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1 Continental weathering and redox conditions during the early Toarcian Oceanic
2 Anoxic Event in the northwestern Tethys: insight from the Posidonia Shale section
3 in the Swiss Jura Mountains
4
5 Jean-Carlos Montero-Serrano 1,2,3,*, Karl B. Föllmi 1, Thierry Adatte 1, Jorge E. Spangenberg 1,4, Nicolas
6 Tribovillard 5, Alicia Fantasia 1, Guillaume Suan 6
7
8 1 Institut des sciences de la Terre, Université de Lausanne, Bâtiment Géopolis, CH-1015 Lausanne, Switzerland
9 2 Institut des sciences de la mer de Rimouski, Université du Québec à Rimouski, 310 allée des Ursulines, Rimouski, Quebec G5L 3A1,
10 Canada
11 3 Geotop Research Centre, Université du Québec à Montréal, Montréal, Quebec H3C 3P8, Canada
12 4 Institut des dynamiques de la surface terrestre, Université de Lausanne, Bâtiment Géopolis, CH-1015 Lausanne, Switzerland
13 5 Laboratoire Géosystèmes, UMR-8217, Université Lille 1, CNRS, Villeneuve d’Ascq, France
14 6 UMR CNRS 5276 LGLTPE, Université Lyon 1, Campus de la Doua, Bâtiment Géode, F-69622 Villeurbanne Cedex, France 15
16 * Corresponding author: J.-C. Montero-Serrano. Phone: +1.418.723.1986 Ext: 1139; Fax: +1.418.724.1842; E-mail addresse:
17 [email protected] (J-C Montero-Serrano), [email protected] (K. Föllmi), [email protected] (T. Adatte),
18 [email protected] (J. Spangenberg), [email protected] (N. Tribovillard), [email protected] (A. Fantasia),
19 [email protected] (G. Suan) 20
21 Highlights
22 -The T-OAE negative CIE reflects a 13C-depleted carbon injection to ocean-atmosphere
23 -Detrital proxies suggest an increase in continental weathering rates during the T-OAE
24 -Redox proxies suggest moderate anoxic bottom-water conditions during the T-OAE
25 -Dysoxic-to-anoxic conditions prevailed until at least to the bifrons Zone
26 -P-enrichments are observed during the T-OAE
27
1
28 Abstract
29 The geochemistry and mineralogy of organic-rich sediments of the Rietheim succession in
30 northern Switzerland were studied to evaluate the main impacts of the early Toarcian Oceanic Anoxic
31 Event (T-OAE; ~183 Ma ago) on the depositional environment in the NW Tethys Ocean. The
13 13 32 geochemical data obtained (notably, G Corg, G CCaCO3, Rock-Eval pyrolysis, TOCcff) support the
33 hypothesis that the negative C-isotope excursion at the onset of the T-OAE mainly reflects a major
34 perturbation in the global carbon cycle rather than local environmental changes (e.g., basinal restriction,
13 15 35 productivity). Rock-Eval pyrolysis and G Corg-G Ntotal crossplots indicate that the organic matter was
36 primarily of marine origin during the T-OAE, and deposited in an epicontinental sea setting. Detrital
37 proxies [notably, detritus index, siliciclastic balance and Ln(Al2O3/Na2O)] suggest that the elevated
38 pCO2 levels and greenhouse climate prevailing during the T-OAE induced significant acceleration of the
39 hydrologic cycle and an increase in continental chemical weathering rates. Redox-sensitive proxies
40 [relative pyrite contents, organic and sulfide balances, V/(V+Ni) ratios, trace elements enrichment
41 factors, trace-element-TOC co-variation patterns and TOC/Ptotal molar ratios] indicate that the Rietheim
42 Posidonia Shale sediments accumulated under oxic-to-dysoxic bottom-water conditions, with more
43 reducing conditions during the T-OAE (anoxic and possibly euxinic). These changes in redox conditions
44 were probably induced by substantial thermohaline stratification driven by enhanced freshwater input
45 and the recurrent influx of brackish Arctic seawater into the NW Tethys through the Viking Corridor, in
46 agreement with previous sedimentological and geochemical evidences from the NW Tethys Ocean. In
47 the studied succession, such conditions were perhaps also responsible for sequestration of most of the
48 regenerated P in the deeper water column by adsorption and/or co-precipitation into authigenic phases.
49 Overall, the major palaeoenvironmental changes documented in this study provide a robust example of
50 Earth’s possible responses to the rapid onset of extreme greenhouse conditions.
2
51 Keywords: Early Toarcian Oceanic Anoxic Event; Posidonia Shale; Geochemistry; Mineralogy;
52 Paleoenvironment; Switzerland.
53
54 1. Introduction
55 The early Toarcian (~ 183 Ma ago) was a brief period in the Early Jurassic marked by an intense
56 global warming to about 5–10°C warmer than at present (e.g., Chandler et al., 1992; Rosales et al., 2004;
57 Dera et al., 2011). The lower Toarcian strata are characterized by a pronounced (4–7‰) near-global
58 negative carbon isotope (G13C) excursion recorded in organic matter, carbonates and fossil wood from
59 open- and marginal-marine sections (e.g., Jenkyns et al., 2002; Hesselbo et al., 2007; Al-Suwaidi et al.,
60 2010; Caruthers et al., 2011; Suan et al., 2011; Izumi et al., 2012; Reolid, 2014). It is currently thought
61 that this intense global warming period and its associated environmental changes were the result of a
62 major perturbation in the global carbon cycle, likely triggered by the combination of several geological
63 processes, such as: (1) greenhouse gas inputs resulting from intensified rift-related tectonics (Bernoulli
64 and Jenkyns, 2009) and large-scale eruptions in the Karoo-Ferrar large igneous province (Duncan et al.,
65 1997; Pálfy and Smith, 2000), (2) thermogenic methane release related with magmatic intrusions and
66 metamorphic alterations of Gondwanan coals and organic-rich shales in the Karoo-Ferrar province (e.g.,
67 Hesselbo et al., 2000; McElwain et al., 2005; Wignall et al., 2006; Svensen et al., 2007), (3) sudden
68 dissociation of large amounts (~2500 Gt; Beerling and Brentnall, 2007) of continental-shelf methane
69 hydrate deposits (e.g., Hesselbo et al., 2000; Kemp et al., 2005), possibly modulated by orbital forcing
70 (Kemp et al., 2005, 2011). Note that thermogenic and biogenic methane are both powerful greenhouse
13 71 gas enriched in isotopically-light carbon (G Cmethane < -60‰; e.g., Hesselbo et al., 2000), whereby a
72 massive liberation of these gases may rapidly disrupt the global carbon cycle.
73 The extreme greenhouse conditions prevailing during early Toarcian also favoured a period of
74 enhanced marine organic carbon deposition termed the early Toarcian Oceanic Anoxic Event or T-OAE
3
75 (e.g., Jenkyns, 1988; Baudin et al., 1990a). However, the expression of the early T-OAE has a wide
76 geographic variability in terms of organic matter content and degree of oxygenation of bottom waters.
77 Indeed, in many epicontinental sections from NW Europe and N Siberia (Fig. 1a), the T-OAE favoured a
78 widespread deposition of organic-rich lithologies, such as black shales (e.g., Jenkyns, 1988; Baudin et
79 al., 1990a,b; Hesselbo et al., 2000; Röhl et al., 2001; Cohen et al., 2004; Pearce et al., 2008; McArthur et
80 al., 2008; Hermoso et al., 2009a,b; Suan et al., 2011). However, in more open marine sections from S
81 Europe (Portugal - Suan et al., 2008; S Spain - Rodríguez-Tovar and Reolid, 2013), northern margin of
82 Gondwana (Oman - Immenhauser et al., 1998; Morocco - Bodin et al., 2010; Algeria - Reolid et al.,
83 2012, 2014a), as well as northern and southern Panthalassa Ocean (W Canada - Caruthers et al., 2011;
84 NE Japan - Gröcke et al., 2011; NW Argentina - Al-Suwaidi et al., 2010), these organic-rich lithologies
85 are absent or extremely condensed, suggesting that intensity of the T-OAE may have been strongly
86 modulated by local and regional factors. In fact, as suggested in McArthur et al. (2008), the degree of
87 water restriction and oxygen depletion varied regionally in Europe, increasing from South to North along
88 the European Epicontinental Seaway. Overall, many of these studies quoted above have led to debate
89 concerning the intensity of marine anoxia and geographic extent of the T-OAE.
90 In this context and in order to improve the characterization of the T-OAE within the NW Tethys
91 Ocean, we present here a new high-resolution multi-proxy record [including the analysis of stable
13 18 13 15 92 isotopes (G CCaCO3, G OCaCO3, G Corg, G Ntotal), bulk mineralogy, elemental geochemistry, and Rock-
93 Eval pyrolysis] of an early Toarcian section located in the northern Swiss Jura Mountains (Rietheim
94 village, Canton Aargau; Fig. 1). In this study, these new geochemical and mineralogical data have been
95 used not only to explore the expression of the early T-OAE in an area not explored yet for its
96 geochemistry, but also to better document the main impacts of this event on the depositional
97 environment in the NW Tethys Ocean.
98
4
99 2. Geological setting
100 2.1. Depositional setting
101 During the Early to Middle Jurassic, the breakup of Pangaea led to the opening of the Central
102 Atlantic and Tethys oceans (Fig. 1a), accompanied by widespread tectonic extension in what is now the
103 Alpine-Mediterranean region, as well as in northern Europe (e.g., Bernoulli and Jenkyns, 2009). In
104 western Europe, the early Toarcian is characterized by widespread organic-rich lithologies deposited in
105 semi-restricted basins surrounded by shallow shelf seas extending across the West European platform to
106 the northern Tethys Ocean (Fig. 1a; Jenkyns et al., 2002; Wignall et al., 2005; McArthur et al., 2008;
107 Suan et al., 2008; Hermoso et al., 2009a,b, 2013). McArthur et al. (2008) suggested that water restriction
108 varied regionally, being greatest in areas deeper and more distant from the Tethys Ocean (e.g., the
109 Cleveland Basin; United Kingdom) and less restricted towards the south and east across the West
110 European platform (e.g., the Paris Basin and SW Germany).
111 The Lower Jurassic stratigraphy in northern Switzerland (tabular and folded Jura; Fig. 1b) has
112 been summarized recently by Reisdorf et al. (2011). The Early Jurassic Swiss deposits accumulated in
113 the relatively slowly subsiding transition zone between the SW Germany Basin and the eastern Paris
114 Basin under fully marine conditions (Reisdorf et al., 2011). Therefore, the lithology and fossil content of
115 the Lower Jurassic deposits in northern Switzerland (such as the Posidonia Shale, renamed as Rietheim
116 Member) are similar to those in SW Germany (e.g., Kuhn and Etter, 1994; Röhl et al., 2001; Röhl and
117 Schmid-Röhl, 2005; Reisdorf et al., 2011).
118
119 2.2. Early Toarcian Posidonia Shale in northern Switzerland
120 The studied lower Toarcian Posidonia Shale is located in the Southern Tabular Jura,
121 approximately 400 m SW of the village of Rietheim (Fig. 1b; Canton Aargau; 47°35’48.75’’N,
122 8°16’39.12’’E; cropping out in the bed of the Steigstrasse Stream). This Posidonia Shale succession
5
123 (~790 cm thick) has been lithostratigraphically and biostratigraphically well studied (Kuhn and Etter,
124 1994) including, in fact, three main ammonite biozones (tenuicostatum, falciferum, and bifrons) and
125 several subzones (exaratum, elegans, falciferum, and commune) (Fig. 2). The transition between the
126 tenuicostatum and falciferum zones is not precisely defined due to a scarcity of index fossils and/or to
127 the presence of a stratigraphic gap in the Pliensbachian-Toarcian transition interval (Kuhn and Etter,
128 1994).
