − 5.54 − 2.33 − 11.2 − 6.13 − 3.83 − 4.15 − 5.79 − 12.96 − 5.53 − 5.14 − 5.19 − 3.59 − 5.47 − 3.29 − 5.45 − 5.76 − 5.21 − 5.42 − 5.57 − 5.65 − 4.85 − 4.16 − 6.7 − 5.71 − 5.47 − 5.67 − 5.91 − 5.52 − 5.94 − 5.45 − 5.94 − 5.61 − 4.58 − 4.44

Calcareous claystone Calcareous claystone Calcareous claystone Calcareous claystone Calcareous claystone Calcareous claystone Limestone Calcareous claystone Calcareous claystone Calcareous claystone Calcareous claystone Calcareous claystone Calcareous claystone Calcareous claystone Calcareous claystone Limestone Calcareous claystone Calcareous claystone Calcareous claystone Calcareous claystone Calcareous claystone Calcareous claystone Calcareous claystone Calcareous claystone Calcareous claystone Limestone Calcareous claystone Calcareous claystone Calcareous claystone Limestone Calcareous claystone Calcareous claystone Limestone (grained) Limestone (neomorphized) Calcareous claystone Calcareous claystone Calcareous claystone Limestone Calcareous claystone Limestone Calcareous claystone Limestone Calcareous claystone Limestone Limestone (grained) Calcareous claystone Limestone Limestone (neomorphized) Calcareous claystone Limestone (grained) Calcareous claystone Limestone Calcareous claystone Limestone Calcareous claystone Limestone Marl Marl Calcareous claystone Limestone Calcareous claystone Calcareous claystone Limestone Calcareous claystone Calcareous claystone Calcareous claystone Dolomitic siltstone Dolomitic siltstone Dolomitic siltstone Dolomitic siltstone Dolomitic siltstone Dolomitic siltstone Dolomitic siltstone Dolomitic siltstone Calcareous claystone Calcareous claystone

Table 2 (continued) Box

Depth

Sample

δ13C VPDB

354 355 356 356 357 358 359 360 361 362 362 363 364 364 365 366 366 367 367 368 369 371 373 375 378 383

853 854.2 855.2 855.2 856.2 857.3 858.3 859.3 860.3 861.3 861.3 862.3 863.3 863.3 864.5 865.5 865.5 867.5 867.5 869.6 870.6 872.6 874.6 878.2 883.7 889.6

354 355 356u 356 lm 357 358 359 sr 360 361 sr 362 362 sr 363 sr 364 364 365 366 sr 366 367 sr 367 368 369 371 373 375 378 383

2 1.87 1.32 − 0.35 2.09 1.56 0.86 1.35 1.04 0.28 0.82 0.66 0.65 0.98 1.06 0.6 1.66 0.67 − 1.22 1.13 0.83 0.9 0.84 0.7 0.96 1.12

δ18O VPDB − 5.11 − 5.07 − 5.98 − 4.57 − 5.01 − 5.38 − 6.03 − 5.45 − 5.2 − 5.72 − 5.55 − 5.44 − 5.24 − 5.19 − 5.31 − 5.61 − 2.82 − 5.75 − 5.09 − 4.06 − 5.07 − 5.57 − 5.53 − 3.82 − 5.44 − 5.14

Lithology Calcareous Calcareous Calcareous Limestone Calcareous Calcareous Calcareous Calcareous Calcareous Limestone Calcareous Calcareous Calcareous Limestone Calcareous Calcareous Limestone Calcareous Limestone Calcareous Calcareous Calcareous Calcareous Calcareous Calcareous Calcareous

claystone claystone claystone claystone claystone claystone claystone claystone claystone claystone claystone claystone claystone claystone claystone claystone claystone claystone claystone claystone claystone

the upper Wenlock (Pristiograptus parvus–Gothograptus nassa Zone) and middle Ludfordian (e.g. Topulos, 1976) i.e. occurring in the vicinity of the Mulde and Lau events. The anomalies have often been used by H. Tomczyk (personal communication, 2001), for identification of his informal stages (see e.g. Tomczyk, 1970) in the case of lack of core. The distinct gamma excursion recognized in the Mielnik IG-1 section (depth interval 858–823 m) is broadly coeval with the middle Ludfordian CIE. Over the whole studied interval the total gamma values (cps—Fig. 3H) show a relatively strong correlation (0.68) with eK values. The eK values show a moderate correlation with eTh values (0.49) and a lack of correlation with eU (−0.05). The eTh and eU values show a slight negative correlation (−0.39) below and above of the gamma excursion interval and a lack of correlation within it. Below and above the gamma excursion the SGR shows a cyclic pattern with distinct coupled positive eU–negative eTh peaks, which are in general coincident with eK depletions (Fig. 3J,K). Additionally the levels of coupled, opposite eU—(eTh, eK) peaks are characterized by abundant graptolites noted in the facies record (Fig. 3E) and increased anhysteretic remanent magnetization (ARM) and isothermal remanent magnetization (IRM) values (Fig. 3M–N, see next section). The negative total gamma excursion begins at a depth of 858 m, slightly below the last appearance of the Neocucullograptus kozlowskii. The eK values (as a main component of the gamma spectrum) begin to decrease earlier, at a depth of 863 m, which is coeval with the beginning of initial increase of δ 13C values and appearance of the “varve-like” texture. The beginning of the pronounced negative gamma shift (depth interval 858–855 m) is related to the decrease of eK and eTh, only partly compensated by eU enrichment; while the next phase of the negative shift (855–850 m) is related to a uniform decrease in the content of all three measured elements. In spite of the uniform decrease, the relative proportion of each element in the total dose in the anomalous interval shows a relative domination of eTh and eK (Fig. 3J) in parallel with relatively scarce eU. The negative gamma ray excursion peak is coincident with the pronounced shift in δ 13C values (Fig. 3).

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initially very low, but gradually restored eU values. The eK values recover faster than eU, however, in the L11–L12 intervals decrease strongly again. Above 823 m, the patterns of total gamma ray intensity and individual element abundances are in general similar to these observed below the excursion, with distinct coupled high-eU, loweTh and eK peaks. It is important to note that in the depth interval 839–823 m the gradual decrease in eK values is parallel to the post-excursion gradual decrease of δ13C and immediately above (823–792 m) the eK fluctuations show moderate positive correlation (0.43) with the carbon isotope ratio trend. 4.4. Rock magnetism (by Katarzyna Sobień)

Fig. 7. A–C) Micrographs of neomorphosed limestones from the Mielnik IG-1 section interpreted as caused by sulphate reduction in anoxic conditions: A) equigranular, xenotopic texture with abundant bunch-like aggregates of (?framboidal) pyrite replacing amorphous organic matter; B) Neomorphosed upper part of graded grained beds with more abundant organic matter content or primary aragonite-rich composition; C) porphyrotopic texture in the bottom of limestone intercalation with preferential recrystallization in the base; note large pyrite replaced organic matter relict below.

In the upper part of the gamma ray excursion interval (850–823 m), total gamma ray intensity is gradually restored. In this interval the eTh values are relatively constant and anomalously low in respect to the

Laboratory measurements show (Fig. 3I) relatively low magnetic susceptibility (MS) values ranging from 10 to 99.7 (×10− 6 SI), with an average value of 72.6 (×10− 6 SI). The measurements performed directly along the core (Fig. 3I—in background) by portable equipment gave, in places, slightly different shape of magnetic susceptibility curve. In these places the magnetic susceptibility fluctuations, obtained using portable equipment, are similar to laboratory measurements of the anhysteretic (ARM) and isothermal (IRM) remanent magnetization (magnetic concentration parameters, Fig. 3M–N); hence they are probably to a significant degree influenced by the content of ferromagnetic minerals. In general, magnetic susceptibility fluctuations show some similarities with the gamma record. In particular they are parallel to the trends in eK values, what suggests that both parameters are driven by a common sedimentary process. The S-ratio expresses a relative contribution of lower-coercivity minerals to the total remanence (Bloemendal et al., 1992). It depends on the type of mineral, grain size and dispersion, but may also reflect variations in the coercivity characteristics of the magnetite component (Kruvier and Passier, 2001). A change in the S-ratio of a few percent can be caused by variations in the magnetic properties of the magnetites, such as the degree of oxidation or titanium substitution, while to cause a drop of several percent in the S-ratio values a substantial amount of hematite is required (Bloemendal et al., 1992). The S-ratio in the whole section (Fig. 3O) is close to 1 (0.91–0.95). It mirrors trends of magnetic susceptibility, reaching maximum values in intervals of low magnetic susceptibility, which is confirmed by a moderate negative correlation (−0.54) between both parameters. In general, there is no correlation between S-ratio, ARM and IRM parameters. In the lithological complexes L1–L2 the magnetic susceptibility values are stable, as well as the S-ratio (~0.925); magnetic concentration parameters (anhysteretic and isothermal remanence—Fig. 3M–N), show three positive excursions, coincident with coupled high-U–low-Th levels. In the succeeding L3–L4 interval magnetic susceptibility values gradually decrease. This decrease is coeval with an initial increase in δ13C values, a gradual decrease in total gamma ray intensity values and with the gradual domination of the “varve-like” texture of sediments. In the middle of the L3 complex anhysteretic and isothermal remanence parameters slightly decrease and they remain stable in the L4 interval. The S-ratio gradually decreases to 0.91 in the L3 complex, and just above the last appearance of the Neocucullograptus kozlowskii extinction it substantially increases in the L4 interval. At the bottom of the dolomitic siltstone (L5) complex (850.5 m) a distinct negative shift in magnetic susceptibility occurs, coinciding with the rapid increase in S-ratio, main carbon isotope shift and natural gamma-ray minimum. In the next 10 m of the L5 complex, magnetic susceptibility values remain low, which is coeval with maximal δ13C values. At the top of the L5 complex the S-ratio, together with ARM and IRM, reaches its maximum (843.5 m), Subsequently (L6), S-ratio values gradually decrease, parallel to pronounced ARM and IRM decrease and gradually restoring magnetic susceptibility values and maximum of the CIE. In fine claystones of the L6 complex, a local minimum of ARM and IRM values occurs, followed by an increase of both parameters throughout calcareous claystones and marls (L7–L9) up to a depth of

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Fig. 8. Stepwise acquisition of the isothermal remanent magnetization—IRM. Analyses made for 11 selected samples.

834 m. In the same interval, a relatively rapid restoration of MS values occurs, during which a pronounced decrease in δ13C values is noted. In the next interval (complexes L10–L11), magnetic susceptibility values and S-ratio are generally stable, parallel to pronounced gradual decrease of ARM and IRM parameters, with a minimum at the level of the first appearance of Pseudomonoclimacis latilobus–Slovinograptus balticus fauna, where the coupled high eU–low eTh levels, large-scale fluctuations in magnetic susceptibility and a substantial drop in the S-ratio appear. Above this level the petromagnetic parameters generally show stabile values, with a small negative anomaly in magnetic susceptibility values at the first appearance of the Monograptus acer–Ps. latilobus fauna; which is, however, not reflected in carbon isotope record. To evaluate magnetic mineralogy, additional IRM saturation experiments for 11 selected samples were carried out (Fig. 8). IRM acquisition curves show identical shape. 90% of saturation was achieved by all of the samples in a field of 0.3 T, which is typical for a mineral of low coercivity (magnetite) as the remanence carrier (in the ferromagnetic spectrum). Its presence is confirmed by a prominent positive correlation (0.77) between ARM and IRM parameters. Plots of MS versus ARM and IRM (not show) display poor interplay of these parameters, with correlation coefficients of 0.30 and 0.17 respectively. However, for samples with the lowest MS values (MS excursion interval), correlation coefficient of MS vs ARM and IRM is very high (0.91). 5. Interpretation 5.1. Sedimentary environment 5.1.1. General sedimentary environment of the Mielnik IG-1 succession and assumptions for interpretation of the facies record The general facies background of the studied interval represents offshore sediments, transitional between carbonate platform and clayey graptolitic shale facies. The limestone intercalations in claystone background represent a common type of sedimentary rock of the periplatform setting noted widely in the Silurian Baltic Basin (Aaloe and Jürgenson, 1977—Tables 3–4; Nestor and Einasto, 1977), dominated by sediments referred to here as detrital wackestone facies. One of the main observed facies changes in the studied succession is variation in the carbonate content of claystones. Based on a general facies model for the Silurian Baltic basin (Nestor and Einasto, 1977), it can be expected that the carbonate-poor intervals should have been deposited in relatively deeper sedimentary environments, while the carbonate-rich intervals record sea-level falls (e.g. Urbanek, 1997). Observations of modern periplatform environments suggest a more complicated model with sequence stratigraphic conditioning of carbonate supply (Schlager et al., 1994, Lantzsch et al., 2007; Jorry et al., 2010), which should be applied also in the case of the

studied succession. In this model the intervals rich in fine-grained carbonates result from highstand shedding, due to high carbonate production and progradation (Schlager et al., 1994) in stillstand conditions. During lowstand the platform tops are exposed and carbonate production ceased, while terrigenous material could have by-passed the platform and been supplied to the periplatform area, forming carbonate-poor intervals. However, in the case of prolonged relative low sea-level, an attached lowstand platform may be formed and shedding of carbonate material may also occur, resulting in similarities to the highstand-related, carbonate-rich record in periplatform setting. In this case the resedimented carbonate material may be expected to be coarser and contain abundant shallow-water grains (Lantzsch et al., 2007). During transgressive intervals, carbonate production and export depend on the rate of sea-level rise and on platform top elevation and relief. The proper sea-level rise causes very high carbonate production and export in the conditions of the “re-flooding window” (Jorry et al., 2010); whereas pronounced flooding causes platform drowning and cessation of carbonate production and export, resulting in a relative condensation of carbonate-poor sediments in the periplatform setting (Schlager et al., 1994). The early stage of the pronounced transgression will be additionally recorded by calciturbidites triggered by an increase of the hydrostatic pressure (Lantzsch et al., 2007). Sea-levels falls may also trigger calciturbidites, however, carbonate resedimentation ceases after the exposure of the platform top. According to these circumstances we interpret the abundance and type of the carbonate component in the sequence stratigraphic framework, with regard particularly to its spectral gamma record (see Section 5.2), as strictly dependent of relative condensation (eU) or terrigenous input (eK, eTh). 5.1.2. Excursion facies record—interpretation The presence of crushed and abraded dolomite grains, as well as vertically observed correlation between grain size and grain abundance of dolomite and quartz, along with the general rock texture, indicate a detrital origin for the dolomite grains in the “dolomite siltstone” facies. However, despite high detrital influx, the facies lacks of any signs of higher energy conditions. The presence of delicate reverse-graded discontinuous lamination indicates lack of reworking of material and excludes sediment homogenization due to bioturbation, which additionally suggests anoxic conditions in the sedimentary environment. The composition of the faunal assemblage remains unchanged; however, graptolites and nautiloids are relatively scarce. All grains in respective horizons are very well sorted and very fine-grained. All these features indicate sedimentation from suspension in dysoxic offshore environment. This conclusion implicates two possible models of detritus transport. 1) Transport in suspension in periplatform setting may occur due to density cascading (Wilson and Roberts, 1995). In this model, fine carbonate detritus is transported by hyperpycnal water flow downslope to its density compensation level, where it forms widely extending nepheloid layer with entrained fine detritus. Carbonate material may be suspended in platform top turbid waters due to storms and entrained into hyperpycnal water mass, which may form as a result of temperature changes or evaporation in wide shoal areas (carbonate platform interior). In our opinion does not aply to the features and the position of the dolomitic siltstone facies in the succession. Material derived from subaqueous, platform top environment should be more varied in its composition and relatively depleted in quartz and mica. Additionally, detrital limestone horizons would be expected to intercalate with significant pure pelite layers. Moreover, in contrast to siliciclastics, carbonate export from platform-top environment is expected during highstand conditions (Schlager et al., 1994; Wilson and Roberts, 1995), while dolomitic siltstone facies occurs during globally recognized

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regression (see below). It is also important to note that the grain assemblage of dolomitic siltstone occurs in the studied succession exclusively in the CIE interval and is not present within any calciturbidite intercalation, nor in the most proximal deposits of maximal progradation surface (see below), where redeposited shallow-water allochems are abundant. 2) During lowstand, a large part of the carbonate platform is emerged, and carbonate material may be transported by wind. The constant and specific grain assemblage (quartz silt, euhedral to rounded dolomite) and accessory grains such as: carbonaceous spherulites (often rounded), minute ooids (often found in recent and ancient carbonate eolianites—Loope and Abegg, 2001), abundant mica flakes, very small bioclasts; is consistent with the provenance of the material from the emerged platform top. Wind transport allows to by-pass the shallow water environment without admixture of a wide spectrum of neritic carbonate grains. Grain sizes and composition is very similar to recent dust deposits derived from arid regions such as the Great Salt Lake (Jones, 1953) or sebkhas and deserts surrounding the Persian Gulf (Al-Bakri et al., 1984). Dolomite crystals formed in the recent sebkha environment of the Persian Gulf show similar euhedral shape and often form spherulites (Bontognali et al., 2010). The Persian Gulf was an important source of wind-transported dolomite grains deposited in the Arabian Sea during the last glacial regressions—when the carbonate platform was emerged (Reichart et al., 1997), but also during arid climate historic events (Cullen et al., 2000). The detrital dolomite–quartz grain association is also abundant in the periplatform realm (both warm and cool) during the last glacial glacieustatic regressions (Eberli, 2000; Rendle et al., 2000; Betzler et al., 2005), however, their possibly aeolian origin has not been emphasized. Numerous amorphous, platy-shaped organic matter aggregates seem to be independent in size of the background grain size tendency; they are, however, distinctly more abundant in the dolomite–quartz coarser intervals (in the case of the “varve-like” mudstones—in the quartz laminae). This suggests a parautochthonous origin of these aggregates, forming during periodical phytoplankton blooms, probably triggered by periodical dust influx. The dolomitic siltstone facies are preceded and followed by claystones laminated with siltstones with “varve-like” texture, appearing at the same levels as the initial, minor δ13C shift (863 m). The facies contain similar graded discontinuous lamination, finegrained quartz silt and similar distribution of the amorphous, platyshaped organic matter aggregates. The lamination gradually becomes prevalent, directly coeval with a gradual increase in δ13C values; and at a depth of 853.5 m the silt material becomes rock-forming, followed by the appearance of dolomite (850.5 m) and the main shift in δ13C values. The texture and composition similarities may indicate a similar aeolian origin of this facies, most probably formed before and after of dolomite grain source was available (emersion of the platform top). The last exceptional facies in the CIE interval occurs in the L7–L9 complexes. The L7 complex contains ostracode wackestone facies with thin intercalations of graded grained grainstones, composed of peloids and bioclasts representing the most proximal calciturbidite intercalations in the studied interval. The L9 complex is composed of ostracode wackestone facies and marls with rich benthic fauna. The dominance of marly-wackestone texture of the fossiliferous sediments indicates autochthonous character of the fossils and suggests temporary colonization of the bottom. These two events, coeval with CIE decline, are probably caused by gradual shoaling due to normal regression and indicate the maximal progradation surface during sea-level lowstand (L7) and the re-flooding window (Jorry et al., 2010) at the beginning of the transgression (L9). 5.1.3. Facies-related point anomalies in the stable carbon isotope record The 1–2‰ lighter δ 13C and 1–2‰ heavier δ 18O (see Table 2) values of the partly neomorphosed detrital wackestone -facies, along with