129 The bottom of the section, a glauconite-rich marl layer (~4 cm thick) and light-grey bioturbated
130 marls, attributed to the spinatum Zone from the upper Pliensbachian, are intercalated with two compact
131 marly-limestone beds about 10 and 18 cm thick, respectively (Fig. 2). The light-grey bioturbated marls
132 contain diverse benthic micro- and macrofauna (including ammonites, belemnites and shell coatings) as
133 well as phosphatic fish debris, disseminated framboidal pyrite, and pyritized fossils. The tenuicostatum
134 Zone is represented by the transition of light-grey to dark-grey marly clays with generally low organic
135 contents (<1%) and relatively abundant benthic fauna (notably bivalves and brachiopods). These marly
136 clays (~ 30 cm thick) become more laminated and dark towards the upper part of the tenuicostatum
137 Zone, where continuous organic-rich mudstone sedimentation started (up to the bifrons Zone). Indeed,
138 black to brownish laminated bituminous marly clays (“Kartonschiefer-Fazies”; with more than 6%
139 organic carbon) including fish and wood debris, pyritized ammonites, and disseminated framboidal
140 pyrite prevail at the base of falciferum Zone. In contrast, weakly laminated black to grey bituminous
141 marly clays (“Tonmergel-Fazies”; 1 to 6 wt.% organic carbon) dominate from the mid falciferum to the
142 bifrons zones (Fig. 2). Silt-sized quartz grains are present in all types of laminations. Phosphate is
143 commonly in the form of skeletal apatite at the base of falciferum Zone (as seen in thin sections; Kuhn
144 and Etter 1994). Likewise, several thin bioturbated layers with moderately diverse benthic fauna occur
145 within this bituminous marly-clayey facies. In fact, in comparison to the sediments at the base of
146 falciferum Zone, the abundance of benthic fauna (e.g., Pseudomytiloides dubius) and trace fossils (e.g.,
6
147 Chondrites sp. and Planolites sp.) increases in the upper part of the section (falciferum to commune
148 subzone).
149 The organic-rich mudstones are separated by several finely laminated limestone beds
150 (Spinatumbank, Unterer Stein, Homogene Kalkbänke, Oberer Stein and Monotisbank) into a lower,
151 middle, and upper part (Fig. 2). These limestone beds are likely of early diagenetic origin (Röhl et al.,
152 2001; Röhl and Schmid-Röhl, 2005). The well-laminated limestone beds Unterer Stein and Oberer Stein
153 can be traced over many kilometres within the epicontinental basin and are important marker beds (e.g.,
154 Kuhn and Etter 1994; Röhl et al., 2001). The micro-laminations are composed of alternating
155 argillaceous, organic-rich, and fecal pellet/coccolith-rich layers (Kuhn and Etter 1994; Röhl et al., 2001).
156 Finally, the top of the section (uppermost commune subzone) is marked by a 10–12 cm thick limestone
157 bed (Monotisbank) with abundant shell debris (mainly the bivalve Pseudomonotis substriata) and
158 ammonites (especially Dactylioceratidae).
159
160 3. Material and methods
161 3.1. Samples
162 The lower Toarcian Posidonia Shale succession of Rietheim was sampled systematically along
163 the approximately 790 cm of the stratigraphic log (Fig. 2). A total of 226 samples were collected: 155
164 samples (vertical sample spacing, 2 cm) from a drill core in the first 316 cm of the stratigraphic log (Fig.
165 2), and 71 outcrop samples (vertical sample spacing, 5–10 cm) in the last 470 cm (Fig. 2). An aliquot
166 (about 30 g) of each sample was crushed and pulverized using a Shatterbox with agate mills. Then, each
167 sample was homogenized by sieving through a 100 mesh screen.
168 3.2. Analytical procedure
169 3.2.1. Stable isotope analyses (carbon, oxygen, and nitrogen)
7
13 18 170 Bulk carbonate carbon (G CCaCO3) and oxygen (G OCaCO3) isotope analyses were carried out on
171 powdered bulk-rock samples using a Thermo Fisher Scientific GasBench II and carbonate preparation
172 device interfaced to a Thermo Fisher Scientific Delta Plus XL continuous-flow isotope ratio mass
13 15 173 spectrometer. Organic carbon (G Corg) and total nitrogen (G Ntotal) isotopes were also measured on the
174 organic fraction of decarbonated (10% HCl treatment) samples using a Carlo Erba 1108 elemental
175 analyser connected to a Thermo Fisher Scientific Delta Plus V isotope ratio mass spectrometer via a
176 Confo III interface. The stable carbon and oxygen isotope ratios are reported in the delta notation as the
177 per mil (‰) deviation relative to the Vienna Pee Dee Belemnite (VPDB) standard. The total nitrogen
178 isotope ratios are also reported in the delta notation as the per mil (‰) deviation relative to atmospheric
13 12 18 16 15 14 179 nitrogen (air N2) standard. δ = [(Rsample/Rstandard) - 1] x 1000 with R= C/ C, O/ and N/ N.
180 Analytical precision and accuracy were determined by replicate analyses and by comparison with
181 international and in-house standards (Carrara marble, USGS-24, USGS-32, IAEA-N3, UREA, glycine,
182 and pyridine), and were better than ±0.1‰ (1σ) for carbon, oxygen, and nitrogen.
183 3.2.2. Rock-Eval pyrolysis
184 The type and thermal maturity of the bulk organic matter (OM) was established by Rock-Eval
185 pyrolysis using a model 6 device (Vinci Technologies) under standard conditions (see Béhar et al., 2001
186 for procedural details). The total organic carbon content (TOC, wt.%), Tmax index (°C; OM thermal
187 maturity indicators), Hydrogen Index (HI, mg HC/g TOC), and Oxygen Index (OI, mg CO2/g TOC)
188 were also determined. The HI and OI signatures allow a rough distinction between algal/bacterial (types
189 I and II kerogen) and terrestrial plant (types III and IV kerogen) OM, as well as the degree of
190 degradation and oxidation. Analytical precision was better than 0.05 wt.% (1σ) for TOC, 1.5°C (1σ) for
191 Tmax, 10 mg HC/g TOC (1σ) for HI, and 10 mg CO2/g TOC (1σ) for OI.
192 3.2.3. Total organic carbon on the carbonate-free fraction
8
193 In order to complement Rock-Eval data on the OM characteristics, we also quantified TOC
194 concentrations on the carbonate-free fraction (cff) using a CHN Elemental Analyser (Carlo Erba Flash
195 EA 1112 CHNS/MAS200). The decarbonated fraction was obtained by a double 10% HCl treatment.
196 Precision was better than 1% based on an internal standard (methionine) and replicate samples. The
197 TOCcff allows a robust signature of the organic matter flux regardless of dilution due to carbonate input.
198 3.2.4. Bulk-rock mineralogy
199 The bulk-rock mineralogy was analysed by X-ray diffraction using a Thermo Scientific ARL
200 X’TRA powder diffractometer. This instrument is fitted with a copper tube (Cu K-alpha = 1.540562Å),
201 operating at 40 kV and 30 mA and a post-diffraction graphite monochromator. The random powdered
202 rock samples were scanned from 2° to 65° two-theta in steps of 0.02° two-theta and a counting time of 1
203 second per step, resulting in 3151 data points. For the semi-quantification of the major mineralogical
204 components, the bulk-rock XRD scans obtained were converted into mineral weight percents (wt.%)
205 using the RockJock computer program (Eberl, 2003). This RockJock technique uses a full-pattern fitting
206 method that permits the semi-quantification of whole-sediment mineralogy with a precision of 1-2%. To
207 check the quality of this fitting procedure, a degree-of-fit (DOF = minimum absolute difference) statistic
208 is calculated between the measured and simulated XRD patterns. The DOF obtained with our samples
209 was good with values ranging from as low as 0.131 to as high as 0.402 and averaging 0.202 for the 187
210 samples. For this study the wt.% of the following minerals were calculated (Fig. 6): quartz, potassium
211 (K) feldspar (microcline + orthoclase), plagioclase feldspar (albite), phyllosilicates (biotite, muscovite,
212 illite, chlorite, and kaolinite), pyrite, dolomite, ankerite, and carbonate fluorapatite (CFA).
213 Because the quartz grains are principally detrital in the whole section (as seen in thin sections;
214 Kuhn and Etter 1994), we determined a detritus index, calculated by summing quartz, phyllosilicates, K-
215 feldspar, and Na-plagioclase contents, to observe changes in detrital influx (e.g., Keller et al., 2004;
216 Westermann et al., 2010; Stein et al., 2012; Bomou et al., 2013). High detritus index values may indicate
9
217 higher delivery of terrigenous material from continental sources and/or decreased carbonate productivity
218 and dissolution.
219 3.2.5. Major, minor and trace element analysis
220 Loss on ignition (LOI) was determined gravimetrically by heating the dried samples up to
221 1050°C for two hours. Major and trace-element concentrations were determined by X-ray fluorescence
222 (XRF) spectrometry using a PANalytical PW2400 with an RX tube (Rh anode). Major (and minor)
223 elements (Si, Ti, Al, Fe, Mn, Mg, Ca, Na, K, P) were measured after sample fusion with tetraborate, and
224 trace elements (Nb, Y, Rb, Zn, Cu, Ni, Co, Cr, V, Nd, La, As) after mixing the samples with Mowiol
225 polyvinyl alcohol (2%) and pressed to pellets. The detection limit for major elements is better than 0.01
226 wt.% and for trace elements is generally 1 to 4 ppm. Analytical reproducibility monitored by replicate
227 analyses of selected samples was lower than ±5% for major, minor, and trace elements. Analysis
228 accuracy was assessed by analyses of international and in-house standard reference materials (TS1-
229 Cement, TS3-Clay, TS4-Limestone, TS5-Marlstone, TS7-Sandstone, 372-Portland cement, 368-
230 Dolomite).
231 A compositional Q-mode cluster analysis (e.g., Egozcue and Pawlowsky-Glahn, 2005;
232 Pawlowsky-Glahn and Egozcue, 2011) was performed on our elemental geochemical dataset with the
233 goal of finding elemental associations with similar relative variation patterns that may be interpreted
234 from a palaeoenvironmental standpoint (e.g., Jaminski et al., 1998; Hoffman et al., 1998; Reátegui et al.,
235 2005; Montero-Serrano et al., 2010). This analysis was carried out using a log-ratio approach (Aitchison,
236 1986; Egozcue and Pawlowsky-Glahn 2005). Likewise, as measure of dissimilarity we use the variation
237 array, and as clustering criterion the Ward method. Importantly, in order to isolate the siliciclastic
238 component of the marls and marly clays, calcium-rich samples (Ca > 25 %, ~71 samples) are excluded
239 from the Q-mode cluster analysis. All statistical calculations were conducted with “R” software using the
10
240 packages “StatDA” (Reimann et al., 2008) and “compositions” (van den Boogaart and Tolosana-
241 Delgado, 2008).
242 In order to infer palaeoenvironmental changes (weathering, sediment supply, redox conditions)
243 during the early T-OAE, the stratigraphic distribution of the elemental geochemical associations
244 obtained by the compositional Q-mode cluster analysis were represented as elemental balances.
245 Elemental balances are log-contrasts resulting from a log-ratio of two geometric means of two non-
246 overlapping elemental groups (Egozcue and Pawlowsky-Glahn, 2005). The general expression of an
247 elemental balance is:
248 b = [√ (r x s / r + s)] * Log[gm(A)/gm(B)] [1]
249 where A, B are two non-overlapping elemental groups of a complete composition x of D elements (e.g.,
250 Nb, Rb, La, K, Ti, Al, Si,…, xD), r and s (r + s ≤ D), are the number of elements in A, B respectively,
251 and gm(.) denotes the geometric mean of the elemental groups A and B.
252 Alternatively, the chemical index of alteration (CIA), defined as: CIA = 100 × Al2O3/(Al2O3 + CaO +
253 Na2O + K2O) in molar proportions, may also be used as a proxy to the degree of weathering in the source
254 areas with higher values suggesting more intense chemical weathering (Nesbitt and Young, 1982).
255 However, samples that considerably vary in CaO due to presence of carbonate grains or cements (calcite
256 or dolomite) and phosphates, such as samples studied here, may suggest misleading conclusions if the
257 CIA is used to infer the degree of weathering (e.g., Cullers, 2000; Goldberg and Humayun, 2010). Under
258 this context and based on a statistically model of linear compositional and weathering trends, von
259 Eynatten et al. (2003) suggest that Ln(Al2O3/Na2O) (in molar proportions) may be used as a weathering
260 index similar to CIA. Therefore, we prefer to use the Ln(Al2O3/Na2O) rather than CIA to infer changes
261 in the degree of chemical weathering during deposition of the Rietheim Posidonia Shale succession.
262 Using Ln(Al2O3/Na2O) has the advantage of avoiding uncertainties concerning the correction for CaO
263 associated with phosphate and carbonate phases (e.g., Cullers, 2000; von Eynatten et al., 2003).
11
264 In addition, to compare the relative enrichment of redox-sensitive trace elements (RSTE: Cu, Ni,
265 Zn, V, Co, As), we calculated Enrichment Factors (EF) by comparing Al-normalized metal concentration
266 to those of average shale (Wedepohl, 1991): XEF = [(X/Al)sample / (X/Al)average shale]. In practical terms, EF
267 > 3 represents a detectable authigenic enrichment of an element over average shale concentrations,
268 whereas EF > 10 represents a moderate to strong degree of authigenic enrichment (e.g., Tribovillard et
269 al., 2006; Algeo and Tribovillard, 2009).
270 All analytical data presented are available electronically in the Supplementary Appendix A.
271
272 4. Results
273 4.1. Stable isotopes
13 274 Carbon isotope values of bulk OM (G Corg ~ -32.6 to -25.3‰; Fig. 3a) and the carbonate
13 275 (G CCaCO3 ~ 1.2 to -2.6‰; Fig. 3b) of marls and marly clays show comparable variations throughout the
276 section, with a pronounced negative carbon isotope excursion (or CIE) within the exaratum subzone
13 13 277 (Fig. 3a-b). The magnitude of this negative CIE varies from ~3‰ in G CCaCO3 to ~5‰ in G Corg records.
13 13 278 On average, the G Corg values are in the range of G Corg records of marine OM deposited during
13 279 pronounced greenhouse phases (e.g., Dean et al., 1986; Jenkyns, 2010), whereas G CCaCO3 values are
13 280 within the range of early Toarcian G CCaCO3 records from epicontinental seas of the Peritethyan Realm
281 (e.g., Röhl and Schmid-Röhl, 2005; van de Schootbrugge et al., 2005; Hesselbo et al., 2007; Suan et la.,
282 2008; Hermoso et al., 2009b).
13 283 Extremely light G CCaCO3 values are found within limestone beds and bioturbated marls (e.g., -
13 13 284 4.57‰ in the Unterer Stein; Fig. 3b).G Corg and G CCaCO3 values decrease slightly up section. Oxygen-
18 285 isotope values of the bulk carbonate (G OCaCO3) shows higher values in the upper Pliensbachian interval
286 (-6 to -2.7‰) compared to the lower Toarcian (-8 to -7‰) (Fig. 3c). However, in the Spinatumbank,
18 287 Unterer Stein, and Homogene Kalkbänke limestone beds, the G OCaCO3 values are significantly higher (- 12
15 288 4 to -2.7‰). Nitrogen-isotope values of OM (G Ntotal) range between +1 and +3‰ throughout the
15 289 section (Fig. 3d), with lower values (~ +0.4‰) in the limestone bed Unterer Stein. These G Ntotal values
290 are within the range of values that characterized the Jurassic-Cretaceous interval (Algeo et al., 2014).
291 4.2. Organic matter characterization
292 Total organic carbon (TOC), total organic carbon from the carbonate-free fraction (TOCcff), and
293 the hydrogen index (HI) show lower values (TOC < 0.5 wt.%; TOCcff < 2 wt.%; HI < 70 mg HC/g TOC)
294 in the upper Pliensbachian interval compared to the lower Toarcian (Fig. 3e-g), whereas oxygen index
295 (OI) show higher values (up to 422 mg CO2/g TOC; Fig. 4a). In fact, in the lower Toarcian interval,
296 TOC, TOCcff, HI and OI values range from 2 to 11.8 wt.%, 5 to 18 wt.%, 400 to 794 mg HC/g TOC, 49
297 to 422 mg CO2/g TOC, respectively. High TOC (6 to 11.8 wt.%) and HI (600 to 794 mg HC/g TOC) as
298 well as low OI (<10 mg CO2/g TOC) values characterize the black to brownish laminated bituminous
299 marly clays (exaratum subzone) recording the early Toarcian CIE, whereas lower TOC (< 2 wt.%) and
300 HI (HI < 360 mg HC/g TOC) values characterize the early-diagenetic limestone beds as well as the
301 bioturbated marl intervals. Rock-Eval Tmax values average 427°C without any significant variation (Fig.
302 3h).
303 4.3. Bulk mineralogical composition
304 The bulk mineralogy in the Rietheim Posidonia Shale is dominated by calcite, phyllosilicates, and
305 quartz, and lower proportions of K-feldspar, Na-plagioclase, pyrite, ankerite, dolomite, and carbonate
306 fluorapatite or CFA (Fig. 6 and 7). An unidentified group of organic material and some poorly
307 crystallized minerals (e.g., iron oxide) are unquantified and not considered here (Fig. 7j). In general,
308 calcite distribution is inversely correlated to phyllosilicates, quartz, K-feldspar and Na-plagioclase (Fig.
309 7a-e). Calcite is the dominant mineral (60–96%) from the finely laminated limestone beds (e.g., Unterer
310 Stein, Homogene Kalkbänke, Oberer Stein, and Monotisbank), whereas phyllosilicates (19–53%) and
311 quartz (5–18%) are the dominant minerals in the marl and marly clay intervals (Fig. 7a-e). A decrease in
13
312 calcite content from 60% to 25% is observed at the base of the exaratum subzone (Fig. 7e). The upper
313 Pliensbachian (spinatum Zone) interval is characterized by high ankerite (up to 18%) and dolomite (up to
314 7.5%) contents (Fig. 7f-g). Indeed, the vertical distribution of dolomite mimics the ankerite record. The
315 CFA ranges between 0% to 1.6%, with higher values at the base of falciferum and bifrons zones (Fig.
316 7h). The pyrite content increases (up to 6.9%) in exaratum subzone (broadly coeval with the negative
317 CIE) and then slowly decrease upsection (Fig. 7i).
318 4.4. Major, minor and trace element analysis
319 The compositional Q-mode cluster analysis of the Rietheim Posidonia Shale succession samples
320 revealed three major elemental geochemical associations (Fig. 8):
321 (1) Siliciclastics (Nb, Rb, La, K, Ti, Al, Si, Cr, Na, Y, Nd, Mg): this geochemical group consists of
322 elements associated with clay and silt particles, mainly as chemical constituents of the detrital
323 mineral grains (e.g., phyllosilicates, quartz, Na-plagioclases, K-feldspar) or adsorbed on their
324 surface (e.g., Calvert and Pedersen, 2007; Montero-Serrano et al., 2010). However, Mg shows a
325 different vertical distribution with respect to the other siliciclastics elements (Fig. 8a), likely
326 suggesting that this element is also associated with dolomite and ankerite (Fig. 7f-g). Overall, the
327 variability of riverine and aeolian inputs related to continental climatic fluctuations as well as the
328 hydraulic partitioning of the sediments mostly control the distribution and accumulation of these
329 elements.