87

the pyrite–organic matter association, indicate early-diagenetic sulphate reduction in anoxic conditions (compare Dickson et al., 2008) of some limestone intercalations. Pelitic limestones, adjacent to the fine-grained intercalations occur within and above the CIE maximum; they show heavier δ 13C (and δ 18O) values in relation to the marly background. The position of the interval at the end of progradational part of the sequence (see Section 5.3), suggests that the isotopically heavy admixture is supplied from more proximal facies belts. The local presence of complete neomorphism of allochthonous sediments may suggest a significant proportion of primary aragonite in their composition (compare Munnecke et al., 1997), which in part may be associated with the heavier signatures of platform-derived sediments (Swart and Eberli, 2005; Gischler et al., 2009). 5.2. General assumptions for the spectral gamma ray record interpretation in the Mielnik IG-1 section Gamma-ray spectrometry is an inexpensive and useful method for sequence stratigraphic interpretation, derived from the oil industry; however, it has recently been increasingly adopted in purely academic research studies. The method relies on the abundance of the three most important radioactive elements occurring in sedimentary rocks: K, Th and U, which are quantitatively distinguishable due to their different gamma-ray spectral ranges, and reported as radiometrically determined abundances (eK, eTh, eU). The elements have different geochemical behavior in sedimentary environments (Adams and Weaver, 1958) and various suits in rock-forming minerals (e.g. Hesselbo, 1996). K and U are regarded as environmentally mobile elements, while Th occurs in various insoluble compounds. Changes in the element ratios in the marine environment reflect variation in the clastic influx type and rate, resulting from changes in proximity to the sediment source and/or hinterland weathering type. Coupled sealevel–climate driving is common for the spectrum gamma ray (SGR) and sequence architecture records (Burgess, 2001). Hence, SGR may be used as a simple (non genetic) sequence stratigraphic tool (Davies and Elliott, 1996) and interpreted as a combined sea-level– palaeoclimatic proxy (e.g. Ruffell and Worden, 2000). The studied mid-Ludfordian event is regarded as a time of rapid global climate changes and sea-level fluctuations, hence the separation of palaeoclimatic and eustatic signals is the main difficulty in SGR interpretation. For this reason, the interpretation of the origin of the SGR record needs to be supported by chemostratigraphical correlation within sea-level change framework, outlined from shallow water settings (Eriksson and Calner, 2008; Kozłowski and Munnecke, 2010). The general pattern of K–Th–U abundance depends also on the sedimentary environment and lithology, because of the differential contribution and behavior of rock-forming minerals in different facies settings, and different sedimentary processes driving the mineral distribution in the profile. In the Mielnik IG-1 core, with clayeycarbonate lithologies the main rock-forming minerals are illite, smectite and mixed layers clays (Langier-Kuźniarowa, 1967). Macroscopic muscovite and biotite occur only in the limited interval of the dolomitic siltstone facies. Other K-rich minerals, such as feldspar and glauconite, are not recorded. Hence, the eK values may be simply interpreted as reflecting changes in illite content (locally + mica). The eK (illite) depletions, observed generally in parallel with finer-grained clayey lithology, are probably related to the smectite enrichment (confirmed by preliminary XRD results); when eK depletion occurs along with coarsening of the sediment (in the excursion interval), it may be interpreted as illite dilution mainly by quartz pelite and carbonates. In clayey shale lithologies, in an offshore setting, uranium, as a redox-sensitive trace element, shows in most cases good correlation with TOC (Lüning and Kolonic, 2003). Increased eU abundances often mark levels of sedimentary condensation, which often coincide

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with flooding surfaces (e.g. Junghans et al., 2003). However, in some cases (e.g. the Silurian of North Africa and the Middle East), organic matter accumulations may be associated with higher rates of burial of organic material, occurring at times of falling or low sea-level (Lüning et al., 2003; Loydell et al., 2009). In contrast to the lowstand TOC–eU enrichments, the increase in TOC and in eU on the flooding surface is often accompanied by low eTh values (Davies and McLean, 1998), resulting from reduced terrigenous input. Hence, the flooding surface sediments often record low eTh/eU ratio (Davies and Elliott, 1996). The main suits for Th are the thorium-bearing heavy minerals, occurring also as small inclusions in quartz grains. Detrital kaolinite is also regarded often as a Th carrier (e.g. Doveton, 1994; Hesselbo, 1996); however, it has not been confirmed by geochemical analysis (Hurst and Milodowski, 1996) and is probably a result of co-occurrence of kaolinite and heavy minerals, because of the relatively large size of kaolinite grains (see below in this section). In the case of the illite–smectite dominant facies, all Th-bearing minerals, being fine-grained and heavy, along with quartz (and eventually kaolinite) are concentrated in proximal facies belts, and hence eTh values in the studied sedimentary rocks may be treated as unequivocally related to the enhanced terrigenous influx. The input is a function of both: sea-level changes and aridity fluctuations in the hinterland area; it may, however, be expected that the climate changes should result in longer-lived trends in the eTh record, compared to sea-level fluctuations (see Ruffell and Worden, 2000). Although low eTh/eU ratio as an indicator of flooding surfaces has a certain limitations, in the case of the studied interval of the Mielnik IG-1 borehole it may be additionally controlled by mass-occurrences of graptolites. Most surfaces defined by coupled low-eTh high-eU values in the studied interval show relatively more abundant or massoccurrence of graptolites (Fig. 3). Surprisingly, the most pronounced (according to the eTh-eU pattern) surfaces are coincidental with first occurrences of a diversified graptolite fauna, whereas Pristiograptus dubius opportunistic fauna (Urbanek, 1997) flourished preferentially between coupled eTh-negative–eU-positive peaks. A eustatically controlled pattern of graptolite occurrence has been emphasized before (e.g. Bates and Kirk, 1984; Kaljo and Märss, 1991) and more recently by Goldman et al. (1999) and Egenhoff and Maletz (2007) from the Ordovician of USA and Sweden respectively. In the latter case, the authors recorded the first occurrences of diversified fauna on the maximal flooding surfaces, and ingressions of long-lived, opportunistic taxa during regressive intervals. The eustatically driven distribution of graptolite faunas may be caused by migration of the oxygen minimum zone, interpreted to have been occupied by specialized graptolite faunas living in the denitrification zone (see Finney et al., 2007). Coupled high-eU low-eTh peaks in the studied section are, in most cases, related to eK-lows. Because the flooding surface intervals are developed as very fine-grained carbonaceous–clayey rocks, we suppose that low eK values are related to dilution of detrital illite, most probably by smectite (with probably some contribution of carbonates). The distribution of clay minerals in continental margin seas depends on the differential size and shape of single crystallites [Kaolinite>Illite>>Smectite (K>I>>S)—Ruffell et al., 2002], differential flocculation (S> K> > I, Parham, 1966; Edzwald and O'Melia, 1975), size of single clay particles causing differences in their settling velocities (I> K> > S—Whitehouse et al., 1960; Gibbs, 1977) and very low density of smectites (I> K> > S—see Totten et al., 2002). In a simple ramp setting kaolinite-illite–smectite fractionation in a seaward direction may be expected (Thiry, 2000), as is observed in modern estuaries (e.g. Edzwald and O'Melia, 1975), deltas (e.g. Gibbs, 1977) and marginal seas (e.g., Ehrmann et al., 2007). In the deep shelf, a clay-dominated environment, the landward facies shift during transgression may cause a decrease of illite/smectite ratio (Holmes, 1987; Daoudi et al., 2008), due to slow settling and authigenic

formation of smectite in condensed surfaces (Thiry, 2000 and references therein). This mineralogical change may be recorded as a decrease of eK values. It is important to note also that the dilution of an illite-dominated lithology by other minerals, such as pelitic quartz or kaolinite, may lead a similar eK record in the case of a regressive event; however this case may be easily distinguished by the presence of parallel eTh enrichment, eU depletion and micro- and macro-facies observations. Due to the presumed similar behavior of eTh and eK in the sedimentary environment of the Mielnik IG-1 Ludfordian succession, resulting predominately from the intensity of detrital influx, we use combined eTh × eK coefficient as a relative measure of detrital input (i.e. proximity) and eTh × eK/eU coefficient as a proxy for relative rate of sedimentation. 5.3. Sequence stratigraphic interpretation The lower part of the studied interval (L1–L2), composed of graptolitic claystones with detrital wackestone facies intercalations (partly neomorphosed), shows generally a simple spectral gamma ray record with flooding surfaces marked by coupled eU-enriched– eTh depleted levels. This pattern allows us to distinguish four parasequences within the Neocucullograptus kozlowskii Zone. The base of the first is coeval with the appearance of N. kozlowskii (Urbanek, 1970), and the next ones are additionally marked in the core by the occurrence of abundant graptolites. Increasing upward eK × eTh values within the first two parasequences and relative increase in total gamma intensity suggest progradational parasequence stacking; hence we interpreted the L1 complex as a part of a highstand (Fig. 3P). The subsequent two parasequences (L2 complex) are coeval with decreased carbonate content and presence of the “varve -like” lamination containing abundant quartz pelite, interpreted here as wind-derived. Relative eTh values and eK × eTh/eU ratio are lower, which, along with the low carbonate content may suggest transgressive conditions up to depth of 861 m. At a depth of 858 m, in the maximal progradation surface of the 4th parasequence in the kozlowskii Zone, intercalations of redeposited sediments marks first sign of regressive conditions. However, subsequently above, increased abundance of graptolites and nautiloids, up to a distinct condensation surface at a depth of 855 m (consisting of the last mass occurrence of Bohemograptus bohemicus, scarce N. kozlowskii, and intercalations of pyrite-bearing, neomorphosed detrital wackestone facies), may be interpreted as a distinct flooding surface. On the other hand, the subsequent HST is missed and the condensed interval contains intercalations of cross laminated silt layers, with coarse quartz grains (Fig. 5A). Hence, the second possibility of interpretation (preferred here) of the L3 complex is as a short lowstand step with emersion of the carbonate platform top and cessation of highstand shedding leading to relative condensation with enhanced supply of terrigenous material. The overlying complex L4 (composed of calcareous mudstones without limestone intercalations), with regard to the eK× eTh/eU ratio (Fig. 3), represents falling-stage systems tract, with beginning coeval with the kozlowskii extinction level. This coincidence may be caused by a eustatically driven disappearance of the denitrification zone in the shelf area (Finney et al., 2007). The following sequence boundary, is interpreted at a depth of 847.5 m—within the dolomitic siltstone complex (L5), according to maximal eTh × eK/eU values and reversion of the total-gamma intensity decreasing trend (beginning of progradation). The next interval (up to a depth of 834 m) shows gradual eTh × eK increase which indicates progradation in lowstand conditions. Dolomitic siltstone facies passing upward into calcareous claystones, interbedded with ostracode wackestones, containing the first occurrence of benthic fauna fossils. At the depth 840.5 occurs distinct flooding event marked by mass occurrence of Pristiograptus dubius in lower part of the L7 complex and distinct negative excursion in eTh and ARM–IRM proxies. The upper part of the complex, above the