330 (2) Organic (P, As, TOC, Ni, V) – sulfide (Zn, Cu, Fe, Co): these elements are commonly associated
331 with the OM and the authigenic sulfides (e.g., pyrite) fraction, whose accumulation is largely
332 controlled by the oxic-anoxic conditions of the sedimentary environment (e.g., Jaminski et al.,
333 1998; Algeo and Maynard, 2004; Tribovillard et al., 2006; Calvert and Pedersen, 2007). Among
334 these elements, V/(V+Ni) ratio and TOC/Ptotal molar ratios are reputed as paleoredox-indices
335 (e.g., Hatch and Leventhal, 1992; Jones and Manning, 1994; Algeo and Maynard, 2004;
14
336 Tribovillard et al., 2006; Algeo and Ingall, 2007). Indeed, V/(V+Ni) ratios lower than <0.46,
337 0.46–0.60, 0.54–0.82, greater than 0.82 may suggest oxic, dysoxic, anoxic and euxinic conditions
338 in the water column., respectively (Hatch and Leventhal, 1992).
339 (3) Carbonates (Ca, Mn): this geochemical group consists of elements related with calcium carbonate
340 (mainly calcite) associated with calcareous cement as well as benthic fauna (including
341 ammonites, belemnites, shell coatings) and coccoliths, which agrees with the petrographic
342 interpretations (Kuhn and Etter 1994). The carbonate-associated elements correlate negatively
343 with the other two elemental associations, and thus likely represent a dilution fraction.
344 Based on these geochemical associations, elemental balances are performed (e.g.,
345 b{siliciclastic/carbonate}, b{sulphide/carbonate} and b{organic/carbonate}). The vertical distribution of siliciclastic (Nb-Rb-
346 La-K-Ti-Al-Si-Cr-Na-Y-Nd), sulfide (Fe-Cu-Zn-Co) and organic (TOC-P-As-V-Ni) balances show
347 similar patterns across the Rietheim Posidonia Shale succession (Fig. 8a-d; Fig. 10c; Fig. 11b-c), with
348 minimum values recorded in the spinatum Zone and maximum values in the negative CIE (exaratum
349 subzone) that then slowly decrease upsection. Similarly, Ln(Al2O3/Na2O) values also increase
350 significantly during the negative CIE (exaratum subzone), then slowly decrease upsection and increasing
351 again during the upper part of commune subzone (Fig. 10b). The V/(V+Ni) ratio range between 0.50 to
352 0.85 with maximum values recorded also during the negative CIE that then slowly decrease upsection
353 (Fig. 11d). RSTE enrichment factors (Cu, Ni, Zn, V, Co, As) show a modest enrichment (2 to 9) within
354 the negative CIE compared to average shale values (Fig. 11e-f). Finally, a good correspondence is
355 observed between the different detrital [detritus index, siliciclastic balance, Ln(Al2O3/Na2O); Fig. 10]
356 and redox [pyrite content, sulfide and organic balances, V/(V+Ni); Fig. 11a-f, i] proxies used here.
357 4.5. Phosphorus contents and TOC/Ptotal molar ratios
358 In the Rietheim succession, profiles of total P contents mimic the stratigraphic trend of the CFA,
359 with higher values during the negative CIE and early bifrons Zone (Figs. 8d and 11g). Furthermore,
15
360 higher P values correspond also to higher TOC values during the negative CIE (Fig. 8d), whereas lower
361 total P contents are characterized in the upper part of exaratum and at the base of elegans subzones.
362 We also determined Corg/Porg molar ratios to evaluate the burial efficiency of P relative to total
363 organic-matter preservation in the sediments (e.g., Algeo and Ingall, 2007; Westermann et al., 2013).
364 Indeed, Corg/Porg molar ratios measured in organic-rich sediments may reflect the importance of
365 preferential P release under dysoxic-to-anoxic bottom-water conditions and as such serve as a
366 paleoredox proxy (e.g., Bodin et al., 2006; Algeo and Ingall, 2007; Mort et al., 2007; Bodin et al., 2010;
367 Stein et al., 2011; Westermann et al., 2013). In this study, we use TOC/Ptotal molar ratios (Ptotal = Preactive
368 + Pdetritic; assuming the Pdetritic is likely to represent well-crystallized authigenic carbonate fluorapatite)
369 rather than Corg/Porg or Corg/Preactive (Preactive = Poxide-associated + Pauthigenic+ Porg; Anderson and Delaney,
370 2001), since as discussed in Algeo and Ingall (2007), the Pdetritic is generally a minor component (mean
371 ~13%) of Ptotal in black shales units. This assumption is in agreement with the compositional Q-mode
372 cluster analysis, where Ptotal is associated with TOC (Fig. 8d-e). The TOC/Ptotal molar ratios calculated
373 for the Rietheim Posidonia Shale succession range from 37 to 383 (median ~ 167) at the base of
374 falciferum Zone, with maximum values within the upper part of the exaratum subzone (Fig. 11h).
375 Conversely, in the spinatum, upper part of the falciferum and base of commune subzones, TOC/Ptotal
376 molar ratios are lower, ranging from 33 to 160 (median ~ 78). In general, the increases in the TOC/Ptotal
377 molar ratios are associated to relatively low P concentrations (Fig. 8d) and CFA contents (Fig. 11g).
378
379 5. Interpretation and discussion
380 5.1. Diagenetic effects
13 18 381 The anomalous G CCaCO3 and G OCaCO3 values observed within limestone beds and bioturbated
382 marls in the Rietheim Posidonia Shale succession likely reflect early diagenetic carbonate precipitation
383 (e.g., Algeo et al., 1992; Röhl and Schmid-Röhl, 2005). Therefore, in order to improve the
16
384 palaeoenvironmental signature of our geochemical records, these limestone beds and bioturbated marls
385 (~ 71 samples of the global geochemical dataset) were excluded from the subsequent interpretation and
386 discussion.
387 Stable oxygen isotopes in carbonates are more affected by diagenesis compared to carbon
18 388 isotopes (Schrag et al., 1995). At the Rietheim Posidonia Shale succession, the observed G OCaCO3
389 pattern (higher values in limestone beds and lower in marls; Fig. 3c) most likely reflects the timing of
390 burial diagenesis, rather than changes in the palaeoceanic isotope signal (such as salinity and
391 temperature). Indeed, during burial diagenesis, limestone cement forms early, in shallow depths and
392 lower burial temperatures, and thus acquires higher G18O values (-5 to -2‰) compared to marls (-9 to -
393 6‰), which cement later (e.g., Algeo et al., 1992; Schrag et al., 1995; Rosales et al., 2001).
18 13 394 In this context and despite the diagenetic alteration of the G OCaCO3 signature, the G CCaCO3 as
13 395 well as the G Corg records probably represent a mainly palaeoenvironmental origin because they are in
396 the same range of values and have similar rates of changes to those found in the Tethys, subboreal and
397 Proto-Atlantic Oceans (e.g., Jenkyns, 1988; Hesselbo and Jenkyns, 1998; Hesselbo et al., 2000; Röhl et
398 al., 2001; Röhl and Schmid-Röhl, 2005; Suan et al., 2008; Hermoso et al., 2009a,b; and references
399 therein).
400 5.2. Maturity, sources and preservation of the organic matter
401 Rock-Eval Tmax values average 427°C without any significant variation, indicating that the OM in
402 the Rietheim Posidonia Shale succession is immature with regards to hydrocarbon generation (Fig. 3h).
403 Therefore, variations in HI and OI cannot be attributed to changes in thermal maturity. The sudden shifts
404 towards lower HI values (<450 mg HC/g TOC) observed through the lower Toarcian interval (e.g., in the
405 thin bioturbated horizons between the falciferum and commune subzones; Fig. 3g and Fig. 4a) likely
406 suggest enhanced OM degradation during intermittent periods of oxygenation (e.g., Röhl and Schmid-
407 Röhl, 2005). The vertical distribution of HI (Fig. 3g) together with HI-OI crossplots (Fig. 4a) indicate
17
408 that OM in the upper Pliensbachian interval mainly consists of recycled and/or terrestrial material (type
409 III/IV kerogen), whereas in the lower Toarcian interval, it is mainly dominated by marine phytoplankton
410 and algal material (type I/II kerogen). Similar OM characteristics have also been reported in nearby
411 lower Toarcian successions from SW Germany (e.g., Röhl et al., 2001; Röhl and Schmid-Röhl, 2005).
13 15 412 The G Corg-G Ntotal crossplots (Algeo et al., 2014) suggest that the Rietheim Posidonia Shale
413 succession was deposited in an open epicontinental sea setting (Fig. 4c). In this depositional setting the
414 variability in sedimentary dynamics might have been driven by fluctuations in sea level (Reisdorf et al.,
415 2011). Indeed, sequence stratigraphic studies from NW Tethys have provided evidence for sea-level fall
416 close to the onset of the early Toarcian negative CIE (e.g., Perilli, 2000; Wignall et al., 2005; Pittet et al.,
417 2014). Accordingly, the OM source changes observed in the Rietheim Posidonia Shale succession may
418 have been related to eustatic sea-level changes. A sea-level lowstand during the upper Pliensbachian
419 interval may explain the increase in terrestrial-derived OM, whereas the enhanced input of marine-
420 derived OM in the lower Toarcian interval may correspond to a rising sea-level trend. Likewise, the
421 relatively high TOC (up to 11.8 wt.%) and TOCcff (up to 18 wt.%) recorded in the lower Toarcian
422 interval (Fig. 3e-f) suggests that likely elevated surface productivity and/or oxygen-depleted bottom-
423 water conditions prevailed in the NW Tethys during the T-OAE. These depositional conditions were
424 very favourable to OM preservation in the Rietheim Posidonia Shale succession. Similar results have
425 been reported by Röhl et al. (2001) and Röhl and Schmid-Röhl (2005) in the Posidonia Shale succession
426 from SW Germany.
427 5.3. Link between organic matter source and organic carbon isotopes
13 13 428 The larger amplitude of the negative CIE observed in G Corg records relative to G CCaCO3 in the
429 Rietheim Posidonia Shale might be related by changes in OM sourcing (e.g., van de Schootbrugge et al.,
13 430 2013). The influence of changes in the type of preserved OM on the G Corg records may be explored by
13 431 comparing pre- and post-excursion G Corg values with coeval HI values (e.g., van de Schootbrugge et
18
13 432 al., 2013). In the Rietheim Posidonia Shale succession,G Corg-HI crossplots data from the spinatum and
13 433 bifrons zones show a significant negative correlation (Fig. 4b), suggesting that part of the bulk G Corg
434 records likely reflect changes in OM sourcing. Following the approach proposed by van de
435 Schootbrugge et al. (2013), the residual changes that are not explained by HI variations may be
13 13 13 436 calculated by subtracting the G Corg values reconstructed from theG Corg-HI crossplots from the G Corg
13 13 437 data. For the studied succession, the obtained residual G Corg record (hereinafter G Corg HI-corr) show an
13 438 overall shape and negative CIE magnitude (~ 3‰) similar to the bulk G CCaCO3 record (Fig. 5b-c). These
13 439 results support the idea that the raw G Corg records cannot be used for reconstructing global carbon cycle
440 perturbations without careful consideration of the OM sources (van de Schootbrugge et al. 2013).
441 5.2. Palaeoenvironmental changes related to the Toarcian Oceanic Anoxic Event
442 The variations observed in the mineralogical and geochemical record of the Rietheim Posidonia
443 Shale section are likely related to the onset of the T-OAE during the Early Jurassic. Hence, we discuss
444 our results in terms of the major palaeoenvironmental changes, which fostered deposition of the
445 Rietheim Posidonia Shale in the NW Tethys Ocean during the early T-OAE.
446 5.2.1. Origin of the negative G13C excursion and global carbon cycle perturbations
13 13 447 G Corg HI-corr and G CCaCO3 records of marls and marly clays show a parallel temporal evolution
448 throughout the Rietheim Posidonia Shale succession, with a pronounced negative CIE (~ 3‰) within the
449 exaratum subzone (Fig. 5). This parallel trend, together with the magnitude of the negative CIE, supports
450 the hypothesis that the early Toarcian negative CIE mainly reflects a major perturbation in the global
451 carbon cycle, likely associated both with an massive output of volcanogenic CO2 and thermogenic
452 methane (CH4) from OM sources in the Karoo-Ferrar large igneous province (e.g., Hesselbo et al., 2000;
453 Svensen et al., 2007; Hesselbo et al., 2007; Suan et al., 2008).
454 However, the magnitude of this negative CIE in the Rietheim Posidonia Shale succession is
455 slightly reduced (~3‰) compared, for example, to lower Toarcian strata from SW Germany
19
456 (Dotternhausen; Röhl and Schmid-Röhl, 2005), the UK (Yorkshire; Hesselbo et al., 2000; Kemp et al.,
457 2005), and northern Siberia (Kelimyar River; Suan et al., 2011) (Fig. 9). In agreement with Kuhn and
458 Etter (1994) and Reisdorf et al. (2011), we attribute this to the presence of a stratigraphic gap in the
13 459 Pliensbachian-Toarcian transition interval. Indeed, most of the tenuicostatum Zone, where G Corg values
460 up to -26‰ are frequently recorded (Röhl and Schmid-Röhl, 2005; Fig. 9b), is missing in the Rietheim
461 Posidonia Shale succession (Kuhn and Etter, 1994). This interpretation is supported by the fact that
13 462 G Corg values recorded just below the limestone bed Unterer Stein in Dotternhausen (Röhl et al., 2001;
463 Röhl and Schmid-Röhl, 2005) and Rietheim are almost identical (~ -32.6‰; Fig. 9a-b). Similar
464 interpretations have also been performed by Röhl and Schmid-Röhl (2005) in the Franconian Posidonia
465 Shale succession from Schesslitz in SW Germany, as well as by Pittet et al. (2014) in the Lusitanian
466 Basin (Portugal) and Ruebsam et al. (2014) in the Lorraine Sub-Basin (NE Paris Basin, France).
467 Therefore, we hypothesize that lower Toarcian strata deposited in a basin-margin setting, such as the
468 Rietheim in northern Switzerland, were probably affected by processes of submarine erosion, reworking,
469 and/or non-deposition due to a fast drop in relative sea-level during the Pliensbachian–Toarcian
470 transition (e.g., Kuhn and Etter, 1994; Hardenbol et al., 1998; Röhl and Schmid-Röhl, 2005; Reisdorf et
471 al., 2011; Pittet et al., 2014; Ruebsam et al., 2014).
472 5.2.3. Continental weathering
473 Changes in the weathering at the source areas and clastic influx that prevailed during deposition
474 of the Rietheim Posidonia Shale succession can be inferred using the mineralogical and geochemical
475 data here obtained (Fig. 10). Indeed, the detritus index, siliciclastic balance and Ln(Al2O3/Na2O) are
476 based on minerals (quartz, phyllosilicates, K-feldspar, Na-plagioclase) and chemical elements, which
477 mainly constitute the detrital component (Nb-Rb-La-K-Ti-Al-Si-Cr-Na-Y-Nd). Consequently, higher
478 values in these detrital proxies could be interpreted here as periods of: (1) stronger rainfall, enhanced
479 chemical weathering and continental runoff, and/or (2) lowered carbonate production in surface water.
20
480 In the Rietheim Posidonia Shale succession, the meaningful increase in the detritus index,
481 siliciclastic balance and Ln(Al2O3/Na2O) observed during the early Toarcian negative CIE (exaratum
482 subzone; Fig. 10) suggest a greater increase in detrital input during this interval relative to the rest of the
483 succession, likely due to an enhanced terrestrial discharge resulting from a strongly accelerated
484 hydrological cycle and higher continental weathering rates induced by intensification of greenhouse
485 conditions during the early T-OAE. Seawater osmium-isotope excursion to more radiogenic values
486 (Cohen et al., 2004) in the NW Tethys Ocean also suggest a warm and humid climate as well as a
487 transient increase in continental weathering rates during the early T-OAE. Likewise, numerical models
488 on regional hydrologic responses to large-scale atmospheric circulation patterns during the Early Jurassic
489 indicate that strong monsoonal circulation dominated along the mid-latitude coasts of Tethys and
490 Panthalassa during the Early Jurassic, including the early Toarcian (e.g., Parrish and Curtis, 1982;
491 Chandler et al., 1992; Dera and Donnadieu, 2012). These monsoonal systems are likely associated with
492 localized pressure cells whose positions are controlled by topography and coastal geography (Chandler
493 et al., 1992). In fact, the large continental areas were mainly located between 30° and 60° latitude in both
494 hemispheres during the Early Jurassic (Fig. 1a), which led to a meridional trade-wind/monsoon-wind
495 circulation pattern (e.g., Parrish and Curtis, 1982; Chandler et al., 1992; Röhl et al., 2001).
496 However, the increase in the detritus index and siliciclastic balance during the negative CIE
497 likely cannot be entirely attributed to a strongly enhanced detrital discharge. The pronounced decrease in
498 both calcite (from 60% down to 14%; Fig. 7e) and Ca (from 20 down to 5%; Fig. 8b) contents at the
499 onset of the negative CIE instead suggest that the increase in these proxies may also reflect a relative
500 decrease in carbonate productivity. Thus, in agreement with several studies on continental weathering,
501 carbonate production and ocean acidification during the early Toarcian (e.g., Röhl et al., 2001; Cohen et
502 al., 2004; Suan et al., 2008; Mattioli et al., 2008, 2009; Hermoso et al., 2009a,b; Dera et al., 2009;
503 Trecalli et al., 2012; Hermoso and Pellenard, 2014; Reolid et al., 2014b), we hypothesize that this drop
21
504 in carbonate content recorded at the onset of the negative CIE in the Rietheim Posidonia Shale
505 succession (Fig. 8b) may likely be the result of a combination of three main factors: (1) dilution by
506 higher detrital input resulting from an accelerated hydrological cycle and enhanced continental runoff,
507 (2) decreasing in carbonate production associated with an enhanced continental runoff and southwards
508 flux of brackish Arctic seawater which controlled turbidity and salinity in surface waters within the NW
509 Tethys Ocean, and/or (3) enhanced carbonate dissolution related to strong seawater acidification
510 conditions driven by a massive injection of CO2 into the atmosphere and then into the ocean.
511 Overall, our mineralogical and geochemical proxies support the hypothesis that the massive
512 injection of CH4/CO2 into the ocean-atmosphere system and subsequent global warming induced during
513 the Early T-OAE an acceleration of the hydrologic cycle and an increase in global chemical weathering
514 rates as well as possibly also decreasing in carbonate production (e.g., Mattioli et al., 2009; Trecalli et
515 al., 2012). However, further investigations are required to improve this latter hypothesis in the Rietheim
516 Posidonia Shale section.
517 5.2.4. Redox conditions
518 In this study, redox conditions that prevailed during the deposition of the Rietheim Posidonia
519 Shale succession are investigated using several proxies (Fig. 11a-f and Fig. 12): pyrite content, sulfide
520 (Fe-Cu-Zn-Co) and organic (TOC-P-As-V-Ni) balances, the V/(V+Ni) ratio, RSTE enrichment factors,
521 relationships between Al-normalized RSTE and TOC and the TOC/Ptotal molar ratios.
522 -Inferences from pyrite, redox-sensitive elements and V/(V+Ni) ratio
523 The stratigraphic distribution of relative contents of pyrite (%), sulfide (Fe-Cu-Zn-Co) and
524 organic components (TOC-P-As-V-Ni) as well the V/(V+Ni) ratio show similar trends across the
525 Rietheim Posidonia Shale succession, with higher values recorded during the negative CIE (exaratum
526 subzone) that then slowly decrease upsection (Fig. 11a-d). These stratigraphic trends, together with
527 threshold values proposed by Hatch and Leventhal (1992) for the V/(V+Ni) ratio (see Section 4.4),
22
528 suggest that sediments in the exaratum subzone represent the onset as well as peak of the T-OAE
529 episode where anoxic (and possibly euxinic) conditions prevailed at the sediment-water interface,
530 whereas the remainder of the falciferum Zone and the overlying bifrons Zone represent a continuation of
531 oxic-to-dysoxic conditions. The modest enrichment (2 to 9), which exhibit the RSTE (Cu, Ni, Zn, V, Co,
532 As) compared to average shale values within the negative CIE also supports the interpretation that
533 anoxic bottom-water conditions developed during the exaratum subzone (Fig. 11e-f). Similar results
534 have been obtained by Brumsack (1991) in the nearby Posidonia Shale succession from the SW
535 Germany.
536 The redox conditions, which prevailed in the sedimentary environment, can also be inferred
537 through the co-variation patterns between the Al-normalized RSTE and TOC (Fig. 12a; Algeo and
538 Maynard, 2004; Tribovillard et al., 2006). Note in Fig. 12 that Al-normalized RSTE (V, Ni, As and Cu)
539 plotted versus TOC allows the discrimination of different redox facies (dysoxic, anoxic and euxinic) in
540 both modern (Irino and Pedersen, 2000; Tada et al., 2000; Piper and Dean, 2002) and ancient (Arthur et
541 al., 1990; Brumsack, 1991; Warning and Brumsack, 2000; Algeo and Maynard, 2004) marine
542 depositional environments. In the Rietheim Posidonia Shale succession, these Al-normalized RSTE-TOC
543 co-variation patterns (Fig. 12b-e) suggest that the predominating palaeo-oxygenation conditions during
544 the deposition of these sediments were (1) oxic during the spinatum subzone, (2) anoxic (possibly
545 euxinic) during the exaratum subzone, and (3) dysoxic, with intermittent anoxia, during the elegans to
546 commune subzones.