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flooding surface contain detrital packstone to grainstone (suggesting enhanced winnowing), containing shallow water grains (peloids, crinoids—depth 839 m), which indicate maximum progradation in regressive conditions. The L9 complex (339–340 boxes) contains abundant trilobites (Scotiella, Leonaspis), small brachiopods and crinoid detritus occurring in a carbonate-bearing, marly matrix; which suggests intense shedding of fine carbonate material at the beginning of transgression and represent the re-flooding window. The progradational conditions, after flooding marked by carbonate-poor L8 complex, are additionally marked by heavy mineral proxies, i.e. high (local maximum) eTh and ARM–IRM values. Above a depth of 834 m the increase of eU values in parallel with the decrease of the eK-content (interpreted as a decrease of illite/smectite ratio) and gradual domination of fine clayey lithology is interpreted as indicating a transgression. The presence of intercalations of the graded grained grainstone facies in the lower part of the interval may reflect upper slope instability due to beginning of the transgression, while the intercalation of the wackestone with biomorpha facies above indicates condensation events due to backstepping and drowning, which is additionally confirmed by carbonate depletion in this interval. The subsequent maximal flooding surface interval (L12 complex) is marked by three distinct condensation surfaces (high-eU–low-eTh), low eK values and intercalations of neomorphosed detrital wackestone– limestones (suggesting anoxic conditions) along with an ingression and mass occurrences of graptolites of the Pseudomonoclimacis latilobus– Slovinograptus balticus assemblage (Urbanek, 1997). The highstand progradation that follows is recorded by an increase in the eK and eTh content, contemporary with a strong decrease in eU (best seen in the eK × eTh/eU curve—Fig. 3L), coeval also with a Pristiograptus dubius faunal bloom (Urbanek, 1997), probably resulting from shoaling conditions. The following sequence boundary, due to a lack of distinct facies changes and indistinct record in the spectral gamma ray data, is interpreted as higher-order one (type 3—Schlager, 1999). Above the highstand, the decrease in eK and eTh content, along with gradual enrichment in eU and carbonate scarcity in the L14 interval, indicates transgression culminating at a distinct maximal flooding surface marked at the depth of the 793.5 m by high eU– low coupled eK and eTh peaks, intercalations of neomorphosed detrital wackestone-facies and the coeval first appearance of the Uncinatograptus acer with spot reappearance of Pseudomonoclimacis latilobus (see Urbanek, 1997). The succeeding carbonate bearing complex (790–793 m), suggests highstand shedding conditions and is followed by a carbonate-poor interval with distinctly increased eK and eTh values and point anomalies in anhysteretic remanent magnetization (ARM) and S-ratio (see Section 5.4), which are interpreted as a minor forced regression event. The subsequent interval with three distinct flooding surfaces (recorded by high eU– low eTh levels), begins the subsequent transgression.

induced variations within the lithogenic fraction of the sediment (Capotondi and Vigliotti, 1999). A similar situation is present in settings with low clastic influx, such as carbonate platforms (Hladil et al., 2006) or periplatform areas. In general, periplatform sediments contain variable admixtures of quartz, clay and carbonate, the composition of which may dominate the petromagnetic signal. Minor contribution of ferromagnetic minerals in this setting may be syngenetic or detrital (both aeolian and riverine) in origin. In the case of the studied section, moderately low magnetic susceptibility, together with a lack of correlation with anhysteretic remanent magnetization (ARM), and moderate positive correlation with eK measured in the core (0.34) imply that a major contribution to the magnetic susceptibility is probably due to paramagnetic illite (Hrouda and Kahan, 1991). According to Ellwood et al. (2000), in the case of rocks composed mainly of diamagnetic minerals, even a very small amount of paramagnetic minerals can elevate their MS. In our studies, one of the most important factors affecting MS values is dilution of illite by weaker paramagnetic and diamagnetic minerals such as quartz pelite, chlorite (English, 1999) and carbonates. The poor relationship between magnetic susceptibility (MS) values and anhysteretic–isothermal remanent magnetization (ARM–IRM) indicates on only accessorial contribution of ferromagnetic admixtures to the magnetic susceptibility signal. However, in the MS negative excursion interval this contribution is significant, which is indicated by distinct coherence observed between the parameters. The baseline ferromagnetic signal, characterized by very high correlation between the ARM and IRM values, suggests strong domination of magnetite in the ferromagnetic spectrum of studied rocks. Fluctuations of magnetite content, observed as overall trends in ARM and IRM values, are probably an interplay resulting from the presence of two independent populations of magnetite grains, reflected by a distinct change of the S-ratio in the anomalous interval. The general range of S-ratio (0.91–0.95) suggests a “homogenous” soft magnetic mineralogy. The relatively lower S-ratio values (around 0.92) in the pre- and post-lowstand part of the section represent finer magnetite populations. The ARM–IRM positive peaks on the flooding surface horizons in this interval (Fig. 3), indicate that the grains are preferentially concentrated on condensation surfaces, which suggests “biogenic fine”, or aeolian-detrital origin of this magnetite grain populations. This signal is disturbed in the excursion interval, showing reduced coercivities (around 0.945), interpreted as significant admixture of the coarser magnetite grains (>10–15 μm) (e.g. Evans and Heller, 2003). The main contradiction is in the coeval decrease of the ferromagnetic content (low ARM–IRM values) and appearance of coarser fraction. This suggests radical changes inside the clastic material delivery system.

5.4. Rock magnetism record (by Katarzyna Sobień)

The record of sea-level changes in a deep shelf setting is not so clearly recorded as in the case of the littoral zone, where a change of a few to several dozen metres results in distinct facies shifts or emersion. However, in a shallow water environment, a sea level drop results in gaps, erosion of older sediments, redeposition and/or rapid infilling events, hence the stratigraphical record of these setting is intermittent and often limited to transgressive intervals. Additionally, the rate of deposition in nearshore environments often changes rapidly, hence thickness–time relations are difficult to interpret. Deep shelf facies contain continuous record, with more stable rate of sedimentation, but eustatic changes are not so clearly reflected in deep, up to several dozen meters, water setting, as in the littoral zone. The interpretation of sea-level changes is one of crucial issues in the discussion about links between the carbon isotope record and eustacy during the mid-Ludfordian CIE (see sequence stratigraphic frameworks

Changes in the magnetic mineralogy of sediments can be caused by palaeoceanographic (e.g. sea-level) and palaeoclimatic variations, which drive changes in the riverine influx, source of detrital input, carbonate content, aeolian dust concentration and oxic–anoxic conditions. In many cases (e.g. Ellwood et al., 2000, 2001) changes in magnetic susceptibility (MS) is interpreted as induced by sea level fluctuations, where rises in MS values result from maximum magnetic (detrital) input due to intensive erosion during sea level fall. In nearshore areas, sea level changes control synchronous riverine influx changes—due to base level changes, as well as aeolian detritus influx—by changes of distance from the land areas. However, in deep-sea sediments, changes in the concentration, mineralogy and grain size of the magnetic fraction reflect instead the climatically

6. Discussion 6.1. Sea-level changes during the mid-Ludfordian CIE event and eustatic correlation of the Mielnik IG-1 section

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and discussions in: Eriksson and Calner, 2008; Kozłowski and Munnecke, 2010; Munnecke et al., 2010; Loydell and Frýda, 2011). The detailed relation of eustacy to other environmental changes during the event is crucial for understanding its causes. There is a general consensus regarding the co-occurrence of a global sea-level fall close to the mid-Ludfordian CIE (reviewed in Munnecke et al., 2010). The position of the sea-level drop, preceding the main carbon isotope shift, marked by distinct emersion-related surfaces with epikarst in the base of Eke Formation on Gotland Cherns, 1982), seems to be similar in the Holy Cross Mts (Kozłowski and Munnecke, 2010), in Podolia (Kaljo et al., 2007), the Barrandian (Lehnert et al., 2007b; Manda et al., 2012) and Australia (Jeppsson et al., 2007). It is important to note that the falling stage sedimentary rock are often preserved below the sequence boundary (e.g. Botvide Member in E Gotland—Eriksson and Calner, 2008; Jadowniki Member in the Holy Cross Mountains— Kozłowski and Munnecke, 2010; Isakivtsy Formation in Podolia—Kaljo et al., 2007), which may suggest the common presence of ramp relief before the sea-level fall, that may be caused by a poorly expressed highstand system tract (HST—normally with shoaling and normal regression). The Mielnik IG-1 succession records the first signs of regression in the upper part of the kozlowskii Zone (at the bottom of the L3 complex). The probably gradual shallowing (Fig. 3R) caused firstly carbonate scarcity with levels of the relative condensation in periplatform setting. This event (L3 complex) may be correlated with the Dayia flags in Gotland (Samtleben et al., 2000), where also typical HST is absent. Rapid transition from this relative stillstand to regressive conditions is recorded as rapid shallowing at the base of Botvide Member interpreted in E Gotland as a falling stage or debatable highstand system tract (see discussion in Eriksson and Calner, 2008). In this position sedimentary rocks representing falling stage systems tract (FSST) are present also in the Mielnik IG-1 section (complex L4, Fig. 3P). In both cases the initial gradual increase of δ13C values (Samtleben et al., 2000) are consistent with sea-level fall and the main shift in carbon isotope ratios seems to be almost coincident with the sequence boundary (SB). Above the SB in the carbonate platform margin of Gotland transgressive rocks of the Eke Formation occur (TST-HST in Eriksson and Calner, 2008). The coeval erosive gap in the platform interior (e.g. Ohesaare, Estonia—Kaljo et al., 2007) indicates that the inner platform (ramp) was simultaneously exposed. The offshore successions (Mielnik IG-1—this study, Vidukle— Martma et al., 2005) record gradual shoaling which suggests sea-level lowstand (lowstand system tract—LST), however the both sections contain internal flooding surfaces. This situation may be explained by prolonged lowstand conditions combined with prograding subsidence and slow sea-level rise (flooding events of Kozłowski and Munnecke, 2010), which would have created accommodation space in the outer inclined platform areas under incipient drowning conditions. Hence, sediments deposited in the outer parts of the carbonate platform (e.g. Gotland) during relative sea-level lowstand, formed relatively thin, shallow-water TST-HST parts of the succession(s) above the type-2 (sensu Vail and Todd, 1981) sequence boundary. The next correlative event is the maximal progradation surface (end of relative sea-level lowstand) which is recorded as normal regressive rock. The regression (second forced regression in Eriksson and Calner, 2008) occurring near the end of the CIE, is noted in most of described profiles such as: Gotland (oolites—Samtleben et al., 2000), Australia (oolites—Jeppsson et al., 2007), the Vidukle borehole in Lithuania (shallowest faunal assemblage in the uppermost Mituva Fm.—Martma et al., 2005), Holy Cross Mountains (lagoonal facies and maximal progradation of red muddy sediments— Kozłowski, 2008; Kozłowski and Munnecke, 2010), and also in the Mielnik IG-1 borehole (appearance of benthic fauna and lowstand to flooding window carbonate shedding in the L9 complex). The formation of a lowstand platform—superposed/lying on the flank of the older offshore ramp during a prolonged relative sealevel fall, is consistent with the occurrence of abundant coated grains (including ooids) in the CIE interval (compare with Schlager et al., 1994).