547 Based on the similar covariation between the detritus index, siliciclastics, organic and sulfide
548 balances during the negative CIE (Fig. 10), we hypothesize that detrital inputs may indirectly represent
549 the main cause for the modest excess of trace-metal content in the exaratum subzone. Indeed, in
550 marginal environments, enhanced continental runoff during the early T-OAE likely provided high
551 amounts of nutrients, which lead to increased productivity and greater settling of OM. In addition,
23
552 according to Mattioli et al. (2009) and Dera and Donnadieu (2012), enhanced freshwater input in
553 conjunction with the recurrent influx of brackish Arctic seawater through the Viking Corridor could have
554 increased the themohaline stratification within the NW Tethys Ocean during the early T-OAE. Overall,
555 the combined effects of these processes likely induced depletion in bottom-water oxygen levels and thus
556 a modest RSTE enrichment. Consequently, from a geochemical point of view, the organic-rich sediments
557 of the exaratum subzone were likely deposited in a sedimentary environment very similar to the more
558 recent Plio-Pleistocene Black Sea and/or Mediterranean sapropels (e.g., Brumsack, 1991). That is, a
559 stratification of the water column due to salinity differences and the prevalence of anoxic to euxinic
560 bottom-water conditions (e.g., Murray et al., 1989; Brumsack, 1991).
561 -Inferences from phosphorus contents and TOC/Ptotal molar ratios
562 P is an important limiting nutrient in the ocean, driving primary productivity (e.g., Tyrrell, 1999;
563 Benitez-Nelson, 2000). The flux of dissolved P into the ocean is mainly controlled by continental runoff
564 (Föllmi, 1996), and the transfer of P into sediments occurs by incorporation into OM and fish debris or
565 by aggregation onto clay particles, iron, and manganese oxyhydroxides (Delaney, 1998). However, P
566 burial within sediments is strongly influenced by changes in the benthic redox conditions (e.g., van
567 Cappellen and Ingall, 1994; Ingall et al., 2005; Algeo and Ingall, 2007; Mort et al., 2007). Indeed,
568 dysoxic-to-anoxic bottom-water conditions not only promote an enhanced preservation of OM but also a
569 remobilization and release of mineral-associated phosphate and organic P into the uppermost sediments
570 and overlying water column (van Cappellen and Ingall, 1994; Ingall et al., 2005; Mort et al., 2007).
571 Under such conditions, the increased regeneration of P in sediments may lead to both enhanced
572 phosphate recycling (low P burial efficiency) and increased primary productivity in a positive feedback
573 loop (van Cappellen and Ingall, 1994; Ingall et al., 2005), resulting in sedimentary Corg/Porg ratios higher
574 than the Redfield ratio of 106 (Redfield, 1958, Algeo and Ingall, 2007).
24
575 In Rietheim Posidonia Shale succession, TOC/Ptotal molar ratios higher than the Redfield ratio
576 (106 to 383; Fig. 11h) at the base of falciferum Zone suggest that the deposition of organic-rich
577 sediments during the early Toarcian may have occurred under more oxygen-depleted conditions. These
578 interpretations are in agreement with the stratigraphic trends of all RSTE proxies (Fig. 11a-f and Fig.
579 12), which suggested a moderate shift toward anoxic conditions. Conversely, TOC/Ptotal molar ratios
580 lower than the Redfield ratio (<106) associated to relatively high P and CFA contents in the spinatum,
581 upper part of the falciferum and base of commune subzones suggest that the depositional environment
582 evolved towards more oxic-to-dysoxic bottom waters. Note that the presence of oxygen not only
583 promotes the oxidative destruction of OM but also retention of remineralized organic P within the
584 sediment (Algeo and Ingall, 2007). The predominance of bioturbated layers in the upper part of the
585 Rietheim Posidonia Shale succession supports this interpretation (Fig. 2). However, because high P and
586 CFA contents are not coincident with maxima in TOC content in the spinatum, upper part of the
587 falciferum and base of commune subzones, these P-rich intervals might also be related with a high
588 abundance of preserved fish debris within the sediment (see Section 2.2). Large amount of fish debris
589 have also been reported in the falciferum-bifrons Zone boundary from lower Toarcian sections of the SW
590 Germany (Rohl and Schmid-Rohl, 2005) and SE France (Suan et al., 2013).
591 On the other hand, the high P contents associated with the high TOC values during the negative
592 CIE (Fig. 8d) suggest that a part of the remobilized P during redox recycling was likely trapped within
593 the sediments in authigenic P bearing minerals (e.g., authigenic apatite). Indeed, we hypothesize that
594 oxygen-depleted bottom-water condition in conjunction with enhanced surface water stratification led to
595 sequestration of a large part of the regenerated P in the deeper water column by adsorption and/or co-
596 precipitation into authigenic phases (e.g., apatite). This has also been observed in Mediterranean
597 sapropels (e.g., Slomp et al., 2004) and organic-rich sediments from the early Aptian OAE (Stein et al.,
598 2011; Westermann et al., 2013) as well as Cenomanian/Turonian OAE (Mort et al., 2007; Kraal et al.,
25
599 2010), which received a sudden increase of OM. However, as discussed above, a part of the total P
600 contents in sediments deposited under anoxic conditions might also be driven by the abundance of
601 phosphatic fish debris (Schenau and De Lange, 2001), which may either accumulate in sediments due to
602 combination of several environmental factors such as high production of fish debris in productive surface
603 waters, abrupt fish mortality and efficient preservation under anoxic conditions. Indeed, according to
604 selective extraction data from eastern Mediterranean (Slomp et al., 2004) and Arabian Sea sediments
605 (Schenau and de Lange, 2000, 2001) anoxic bottom waters increase the preservation of phosphatic fish
606 hard parts, becoming an important reactive P sink within the sediments (Slomp and Van Cappellen,
607 2007; Paytan and McLaughlin, 2007; März et al., 2014). In the Rietheim Posidonia Shale succession,
608 thin sections analysis of black shale samples show the presence of phosphatic fish debris (Kuhn and Etter
609 1994). This would explain the good correlation between Ptotal and CFA contents through the entire
610 section (Fig. 8d and Fig. 11g-h). Overall, our results suggest that changes in TOC/Ptotal molar ratios
611 cannot be entirely attributed to changes in redox conditions in the Rietheim Posidonia Shale succession.
612 Alternatively, high P contents in organic-rich sediments deposited under anoxic conditions may
613 also be related to the development of benthic microbial mats, such as coccoid bacteria and cryptalgal
614 (Tribovillard et al., 2010; Westermann et al., 2013). In fact, the presence of abundant biomass in the
615 sediment together with the development of microbial mats at the sediment–water interface might have
616 limited the exchanges between the water column and the pore space within sediment, hampering pore
617 water renewal and promoting the quantitative trapping of remobilized P in the sediment during redox
618 recycling. This environmental process has been recently proposed to operate during the deposition of
619 Jurassic organic-rich carbonates from the French Jura Mountains (Tribovillard et al., 2010).
620
621 6. Conclusions
26
622 The results of this high-resolution geochemical-mineralogical research allow for the following
623 generalizations and conclusions:
13 13 624 1. Carbon isotope (G Corg-corr, G CCaCO3), Rock-Eval pyrolysis, and TOCcff data provide clear
625 evidence that the pronounced ~3‰ negative CIE recorded in the Rietheim Posidonia Shale
626 succession during the onset of the T-OAE mainly reflects a major perturbation in the global
627 carbon cycle, likely induced by a massive injection of 13C-depleted carbon into the whole
628 exchangeable carbon reservoir, rather than changes in the type of OM and/or degree of
629 stratification of bottom waters.
13 15 630 2. The Rock-Eval pyrolysis data and G Corg-G Ntotal crossplots suggest that the OM in the upper
631 Pliensbachian and lower Toarcian interval is primarily of terrestrial and marine origin,
632 respectively, and deposited in an epicontinental sea setting.
633 3. The detrital proxies [notably, detritus index, siliciclastic balance and Ln(Al2O3/Na2O)]
634 corroborate that the elevated pCO2 levels and greenhouse climate prevailing during the T-OAE
635 induced an acceleration of the hydrologic cycle and an increase in chemical weathering rates.
636 4. The redox-sensitive proxies [notably, relative pyrite contents, organic and sulfide balances,
637 V/(V+Ni) ratios, RSTE enrichment factors, Al-normalized RSTE versus TOC co-variation
638 patterns, and TOC/Ptotal molar ratios] suggest that the sedimentary deposition of the Rietheim
639 Posidonia Shale succession took place under oxic-to-dysoxic bottom-water conditions, slightly
640 shifting toward anoxic and possibly euxinic conditions during the negative CIE. The T-OAE
641 influence on the sedimentary environment prevailed until at least to the bifrons Zone. Overall,
642 our results further support the idea that the influx of brackish Arctic seawater through the Viking
643 Corridor could have facilitated the thermohaline stratification and subsequent dysoxic-to-anoxic
644 bottom-water conditions within the NW Tethys Ocean (Dera and Donnadieu, 2012).
27
645 5. Enhanced rainfall and river freshwater inputs together with recurrent discharges of boreal
646 brackish seawaters into the NW Tethys through the Viking Corridor (Dera and Donnadieu,
647 2012), appear to be a plausible scenario for explaining the P-enrichments (mainly authigenic P)
648 in the Rietheim Posidonia Shale sediments during the early Toarcian warming event.
649 6. The geochemical characteristics of organic-rich sediments from the exaratum subzone suggest
650 they were likely deposited in a sedimentary environment very similar to the Black Sea and/or
651 Mediterranean sapropels.
652 Taken as a whole, our high-resolution geochemical and mineralogical data provide a solid record
653 of the major palaeoenvironmental changes that fostered deposition of the Rietheim Posidonia Shale
654 succession within the NW Tethys Ocean during the development of the T-OAE, and corroborate that
655 deposition of the lower Toarcian black shales was mainly driven by a global climate change even though
656 it was influenced by local environmental changes.