In the present interpretation the first main transgressive (prominent relative sea-level rise) event is related to the end of the interval of maximum values of the CIE (see also Eriksson and Calner, 2008; Kozłowski and Munnecke, 2010). The first maximum flooding surface (coeval with the first appearance datum (FAD) of the Pseudomonoclimacis latilobus–Slovinograptus balticus assemblage) is marked by a negative peak in the CIE. According to sequence stratigraphic correlation at the first post-regressive MFS level, low values in periplatform profiles— Mielnik IG-1 (+0.1‰—this study) and Vidukle (−0.85‰—Martma et al., 2005), are coeval with a relatively weaker decrease observed in the case of the Gotland (Burgsvik-Hamra transition; +4.06‰) and Rzepin section (Bełcz-Słupianka transition +4.01‰—Kozłowski and Munnecke, 2010), hence the CIE decline probably differs strongly between periplatform areas and the platform interior, possibly reflecting bathymetric gradient of the δ13C values. After the MFS, sea-level stillstand conditions (HST) led to slight regression. The slight increase of δ13C values in the Mielnik IG-1 (+1.11‰) and Vidukle (+1.56‰) cores, along with a common shallowing trend (this study; Martma et al., 2005), allow correlation of the upper part of complex L13 in Mielnik IG-1, with the lower part of the Ventspils Formation in Lithuania. Correlation between sea-level fluctuations (recorded by the illite/ smectite ratio reflected in eK values) and δ13C values, in the interval of δ13C decline in the Mielnik IG-1 succession up to the FAD of Uncinatograptus acer, implies the presence of a bathymetric gradient of δ13C values in the basin, as recognized earlier in the case of the other Silurian CIEs (see Loydell, 2007). In this interval, due to gradual sealevel rise, regeneration of the carbonate ramp relief probably occurred, which may have caused a strong bathymetry-related δ13C gradient across the ramp (see Gischler and Lomando, 2005; Gischler et al., 2009). Hence, according to our sequence stratigraphical correlation, the pronounced differences in the maximal δ13C values between the platform and periplatform setting are noted in the CIE expire—HST interval (Gotland, lower Hamra +6.86‰—Samtleben et al., 2000; Rzepin, lower Słupianka +5.72‰—Kozłowski and Munnecke, 2010; Mielnik +1.2‰— this study; Vidukle—rise to +1.56‰ in Ventspils Fm.—Martma et al., 2005). The next maximal flooding surface (MFS) distinguished in the Mielnik IG-1 section is coincident with the FAD of the Uncinatograptus acer. The end of δ13C fluctuations in Mielnik IG-1 section indicates that this level corresponds to the CIE decline in the inner carbonate platform setting, hence it can be referred to the MFS interval in the top part of the Słupianka Member in the Rzepin section (Kozłowski and Munnecke, 2010), as a first post-excursion MFS. Subsequent stronger (?forced) regression (?FRST/HST) corresponds probably to the Bronkowice sandstone and oolites in the Rzepin section (Kozłowski and Munnecke, 2010) and may be related to the late Pseudomonoclimacis latilobus Zone shallowing event in Bohemia (Manda and Kříž, 2006). The subsequent pronounced transgression in the top part of studied interval in the Mielnik IG-1 section corresponds probably to the distinct flooding event in the bottom of the Sarnia Zwola Formation (Kozłowski and Munnecke, 2010) and appearance of “faro-type” reefs in the Sundre Formation of Gotland (Samtleben et al., 2000). It is important to note that blooms of Pristograptus dubius are located in stillstand (lowstand or highstand) regressive deposits, while diversified graptolite faunas preferentially appear in MFS intervals (compare with Urbanek, 1997). This may suggest that well dated biostratigraphical levels in successions representing the shelf interior may correspond only to the flooding events occurring within part of the total range of the index graptolite species. For example, according to the presented sequence stratigraphic framework and Urbanek's (1997) data, Ps. latilobus in the Mielnik IG-1 profile (Fig. 2) appeared in two consecutive MFS levels. 6.2. Magnetic susceptibility and natural gamma ray excursions The Mielnik IG-1 succession records distinct magnetic susceptibility (MS) and natural gamma (NG) negative anomalies coincident with the

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CIE (Fig. 3H–I). Analysis of these events indicates that the signals have a similar source, resulting from the dilution effect of diamagnetic, lowradioactivity mineral(s). Both of the excursions begin in the middle part of the Neocucullograptus kozlowskii Zone, with ingression of “varve -like”, carbonate-pure mudstone (Fig. 3H–I). The ingression of dolomitic siltstone facies is coeval with another rapid negative shift in MS values, however the NG minimum precedes the dolomite siltstone interval; probably because of abundant (K-bearing) mica observed in this facies, which counter-balances the illite dilution. These observations, along with microfacies analysis, indicate that the main diluting minerals in the excursion interval are very fine-grained quartz and dolomite. As the observed negative excursions are present in the sea-level fall to lowstand interval, the supply of terrigenous material may be caused by enhanced riverine influx due to sea-level fall, however the dolomite–quartz grain association suggest rather aeolian route of transport. Because wind-related mineralogical signal during the sea-level fall is expected to be increased by the basinward shift of the shoreline, the consequent mineralogical record (dilution effect), observed in the studied section, should be treated as caused by a tight interplay between eustacy and presumably enhanced gustiness, and therefore an unequivocal separation of the two signals in the MS and natural gamma records is very difficult. We suppose that the quantification of aeolian input in the studied section may be reflected in the S-ratio (Fig. 3O), which in the studied case probably reflected the average size of magnetite grains (see Section 5.4). S-ratio values in the studied interval show some independence of the sequence stratigraphic framework and, in places, mirror trends expected from other clastic influx proxies. The observed excursion in S-ratio values is very similar to and strictly coeval with a rapid increase of abundance of the terrestrial spores in this interval on Gotland, interpreted by Stricanne et al. (2006) also as due to aeolian influx (see also Loydell, 2007). Similarly, as in the case of Gotland, the S-ratio excursion does not follow reconstructed sea level fluctuations; particularly, it is not influenced by sequence boundary (coincident with the transgression recorded in the Eke Formation on Gotland); or the post-CIE-event transgressive surface at the Burgsvik-Hamra transition. The petromagnetic record of the studied interval in the Mielnik IG-1 section shows prominent similarities with the record of enhanced aeolian influx in low-latitudes during the Pleistocene. In the sediments of the South China Sea (Kissel et al., 2003) periods of enhanced winter monsoon activity during glacial intervals are reflected by relatively lower magnetic susceptibility, depletion in magnetic grains (low ARM and IRM values) and reduced coercivities indicating the simultaneous appearance of the coarser magnetic fraction, interpreted as being derived by strong winds.

6.3. Record of the mid-Ludfordian CIE in Mielnik IG-1 in relation to other sections 6.3.1. CIE amplitude The maximal δ13C values are in general reported from shallow carbonate facies deposited in foreland basin type tectonic setting, with a very high rate of subsidence (i.e. Scania +11.2‰—WigforssLange, 1999; Gotland +10.54‰, Australia, Broken River Region— +9.67‰—Jeppsson et al., 2007; Holy Cross Mountains, Poland +8.9‰—Kozłowski and Munnecke, 2010). In shallow carbonate platform areas, characterized by weaker subsidence regime, maximal δ13C values are probably omitted, because of gaps, which is consistent with the occurrence of the CIE during sea-level lowstand (i.e. Podolia, Isakivtsy–Prigorodok transition (probably emersion)—+6.9‰ (Kaljo et al., 2007); Estonia, Ohesaare (emersion and erosion, gap)—lack of the CIE record (Kaljo et al., 1997; Kaljo and Martma, 2006); Czech Republic, Požary and Mušlovka Quarry (SB, paleokarst)—max. ~+3‰

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and +4.6‰ respectively (Lehnert et al., 2007b), USA (SB) up to +5.17‰ in the distal platform setting—Barrick et al., 2010). Proximal deep shelf facies with a complete record show slightly lower maximal δ13C values (e.g. Kosov section, Bohemia 8.0% Lehnert et al., 2007b; Vidukle borehole, Lithuania 8.17%—Martma et al., 2005) in relation to the maximal values noted in coupled shallowmarine–high subsidence areas. In comparison, the peri-platform Mielnik IG-1 section, with its supposedly complete succession and high subsidence rate, shows lower maximal δ 13C values (+6.55‰, +6.74‰) in autochthonous and allochthonous sedimentary rock respectively, that correspond to the maximal δ13C values reported from the deep shelf (probably also complete) succession of the Priekule borehole (+5.9‰—Kaljo et al., 1997). An important feature is the different isotope signature of autochthonous and allochthonous sediments in the Mielnik IG-1 section, noted in the uppermost part and decline of the CIE. In the present interpretation the redeposited rocks represent material of the same age derived from shallower facies belt, which indicates a facies/bathymetric-related carbon isotope gradient in the basin. Both observations confirm that the amplitudes of the mid-Ludfordian CIE are probably higher in shallow than in the deeper part of the basin (see Munnecke et al., 2003; Loydell, 2007). 6.3.2. CIE beginning in Mielnik IG-1 is recorded as a slight, gradual increase of carbon isotope ratio values noted in the upper part of the Neocucullograptus kozlowskii Zone, similar to its position in Barandien area (Lehnert et al., 2007b; Manda et al., 2012). The kozlowskii graptolite extinction event, well defined in the section, occurs, similarly as the conodont extinction in Gotland and Australia, in the interval preceding the main CIE shift (Jeppsson et al., 2007); however, the detailed relation of the graptolite and conodont extinction is uncertain due to the lack of conodont data in the studied interval of the Mielnik IG-1 section. Single specimens of Bohemograptus bohemicus have been observed immediately below the main CIE shift, i.e. probably above the conodont extinction level. 6.3.3. CIE main shift The mid-Ludfordian CIE maximum starts with a very rapid shift in δ13C values occurring in the Mielnik IG-1 section in the lower part of the Neocucullograptus kozlowski–Pseudomonoclimacis latilobus interzone, characterized by scarce occurrences of Pristiograptus dubius. The coincidence of the pronounced carbon isotope shift with a sequence boundary in shallow-water settings may be an artefact of the incomplete sedimentary record or condensation due to winnowing, hence the duration of the shift (600 years in Jeppsson et al., 2007) is difficult to assess and its rapidity may be illusory. The Mielnik IG-1 section lacks signs of erosion in this interval; succession reveals a high overall rate of sedimentation (1 m in ca. 9500 years on average) and its periplatform position, together with sedimentary record of forced regression, suggest even an increase of this value in the CIE interval. Coring in the interval is almost complete (4.5 m of core in 5 m long core interval) and the rapid δ13C shift is documented in a single box. Hence the record confirms an abrupt shift in carbon isotope values (Jeppsson et al., 2007), [(2.7‰ in 0.4 m—according to the average thickness–time relation in the Ludfordian part of the section −less than 4000 years)]; in the studied case—strictly coincident with the abrupt appearance of dolomitic siltstone facies and slightly preceding the sequence boundary (see also Martma et al., 2005). 6.3.4. The CIE maximum is represented by a 13 m thick interval (according to the average rate of deposition: maximum ~ 120,000 years) of constantly high carbon isotope ratios, coeval with: progradation (here: sea-level lowstand); high carbonate content in spite of the low base level (here interplay of relative aridity, lowstand shedding and

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antiestuarine–evaporation triggered circulation—Bickert et al., 1997); high admixture of silt-sized dolomite and quartz (here: aeolian influx). These features seem to be identical to the characteristics of the CIE maximum interval in Gotland (Jeppsson et al., 2007; Eriksson and Calner, 2008). 6.3.5. The CIE decline is coeval with transgression, as in Gotland (Eriksson and Calner, 2008). Moreover in the studied section, more rapid decrease of δ13C values is noted in the carbonate-poor transgressive (according to spectral gamma-ray data) intervals, which suggests platform drowning as the important driving mechanism of the CIE decline. Correlation of carbon isotope ratios with eK values is observed up to 792 m. and above this depth δ13C values are constant despite fluctuations in eK values. Because the eK values depend on the proximity to the land (indirectly: bathymetry), this correlation indicates the presence of a δ13C paleobathymetric gradient and continuation of the CIE in the shallower part of the basin up to the end of this correlation at the first appearance datum of Uncinatograptus acer and the related maximal flooding surface. Hence the drop in carbon isotope values seems to be rapid in the basin (Martma et al., 2005; this study), while a more gradual decrease is recorded in the carbonate platform interior (Gotland - Samtleben et al., 2000; Podolia – Kaljo et al., 2007). 6.4. The significance of the dolomite siltstone facies in the discussion on climate changes during the mid-Ludfordian CIE The rapid regression at the beginning of the mid-Ludfordian CIE is regarded as glacially induced (e.g. Eriksson and Calner, 2008), which has been recently supported by the record of climate cooling at this time (Lehnert et al., 2007a). These global climate changes are probably also reflected by at least local, low-latitude aridity (Samtleben et al., 2000). Climate aridity may be reflected by high carbonate content in the CIE interval in the studied section, occurring despite base-level fall due to forced regression, which should have caused reduced carbonate shedding parallel with enhanced terrigenous input. On the other hand, the main carbonate component in the CIE interval is detrital dolomite, interpreted as wind-derived. The dolomite–quartz terrigenous input caused changes in the average mineral composition in the basin and significantly influenced natural total gamma and MS record in the CIE interval. The detrital dolomite input is also noted in the studied interval in other locations. Lehnert et al. (2007b) mentioned abundant “dispersed dolomite” in the pre-regression interval of the Požary section (Czech Republic). The similar dolomite is present in redeposited carbonate clasts in distal Všeradice section, indicating its (at most) early diagenetic origin (Manda et al., 2012). Samtleben et al (2000) noted up to 50% of quartz, detrital dolomite and mica in the Dayia flags of Gotland (correlated with “varve-like” mudstones, and initial increase in δ 13C values in Mielnik IG-1 section). The Daya flags show distinct parallel lamination with aligned, well-sorted grains, lacking signs of hydrodynamic transport, what had led Samtleben et al. (2000, p. 6) to favor aeolian transport of its terrigenous components. Eriksson and Calner (2008) described very fine-grained, beige, laminated dolomitic fissile mudstone from the Botvide Member in Gotland, which also contain minute dolomite crystals. Aeolian transport of the terrigenous material in Gotland has been suggested by Stricanne et al. (2006) based on distinct enrichment in terrestrial palynomorphs observed in various facies spectrum. Loydell (2007 p. 542) suggested that this feature may be simply explained by sea-level lowstand and proximity to land, however, Stricanne et al. (2006) rejected this interpretation because the variations of spore abundances did not reflect sea-level fluctuations in the CIE interval (in both Eriksson and Calner, 2008 or Kozłowski and Munnecke, 2010 interpretations). A similar lack of correlation between the terrigenous influx and sealevel fluctuations is noted in the Mielnik IG-1 section. The dolomitic

siltstone facies appears at the end of the sea-level fall and persists above the sequence boundary, in the interval (based on/according to isotope correlation) of gradual reflooding noted in Gotland (Eriksson and Calner, 2008) and in the Holy Cross Mountains (Kozłowski and Munnecke, 2010). The facies also suddenly disappear before the maximal progradation surface of the lowstand interval. The S-ratio, probably reflecting the average diameter of detrital magnetite, shows a similar lack of coincidence with eustacy. The general facies development of the studied event in the Mielnik IG-1 succession shows also distinct similarities with the Homerian Mulde event. In the south Baltic region “varve -like” microlamination occurs in the lundgreni-testis zones (Paskevicius, 1986), or immediately above in the lower part of dubius Interzone (Porębska et al., 2004). In the deep shelf succession, Ulst (1974) noted a 1 m-thick laminated dolomite below the parvus Zone (Ancia Member of the Riga Formation, Latvia). In the studied section, the appearance of significant terrigenous influx (triggered by the sea-level alone or the sea-level plus climate), manifested by magnetic susceptibility and gamma ray dilution effects, is parallel to the slight increase in δ 13C values (pre-excursion step). The rapid appearance of detrital dolomite facies, interpreted as aeolian, is isochronous with the main shift in δ13C values. This suggests that the CIE and the terrigenous input has a common driving mechanism, or the CIE is triggered by terrigenous (aeolian) influx. The simplest common driving mechanism is sea-level changes (regression), leading to emersion of the carbonate platform top and creating a proximal (dolomite bearing) source of wind-derived material (eustatically enhanced dustiness). However, the very rapid appearance of the dolomite siltstone facies, and lack of matching of the facies occurrence with the detailed sequence stratigraphic framework (see also Stricanne et al., 2006) may indicate global climate changes, consisting particularly in the global glacigenic gustiness resulting in an enhanced dustiness (McGee et al., 2010), which would coincidence with platform top emersion. Regardless of the scenario, enhanced dustiness can be a link between the climate-eustacy and the carbon isotope excursion in the case of the studied event. 6.5. Causes of the mid-Ludfordian CIE—globally driven, locally induced? According to classic models (e.g. Kump and Arthur, 1999), positive carbon isotope excursions are often linked with perturbations in the global carbon cycle, particularly with enhanced bioproductivity and enhanced burial of 12C-rich organic matter in the deep-water setting. In the case of the studied event these models are very difficult to apply (e.g. see discussion in Munnecke et al., 2003 and Stricanne et al., 2006). Enhanced black shale deposition in deep-water settings has not been recognised and available palynological data from Gotland are interpreted by Stricanne et al. (2006) as reflecting decreased bioproductivity in this time interval. On the one hand, some features suggest local environmental isotope fractionation: the excursion is very strong and in the light of presented data it is, at least in some regions, very rapid. The amplitude of the event seems to depend on bathymetry. The CIE termination seems to be more rapid in deeper shelf facies and more gradual in platform interiors. On the other hand, the discussed CIE is correlable between different palaeocontinents (Jeppsson et al., 2007; Lehnert et al., 2007b). All these features indicate a global driving mechanism of the environmental changes, which induces in distant places similar local carbon isotope fractionation mechanisms (see Holmden et al., 1998). Up to day, eustacy seems to be one of the main drivers of the mid-Ludfordian (and probably some other Phanerozoic) CIEs (Loydell, 2007; Eriksson and Calner, 2008; Kozłowski and Munnecke, 2010). Lowstand conditions may cause supply of dissolved isotopically heavy carbonates due to carbonate platform weathering (Melchin and Holmden, 2006; Loydell, 2007); however, the presence of a pronounced mid-Ludfordian CIE in of the Holy Cross Mountains (Kozłowski and