657
658 Acknowledgements
659 This research was funded through the GeoNova project and Université de Lausanne, with the
660 financial support of the Swiss University Conference (CUS). We acknowledge Jean-Claude Lavanchy
661 (Centre d’Analyse Minérale, UNIL), Hassan Saleh Khozyem/Tiffany Monnier (ISTE-UNIL) and
662 Brahimsamba Bomou (ISTE-UNIL) for their support with the XRF, Rock-Eval and elemental analyses,
663 respectively. We thank Thomas Algeo (University of Cincinnati) for his comments and suggestions,
664 which significantly helped to improve the early version of this work. We further thank Christine Laurin
665 for reviewing the grammar. Lastly, we specially thank the editor Finn Surlyk and the reviewer Matías
666 Reolid (Universidad de Jaén) for their constructive and helpful comments, which significantly improved
667 this work.
668
28
669 Appendix A. Supplementary material
670 Supplementary data associated with this article can be found in the online version at http…
671 File Excel. Rietheim Posidonia Shale dataset.exe
672
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992
993
994
995
42
996 Figure captions
997
998 Figure 1. (a) Palaeogeographical maps for the early Toarcian (redrawn from Vrielynck and Bouysse,
999 2003) showing the main large igneous provinces (Karoo-Ferrar and North Sea), the Viking Corridor, the
1000 NW Tethyan region, and the Rietheim Posidonia Shale succession studied herein. VC indicates the
1001 location of the Viking Corridor. Locations of the Lower Toarcian successions discussed in this study are
1002 noted as black circles: Dotternhausen, SW Germany (D; Röhl et al., 2001); Schesslitz, SW Germany (S;
1003 Küspert, 1982; Prauss et al., 1991); Yorkshire, England (Y; Cohen et al., 2004; Kemp et al., 2005); and
1004 Kelimyar River, N Siberia (KR; Suan et al., 2011). (b) Map showing the location of the Rietheim
1005 Posidonia Shale succession and geological setting in northern Switzerland.
1006
1007 Figure 2. Biostratigraphy and lithostratigraphy for the upper Pliensbachian to lower Toarcian interval at
1008 the Rietheim Posidonia Shale succession. The ammonite biostratigraphy is from Kuhn and Etter (1994).
1009 Photographs of the drill core and its stratigraphic position are also shown.
1010
13 13 1011 Figure 3. Stratigraphic distribution of the: isotopic compositions of (a-b) carbon (G Corg, G CCaCO3), (c)
18 15 1012 oxygen (G OCaCO3), (d) nitrogen (G Ntotal), (e) total organic carbon (TOC), (f) total organic carbon from
1013 the carbonate-free fraction (TOCcff), (g) hydrogen index (HI), and (h) Rock-Eval Tmax from the Rietheim
1014 Posidonia Shale succession. The shaded interval indicates the negative carbon isotope excursion (CIE)
1015 related to the T-OAE.
1016
1017 Figure 4. (a) Hydrogen Index (HI) vs. Oxygen Index (OI) crossplot illustrating the organic matter origin
1018 of the Rietheim Posidonia Shale samples. Note that the organic matter was primarily of marine origin
13 1019 during the lower Toarcian (exaratum to commune subzones). (b) G Corg-HI crossplots. The least square
43
13 1020 linear regression is based on samples from outside of the T-OAE. Note that G Corg values slightly
13 15 1021 decrease with increasing HI values. (c) G Corg-G Ntotal crossplots showing the depositional settings of
13 15 1022 the Rietheim Posidonia Shale sediments. The average G Corg-G Ntotal values of the five depositional
1023 settings are from Algeo et al. (2014).
1024
13 13 1025 Figure 5. Comparison between the bulk G Corg records (a) with G Corg corrected for changing OM
13 1026 sourcing (b) and G CCaCO3 (c). Note that the bulk organic C-isotope data were corrected using the HI and
1027 raw C-isotope data from the spinatum and bifrons zones. The blue, red and black lines represent the 9-
1028 point moving average of the isotopic data.
1029
1030 Figure 6. X-ray diffractograms of typical bulk rock samples from the early Toarcian negative carbon
1031 isotope excursion (CIE) in the Rietheim Posidonia Shale succession.
1032
1033 Figure 7. Stratigraphic distribution of the bulk mineralogical semiquantification from the Rietheim
1034 Posidonia Shale sediments. The shaded interval indicates the negative carbon isotope excursion (CIE)
1035 related to the T-OAE.
1036
1037 Figure 8. (a-d) Stratigraphic profiles of the major geochemical associations obtained by the
1038 compositional Q-mode cluster analysis. The shaded interval indicates the negative carbon isotope
1039 excursion (CIE) related to the T-OAE. The red lines represent the 9-point moving average of the data. (e)
1040 Dendrogram of association of variables obtained by applying Ward clustering algorithm based on the
1041 variation matrix of elemental geochemical dataset.
1042
44
1043 Figure 9. Overview of upper Pliensbachian to lower Toarcian carbon-isotope records from representative
1044 Tethyan and Boreal successions. Carbon-isotope data from (a) Rietheim Posidonia Shale, N Switzerland
1045 (this study; the black line represent the 9-point moving average of the isotopic data), (b) Dotternhausen
1046 and Zimmern, SW Germany (modifed from Küspert, 1982; Prauss et al., 1991 and Röhl et al., 2001), (c)
1047 Yorkshire, England (Cohen et al., 2004; Kemp et al., 2005) and (d) Kelimyar River, N Siberia (Suan et
1048 al., 2011). See Fig. 1a for the geographical location of these stratigraphic sections. The shaded interval
1049 indicates the negative carbon isotope excursion (CIE) related to the T-OAE. Note that carbon-isotope
1050 profiles in geographically distant sections exhibit a similar trend in the lower Toarcian interval.
1051
1052 Figure 10. Stratigraphic distribution of the main detrital proxies obtained from the Rietheim Posidonia
1053 Shale succession. (a) detritus index (summing of quartz, phyllosilicates, K-feldspars, Na-plagioclases);
1054 (b) Ln(Al2O3/Na2O); (c) XRF siliciclastic balance (Nb, Rb, La, K, Ti, Al, Si, Cr, Na, Y, Nd). Balances
1055 are represents according to equation 1 (see Section 3.2.5). The red lines represent the 9-point moving
1056 average of the data. The shaded interval indicates the negative carbon isotope excursion (CIE) related to
1057 the T-OAE.
1058
1059 Figure 11. Stratigraphic distribution of the main redox proxies obtained from the Rietheim Posidonia
1060 Shale succession. (a) relative content of pyrite (%); (b-c) sulphide (Zn, Cu, Fe, Co) and OM (P, As,
1061 TOC, Ni, V) balances. Balances are represents according to equation 1 (see Section 3.2.5); (d) V/(V+Ni)
1062 ratios. Redox zones are from Hatch and Leventhal (1992); (e-f) elements enrichment factors (EF) for the
1063 OM and sulphide associations; (g) relative content of fluorapatite (%); (h) TOC and total phosphorus
1064 (Ptotal) relationship (TOC/Ptotal). The red lines represent the 9-point moving average of the data. The
1065 shaded interval indicates the negative carbon isotope excursion (CIE) related to the T-OAE.
1066
45
1067 Figure 12. Al-normalized redox-sensitive trace-elements (V, Ni, As and Cu) plotted versus total organic
1068 carbon from the Rietheim Posidonia Shale sediments. This plot is used to discriminate between different
1069 redox facies (dysoxic, anoxic and euxinic; e.g., Algeo and Maynard, 2004; Trivovillard et al., 2006).
1070 Redox-sensitive trace-metal data from the Cariaco Basin (Piper and Dean, 2002), the Southern California
1071 margin (Irino and Pedersen, 2000; Tada et al., 2000), Posidonia Shale from SW Germany (Brumsack,
1072 1991), upper Pennsylvanian black shales (Algeo and Maynard, 2004), and Mediterranean and Black Sea
1073 sapropels (Warning and Brumsack, 2000) are also plotted.
46
*Highlights (for review)
Highlights
-The T-OAE negative CIE reflects a 13C-depleted carbon injection to ocean-atmosphere
-Detrital proxies suggest an increase in continental weathering rates during the T-OAE
-Redox proxies suggest moderate anoxic bottom-water conditions during the T-OAE
-Dysoxic-to-anoxic conditions prevailed until at least to the bifrons Zone
-P-enrichments are observed during the T-OAE Figure 1
Figure 1 (Double column - full page width)
Legend (a)Rietheim (b) Present day coastline Continental/oceanic canton Basel Aargau crust transition N Zürich Y Fault 0 50 km Oceanic ridge Bern S Thrust D Oceanic subduction zone Switzerland Igneous provinces Emergent land Shallow marine Geneva Deep marine N 1000 km Rietheim 7°E 8°E 9°E Posidonia Shale
Rietheim Early Toarcian (~180 Ma) Rhine BOREAL KR ASIA Basel 60°N OCEAN VC EUROPE Zürich LAURASIA North SE ASIA Sea Aar 30°N NORTH AMERICA Mega Lhassa
47°N Aar Bern N 0° PANTHALASSA OCEAN TETHYS OCEAN AFRIQUE GONDWANA Legend SOUTH INDIA Tabular Jura Molasse s.l. in the Jura Mountains 30°S AMERICA Folded Jura Subalpine molasse and folded molasse Karoo AUSTRALIA Tertiary rift basins Variscan basement and Tertiary Ferrar volcanites (Hegau) 60°S Upper freshwater molasse Subalpine chains, Helvetic nappes Upper marine molasse ANTARCTICA and Prealps klippen Faults: Thrusts Major thrusts Figure 2
Figure 2 (1.5 column)
Rietheim Posidonia Shale
tage Height S A-ZoneA-subzone (cm) 480 Monotisbank Legend
440 Black to grey bituminous marly clay ()Tonmergel-Fazies 400 Black to brownish laminated bituminous 360
bifrons marly clay (Kartonschiefer-Fazies )
commune 320 Light-grey to dark-grey marly clay 280 Light-grey bioturbated marl 240 Limestone beds 200 Bioturbated layers
falciferum 160
120 S1 S2 S3 S4 S5 S6 S7 S8 S9 S10 80 0 Top 40 5
Early Toarcian 0
elegans
falciferum
Spinatumbank
10 Oberer Stein -40 Oberer Stein 15
-80 Homogene Kalkbänke
Unterer Stein -120 Homogene 20 Kalkbänke -160 25 Base
exaratum -200 Unterer Stein 30 -240 ? semi-? ten. celatum 35 -280 Spinatumbank
-320 40
Pliens.
spinatum
spinatum
45 cm Homogene Kalkbänke Figure 3
Figure 3 (Double column - full page width)
Rietheim Posidonia Shale 15 d13 d N(‰ N air) C(CaCO3 ‰ VPDB) total 2 TOCcff (wt.%) Tmax (°C) 01234 StageA-ZoneA-subzoneHeight(cm) -5 -4 -3 -2 -1 0 1 2 01015520400 420 440 460 480 440 (a) (b) (c) (d) (g) (h) Monotisbank (e) (f) 400 360
bifrons
commune 320 280 240 200
falciferum 160 120
mature
80 Immature 40
Early Toarcian
elegans
falciferum 0
-40 Stein Oberer -80 -120
Kalkbänke -160 Homogene
exaratum
-200 CIE
Stein
? semi-? -240 Unterer ten celatum -280
spin. spin. -320 Pliens. -34 -32 -30 -28 -26 -24 -10-9-8-7-6-5-4-3-2 024681012 0 200 400 600 800 d13 d18 TOC (wt.%) Hydrogen Index Spinatumbank C(‰ VPDB) O(‰ VPDB) org CaCO3 (mgHC/gTOC) Figure 4 Figure 4 (single column)
(a) 900 Type I Legend marls - marly clays 800 commune falciferum elegans 700 exaratum semicelatum Type II spinatum 600 limestones commune elegans exaratum 500 semicelatum spinatum 400 Mixtures
(mgHC/gTOC) Type II+III
Hydrogen Index 300
200
100 Type III Type IV 0 0 100 200 300 400 500 Oxygen index
(mgCO2 /gTOC) (b) -24 d13C = -0.0016*HI - 27.741 org T-OAE -26 R²=0.42 background -28
‰ VPDB) -30
org
C( -32
13
d -34 0 100 200 300 400 500 600 700 800 Hydrogen Index (mgHC/gTOC) (c) 8
6 Shelf Oceanic
2 4 Upwelling 2
‰ N air)
total 0 Epeiric
N( Oceanic-Mediterranean
15
d -2
-4 -38 -36 -34 -32 -30 -28 -26 -24 -22 -20 -18 d13 C(org ‰ VPDB) Figure 5 Figure 5 (single column)
Rietheim Posidonia Shale
d13 C(org HI-corr ‰ VPDB) tageA-Zone Height(cm) -4 -2 0 2 S A-subzone480 440 (a) (b) (c)
Monotisbank 400 360
bifrons
commune 320 280 240 200
falciferum 160 120 80 40
Early Toarcian 0
elegans
falciferum -40 Stein Oberer -80 -120
Kalkbänke -160 Homogene
exaratum -200 CIE
Stein
Unterer ? semi-? -240 ten celatum -280
spin. spin. -320
Pliens. ~5‰ ~3‰ ~3‰
Spinatumbank -32 -30 -28 -26 -413 -2 0 2 d13 d C(‰ VPDB) C(org ‰ VPDB) CaCO3 Figure 6
Figure 6
Mu K-Feld Bi Ka An MuQz CFA
Illi Chl MuQz CaQz Plag Ca Do CFA CaPy Ca Py Ca CaQz Py Qz 3000
2500 Top -123 cm 2000 -150 cm 1500 -220 cm
CIE 1000
Counts second per -236 cm 500 -246 cm 0 Base 5 101520253035404550556065 2(CuK)θα Figure 7
Figure 7
Rietheim Posidonia Shale Phyllosilicates (%) K-Feldspar (%) Dolomite (%) Fluorapatite (%) Unquantified (%) StageA-ZoneA-subzoneHeight(cm) 02550 02468 02468 0.00.61.21.8 012345 480 440
Monotisbank 400 (a) (b) (c) (d)(e) (f) (g) (h) (i) (j) 360
bifrons
commune 320 280 240 200
falciferum 160 120 80 40
Early Toarcian
elegans 0
falciferum -40 Stein Oberer -80 -120
Kalkbänke -160 Homogene
exaratum -200 CIE
Stein
Unterer ? semi-? -240 ten celatum -280
spin. spin. -320 Pliens. strata 0 6 12 18 01234 0 255075100 0 5 10 15 20 02468
missing Spinatumbank Quartz (%) Na-Plagioclase (%) Calcite (%) Ankerite (%) Pyrite (%) Figure 8
Figure 8
Rietheim Posidonia Shale (a) Siliciclastics group (b) Carbonate group Si (%) K(%) Cr (μg/g) Nb (μg/g) Y(μg/g) Mg (%) Ca (%) Height(cm) 0 5 1015 20 25 0123 40 80 120 160 0 5 1015 20 25 10 20 30 40 50 024 0102030 StageA-ZoneA-subzone480 440
Monotisbank 400 360
bifrons commune 320 280 240 200
falciferum 160 120 80 40 Early Toarcian 0
falciferum
elegans -40 Stein -80 Oberer -120
Kalkbänke
-160 Homogene
CIE exaratum -200
Stein ? ? -240 Unterer ten semi. -280
spin. spin. -320 Pliens. strata 05100.0 0.2 0.4 0.6 0.1 0.2 0.3 0408012010 30 50 70 90 20 40 60 80 0 500 1000 1500 missing Spinatumbank Al (%) Ti (%) Na (%) Rb (μg/g) La (μg/g) Nd (μg/g) Mn (μg/g)
Rietheim (c) (d) (e) Posidonia Shale Sulfide group Organic group Variance ln(xij /x ) Elemental Cu (μg/g) Co (μg/g) P(μg/g) V(μg/g) 0.0 0.2 0.4 0.6 0.8 1.0 associations Height(cm) 050100 10 30 50 0 1000 2000 0200400 StageA-ZoneA-subzone480 P 440 As
Monotisbank 400 TOC 360
bifrons Zn* commune 320 Cu* 280 Fe* 240 Co* 200 rganic + oNi sulfide* falciferum 160 V 120 Mn 80 Ca 40 Early Toarcian 0 Mg
falciferum
elegans -40 carbonate Nd Stein Oberer Y -80 Na -120 Cr
Kalkbänke
-160 Homogene Si
exaratum -200 CIE
Stein
Unterer Al ? ? -240 ten semi. -280 Ti
siliciclastics
spin. spin. -320 K Pliens. strata 1357 20 80 140 200 036912 05010010 70 130 190 La missing Spinatumbank Fe (%) Zn (μg/g) TOC (%) As (μg/g) Ni (μg/g) Rb Nb Figure 9
Figure 9
Rietheim Dotternhausen & Zimmern (c)Yorkshire (d) Kelimyar River (a)N Switzerland (b) SW Germany UK N Siberia
tage Height(cm) Northwestern Tethys Boreal S A-Zone480 440 400 360 Height bifrons StageA-Zone (m)
320 Height(m) Height(m) StageA-Zone StageA-Zone 30 280 U. T. 12 14 240 25 200 10 160 12 braun. 120 20 8 80 bifrons 10 40 15 Lower Toarcian 6 falciferum 0 8
falciferum -40 commune 10 -80 ? 4
6 Lower Toarcian -120 5 -160 Lower Toarcian 2 -200 Lower Toarcian 4
falciferum
falciferum
? -240 0 CIE ten a. 0 -280 strata 2 missing
-320 ten. -5 -34 -32 -30 -28 -26 -24 -2
Pliens.
spinatum 0 d13
C(‰ VPDB) viligaensis org -10 -4
Pliensbachian -2 tenuicostatum -34 -32 -30 -28 -26 -24 -34 -32 -30 -28 -26 -24 13
spinatum -15 d C(‰ VPDB) d13 org
Pliensbachian C(org ‰ VPDB) -34 -32 -30 -28 -26 -24 13
Pliens. d C(org ‰ VPDB)
spinatum Figure 10 Figure 10 (single column)
Detrital proxies Rietheim Posidonia Shale Clastic influx Clastic influx
Detritus (%) b{Siliciclastic/Carbonate} StageA-ZoneA-subzoneHeight(cm) 25 50 75 -2.2 -1.6 -1.0 480 440 (a) (b) (c) 400 360 Monotisbank
bifrons
commune 320 280 240 200
falciferum 160 120 80 40
Early Toarcian
elegans 0
falciferum -40 Stein Oberer -80 -120
-160 Kalkbänke
Homogene
exaratum -200 CIE
? ? -240 Stein semi- Unterer ten celatum -280
spin. spin. -320
Pliens. 2.8 3.0 3.2 3.4 strata missing Ln(Al O /Na O)
Spinatumbank 23 2 chemical weathering Figure 11
Figure 11 (Double column - full page width)
Redox proxies
Rietheim anoxia/euxinia Posidonia Shale OM preservation b Pyrite (%) {Organic/Carbonate} EFaverage shale Fluorapatite (%) StageA-Zone Height(cm) 02468-3.0 -2.4 -1.8 -1.2 -0.6 0.1 1 10 0.0 0.6 1.2 1.8 A-subzone480 440
Monotisbank (a) (b) (c) (d) (e)(f) (g) (h) 400 360
bifrons
commune As Co 320 EF EF VEF CuEF 280 NiEF ZnEF 240 200
falciferum 160
EF = 1 EF = 1 EF = 3 EF = 3 120 Oxic
Anoxic Euxinic
Dysoxic 80
40 Redfield ratio Early Toarcian 0
falciferum
elegans -40 Stein Oberer -80 -120 -160
Kalkbänke
Homogene
CIE exaratum -200
? ? -240 Stein semi- Unterer ten celatum -280
spin.
spin. -320 Pliens. strata -2.5 -2.1 -1.7 -1.3 -0.9 0.4 0.6 0.8 1.0 0.1 1 10 0 100 200 300 400 missing b{Sulfide/Carbonate} V/(V+Ni) EFaverage shale TOC/P
Spinatumbank total
euxinia decreased oxygenation primary productivity Figure 12 Figure 12 (single column)
Anoxic Euxinic euxinic facies threshold threshold (weak TOC-TE correlation) (a)
TE/Al
anoxic facies (strong TOC-TE correlation) dysoxic facies background concentration 2.5 10 40 60 TOC (wt. %) 50 (b)
-4 40 30 20
V/Al (x10 ) 10 0 012345678910111213 40 TOC (wt. %)
30 (c)
-4
20
Ni/Al (x1010 )
0 012345678910111213 TOC (wt. %) 18 15 (d)
-4 12 9 6 As/Al (x103 ) 0 012345678910111213 TOC (wt. %) 25
20 (e)
-4 15
10
Cu/Al (x105 )
0 012345678910111213 TOC (wt. %)
Legend Redox facies Posidonia Shale Posidonia Shale Sapropels Rietheim SW Germany Anoxic - Euxinic semicelatum Base Anoxic Cariaco Basin exaratum Upper Pennsylvanian Anoxic elegans black shales falciferum Southern California Anoxic commune Top margin Dysoxic Dysoxic