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Munnecke, 2010), without carbonates or emerged carbonate platform in its hinterland (Kozłowski, 2008), does not support the dominant importance of this mechanism in the case of the studied CIE. In our opinion, the most important effect of the sea-level fall is remodelling of the proximal part of the shelf by seaward relocation of the carbonate factory (see eg. Pomar, 1991; Schlager, 1992). In the case of subsequent late lowstand, the initial flooding on the previously exposed platform surface (due to e.g. subsidence or higher order sealevel fluctuation) is recorded as transgressive, but often very shallow water or even non-marine facies (eg. Podolia—Kaljo et al., 2007; Gotland—Eriksson and Calner, 2008; Holy Cross Mountains—Kozłowski, 2008; Kozłowski and Munnecke, 2010). The flooding on inherited and newly attached relief causes spreading of shallow-water carbonate platform environments and subsequent extension due to progradation by carbonate shedding. Flat platform relief results in widely uniform δ13C values of sedimentary rocks, the ratio of which may be influenced by local shallow-water fractionation processes such as photosynthesis or evaporation; this contrasts with the former ramp-shaped relief with bathymetric gradient of carbon isotope ratios in sediments (Gischler et al., 2009). Despite the observed coincidence of Silurian CIEs with sea-level lowstands, not all recognized Silurian global sea-level falls are associated with CIEs (Kaljo et al., 1998). Hence, the eustatic explanation alone is insufficient and the presence of additional specific triggers for the CIE is needed. These triggers are often looked for in aridity changes (Samtleben et al., 2000), which cause the anti-estuarine type of oceanic circulation (Bickert et al., 1997). However, changes in aridity alone are also not the cause of the CIEs, which is indicated by the only periodical correlation between the two in the geological past. Additionally, changes in aridity are not synchronous on the global scale, because they may show opposite patterns in different regions during a single trend in the “global climate” (see e.g. McGee et al., 2010). According to Swart (2008), the carbon isotope record of recent neritic carbonate platforms is disconnected from the global carbon

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cycle. Moreover, similarly as in the case of the Silurian CIEs, the amplitude of Quaternary CIEs varies in different areas and in general decreases basinward (Gischler and Lomando, 2005; Swart, 2008; Gischler et al., 2009). Despite this fact, global-scale correlation between various areas using the δ 13C record is possible (Swart and Eberli, 2005; Swart, 2008; Gischler et al., 2009). These observations indicate globalscale processes, such as eustacy and world-wide climate attributes, acting as driving mechanisms of the globally synchronous (but probably locally induced) carbon isotope record. In the case of the Quaternary CIEs, climate aridity seems to influence carbon isotope fluctuations significantly only locally (Gischler and Lomando, 2005). The “global climate” has few attributes coherent at the global-scale. One of them is gustiness (McGee et al., 2010), which is strictly related to an increased temperature gradient between low and high latitudes during cooling events, which is consistent with the recently discussed climate cooling during the mid-Ludfordian (Kaljo et al., 2003; Lehnert et al., 2007a) and the presence of aeolian admixture in the CIE interval in the studied section (and probably also in others). Enhanced aeolian input during the mid-Ludfordian CIE was postulated previously by Stricanne et al. (2006), based on the increase in abundance of terrestrial spores in the interval on Gotland. It is important to note that, according to Stricanne et al. (2006), “the absolute abundance of terrestrial spores closely mirrors the marine δ 13C”, which is similar to our observations of links between aeolian input (DS-facies transgression, S-ratio, MS, natural gamma record) and carbon isotope development. According McGee et al. (2010, and references therein) enhanced gustiness causes global dustiness due to the long time of residence of dust in the air and unlimited transport ways. It is important to note that the gustiness needn't be strictly synchronous with glacioeustatic fluctuations (McGee et al., 2010). Another factor, which may additionally contribute to enhanced dustiness during gustiness periods, is synchronous low-latitude desertification (aridity) due to local arid climate conditions preferentially occurring during cooling events. Additionally, we may expect that in the early Paleozoic the dustiness induced by enhanced gustiness, was generally

Fig. 9. Oceanographic model for dust fertilized lowstand, arid platform conditions. Antiestuarine circulation model after Bickert et al. (1997).

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much greater, because of the minimal vegetation cover on land. This is consistent with the fact that Precambrian–Early Paleozoic CIEs have higher amplitudes than those in the Mesozoic and Cenozoic. We suggest that during the mid-Ludfordian CIE coupling of three interlinked factors occurred: prolonged sea-level lowstand conditions, enhanced global dust influx to shelf areas and local arid climate causing anti-estuarine circulation in low-latitudes shelves (Bickert et al., 1997). 6.6. Dust-fertilized platforms—possible links between global cooling and widespread, locally induced carbon-isotope record Antiestuarine circulation driven by evaporation, caused by aridity during the studied event (Bickert et al., 1997), could be significantly intensified by specific eustatic conditions. The event begins with a distinct sea-level drop, recorded as a sequence boundary (Eriksson and Calner, 2008), but the main part of the CIE is coeval with the prolonged incipient flooding occurring during persistent low sealevel conditions (Kozłowski and Munnecke, 2010). These specific circumstances may have formed extensive shallow areas with high evaporation and reduced circulation between the shelf interiors and open marine realm; which constituted an “engine” of the ariditydriven anti-estuarine circulation (Fig. 9). Circulation changes in concert with relatively low riverine influx due to climate aridity (Bickert et al., 1997) would cause an important reduction of nutrient supply to the shallow areas by rivers and sea-currents. However, despite the reduction in the abundance of acritarchs on Gotland, which would suggest oligotrophic conditions (Stricanne et al., 2006) the mid-Ludfordian CIE is coeval with the enhanced occurrence of microbial carbonates (Wigforss-Lange, 1999; Samtleben et al., 2000; Munnecke et al., 2003; Calner, 2005; Kozłowski and Munnecke, 2010), which indicates high primary productivity of (benthic) organisms like Girvanella and Rothpletzella (Eriksson and Calner, 2008), commonly interpreted as cyanobacteria. In the case of the earlier, but very similar, Mulde event in the late Wenlock Porębska et al. (2004) noted a similar acritarch scarcity, coeval with relative increase in the abundance of geohopanoids, which may also be indicative of a similar (but planktic) cyanobacteria bloom. In our opinion, these qualitative changes in the composition of primary producers may reflect perturbations in nutrient composition and availability (e.g. Subba Rao and Al-Yamani, 1999), particularly changes in the N:P (Radfield) ratio (see Arrigo, 2005 and references therein). As a result of probably limited nutrient supply by rivers and antiestuarine circulation, the potentially strong influence of aeolian nutrient delivery on nutrient composition and availability should be considered in our model. The main difference in the case of enhanced aeolian dust conditions is the significant increase of bioavailable iron (Fan et al., 2006, see also Soreghan and Soreghan, 2002; LaPorte et al., 2009); which in oligotrophic conditions stimulates cyanobacterial fixing of nitrogen (e.g. Mills et al., 2004). In our model (Fig. 9), in the case of antiestuarine circulation and low riverine nutrient input; the scarcity of phosphorus is an expected effect (see Wu et al., 2000). Simultaneously, local denitrification of fixed N2 should be limited (Fig. 9), because of prevailing aerobic conditions; due to the influx of surface oxygen-rich water masses caused by anti-estuarine circulation. Karl et al. (2008) showed that the P-stressed–N-available conditions (e.g. due to Fe-stimulated nitrogen fixation), are responsible for enhanced aerobic production of methane by modern ocean surface waters, as a by-product of methylphosphonate decomposition. In the present day, ocean surface water is in major part oversaturated in methane and much of this gas is generated in shelf areas (e.g. Bange et al., 1994). Despite this fact, the impact of this process on carbonate isotope signatures of DIC has not been emphasized. Aerobic methanogenesis, as a surface-related carbon isotope fractionation process, eventually should influence the isotopic composition in extensive shallow-water environments. The inflowing surface layer due to antiestuarine circulation may have additionally

amplified the effect, because of its enrichment in residual 13C due to primary production (see Bickert et al., 1997) and, additionally, depletion in 12C due the ocean-surface methane export. Because whole fractionation process occurs in the surface layer, correlation between depth and the degree of local average enrichment in residual 13C, in the entire water column, would be expected (Fig. 9). Because of the flat platform relief the heavy carbon isotope signatures in the shallows should be widespread and laterally constant. In turn, because of the circulation, the coastal water with heavy signatures of DIC could be exported offshore along the bottom, and influence (in various degree – depending on the mixing) the δ 13C of deeper water carbonate sediments (Fig. 9); however in the case of oxic conditions this process may be concurrent to 12C influx from decomposition of organic matter, which would have occurred in offshore environments as even negative δ 13C values (see e.g. Fig. 5 in Martma et al., 2005). With the end of the driving conditions (stillstand plus aridity plus aeolian fertilization), the decline of the isotope excursion should be slower in the inner shelf, because of: (1) an inherited high N:P ratio due to gradually opening circulation with the open ocean and slower denitrification; and (2) local buffering of carbon isotope signatures by partial reworking of carbonate sediments. Another, probably very important factor, contributing to the fast CIE decline in seaward areas would be restoration of estuarine circulation (humid climate); which should first influence the periplatform realm through the influx of deep anoxic water—allowing denitrification and inflow of dissolved phosphorus. Equally important are parallel transgressive conditions, which simply rapidly change bathymetry, and would cause gradual ramp relief regeneration, because of backstepping of the carbonate factory. In the last phase of the CIE, with recreated ramp topography, distinct correlation between the carbon isotope ratios in sediments and bathymetry may be expected. The subsequent gradual cease in restriction of circulation leads to gradual end of the CIE also in the inner shelf. Our model does not exclude the contribution (or even local domination) of additional, aeolian-aridity and circulation driven fractionation mechanism(s) in the formation of the mid-Ludfordian CIE. Moreover, the large amplitude of the event simply suggests an extraordinary interplay of several processes and/or signal amplification. The most important factors in respect to the assumed event conditions, which would act together with global “carbonate platform methanogenesis”, are: 1) additional methanogenesis resulting from bacterial fermentation of organic matter in bottom sediments of shallow and restricted water environments (Gu et al., 2004) 2) internal methanogenesis in microbial mats (e.g. formed by bluegreen alga—Behrens and Frishman, 1971). This mechanism may also be driven or supported by aeolian influx. 3) high intensity of the photosynthesis–remineralization cycle, caused by cyanobacterial colonization (Wigforss-Lange, 1999) and blooms (which may be also dust induced), caused bioinduced precipitation of carbonates characterized by heavier (residual) isotope signatures during productive periods (Merz, 1992; Thompson et al., 1997). This mechanism also may be supported or driven by aeolian influx, due to fertilization. 4) carbon isotope fractionation due to degassing and evaporation in coastal areas (Stiller et al., 1985; see also Swart, 2008) The influence of degassing and evaporation may be reflected by significant covariation of δ 18O and δ 13C in whole Mulde-Sundre interval in Gotland (Samtleben et al., 2000), and in the heaviest carbon isotope signatures being present in oolite facies (WigforssLange, 1999; compare Gischler and Lomando, 2005). On the other hand, extreme enrichment in oolites may also be caused by additional enrichment in 13C due to cyanobacterial alternation of the ooid cortex (Duguid et al., 2010), probably caused by the anoxic remineralization

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(methanogenesis) of cyanobacteria-related organic matter in the ooid grain microenvironment (see Ferguson and Ibe, 1981). Additional storage of the light organic carbon in the deep oceans also cannot be excluded; however, this process seems not to be necessary to explain the mid-Ludfordian CIE. 7. Conclusions 1) The periplatform Silurian succession of the Mielnik IG-1 core records the mid-Ludfordian CIE with maximal δ 13C values of +6.74‰ PDB; 2) The CIE in this section is coeval with significant natural gamma-ray and magnetic susceptibility negative anomalies. Both geophysical proxies appear to be very useful tools for detection of the event, particularly in the case of carbonate-poor successions; 3) Both anomalies are caused by conspicuous admixture of quartz and dolomite pelite, interpreted as wind-derived. Both minerals are diamagnetic and not radioactive and cause dilution of the common, facies-shift related, mineralogical record dominated by clay minerals; 4) The main positive shift in carbon isotope ratios is very rapid (estimated at 4000 years) and appears simultaneously with the appearance of a “dolomitic siltstone facies”; which records rapid increase of climate dustiness; 5) The pre-excursion gradual increase in δ 13C begins in the middle part of the Neocucullograptus kozlowskii Zone, while the main isotope shift occurs in the lowermost part of the N. kozlowskii– Pseudomonoclimacis latilobus Interzone; 6) During late phase of the maximal δ 13C values, coeval with sea-level lowstand, a P. dubius faunal bloom occurs; 7) The facies and geophysical records of the studied section confirm the coincidence of the beginning of the CIE with sea-level fall. The second shallowing event at the end of the CIE (Eriksson and Calner, 2008) is probably caused by normal regression due to the prolonged sea-level lowstand; 8) The CIE is coincident with sea-level lowstand, and the subsequent pronounced sea-level rise is coeval with a rapid decline in δ 13C values in the periplatform setting and a relatively slower decline in proximal shelf areas; 9) The heavier carbon isotope signatures of allochthonous sediments in the Mielnik IG-1 succession and the relatively low maximal CIE amplitude confirm the existence of a bathymetry-related sediment carbon isotope ratio gradient during the event as in the other Silurian CIEs; 10) The first appearances of the Pseudomonoclimacis latilobus– Monograptus balticus and Uncinatograptus acer graptolite faunas are coeval with MFS-s in post-CIE transgressive conditions. The events are separated by an ingression of the dubius fauna (Urbanek, 1997), appearing in stillstand conditions. 11) The MFS' sedimentary rocks record rapid carbon isotope declines. The pronounced part of the CIE in the offshore setting (Mielnik IG-1) ends at the first (latilubus–balticus) MFS. However, in the following interval, up to the second (Uncinatograptus acer) MFS, correlation between the carbon isotope ratios and reconstructed bathymetry (eK), suggest presence of a bathymetry-related carbon isotope gradient in rebuilding a carbonate ramp setting. This phenomenon is consistent with a more gradual CIE decline in proximal setting (e.g. Gotland—Samtleben et al., 2000) and indicate that in the inner shelf the CIE ends at the FA-Uncinatograptus acer MFS (c. 1 Ma before the end of Ludfordian—according to the thickness-to-time relation in the Mielnik IG-1 section). 12) The mid-Ludfordian CIEs shows similar features as the modern glacial-end isotope records of carbonate platforms (Swart and Eberli, 2005; Swart, 2008; Gischler et al., 2009). Both cases seem to be globally driven, but locally (environmentally) induced. We believe that the storage of light organic carbon in the deep oceans

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is not necessary for explanation of the discussed CIE (compare Swart, 2008). 13) The interplay of lowstand conditions, dustiness and aridity might have changed the availability of nutrients in an important way, particularly the N:P (Radfield) ratio in shallow shelf environments during the studied event. We postulate that aerobic methanogenesis in dust-fertilized platforms may significantly influence the carbon isotope ratios of carbonate sediments. Moreover, enhanced degassing and evaporation; internal methanogenesis in microbial mats and organic matter particles, and high carbonate precipitation bioinduction in periods of high productivity probably supported carbon isotope fractionation due to aerobic methane generation. 14) The coincidence of lowstand, low-latitude aridity, and gustiness of the global climate a supports hypothetical, short-lived (~250,000 years after estimated duration of the lowstand in Mielnik IG-1 section) glaciation in the mid-Ludfordian. Acknowledgments We thank our colleague Emilia Jarochowska for language support, fruitful discussions and very valuable comments. The paper benefited greatly from the reviews of Prof. Dimitri Kaljo, Tallin University of Technology and Prof. David Loydell, University of Portsmouth, who also greatly helped in language improvements. Research has been financed by IGP Grant n. 12/2010 and in part by the Polish Ministry of Education and Science Grant no. NN 307 013 237. The paper is a contribution to the IGCP 580 project. Appendix A. Supplementary data Supplementary data to this article can be found online at http:// dx.doi.org/10.1016/j.palaeo.2012.04.024. References Aaloe, A., Jürgenson, E., 1977. General rock types of the Baltic Silurian. In: Kaljo, D. (Ed.), Facies and Fauna of the Baltic Silurian. Academy of Sciences of the Estonian S.S.R. Institute of Geology, Tallinn, pp. 14–45 (In Russian with English summary). Adams, J.A.S., Weaver, C.E., 1958. Thorium to uranium ratios as indicators of sedimentary processes; example of concept of geochemical facies. AAPG Bulletin 42 (2), 387–430. Al-Bakri, D., Khalaf, F., Al-Ghadban, A., 1984. Mineralogy, genesis, and sources of surficial sediments in the Kuwait marine environment, northern Arabian Gulf. Journal of Sedimentary Research 54 (4), 1266–1279. Arrigo, R., 2005. Marine microorganisms and global nutrient cycles. Nature 437 (15), 349–355. Bange, H.W., Bartell, U.H., Rapsomanikis, S., Andreae, M.O., 1994. Methane in the Baltic and North Seas and a reassessment of the marine emissions of methane. Global Biogeochemical Cycles 8 (4), 465–480. Barrick, J.E., Kleffner, M.A., Gibson, M.A., Peavey, F.N., Karlsson, H.R., 2010. The midLudfordian Lau Event and carbon isotope excursion (Ludlow, Silurian) in southern Laurentia—preliminary results. Bollettino della Società Paleontologica Italiana 49 (1), 13–33. Bates, D.E.B., Kirk, N.H., 1984. Autecology of Silurian graptoloids. In: Bassett, M.G., Lawson, J.D. (Eds.), Autecology of Silurian organisms: Special Papers in Palaeontology, 32, pp. 121–139. Behrens, E.W., Frishman, S.A., 1971. Stable carbon isotopes in blue-green algal mats. Journal of Geology 79, 94–100. Betzler, C., Saxena, S., Swart, P.K., Isern, A., James, N.P., 2005. Cool-water carbonate sedimentology and eustasy; Pleistocene upper slope environments, Great Australian Bight (Site 1127, ODP LEG 182). Sedimentary Geology 175, 169–188. Bickert, T., Pätzold, J., Samtleben, C., Munnecke, A., 1997. Paleoenvironmental changes in the Silurian indicated by stable isotopes in brachiopod shells from Gotland, Sweden. Geochimica et Cosmochimica Acta 61, 2717–2730. Bloemendal, J., King, J.W., Hall, F.R., Doh, S.-J., 1992. Rock magnetism of Late Neogene and Pleistocene deep-sea sediments: relationship to sediment source, diagenetic processes and sediment lithology. Journal of Geophysical Research 97 (B4), 4361–4375. Bontognali, T.R.R., Vasconcelos, C., Warthmann, R.J., Bernasconi, S.M., Dupraz, C., Strohmenger, C.J., Mckenzie, J.A., 2010. Dolomite formation within microbial mats in the coastal sabkha of Abu Dhabi (United Arab Emirates). Sedimentology 57 (3), 824–844. Burgess, P.M., 2001. Modelling carbonate sequence development without relative sealevel oscillations. Geology 29, 1127–1130. Calner, M., 2005. A Late Silurian extinction event and anachronistic period. Geology 33 (4), 305–308.

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