Oxygen isotope evolution

Measured δ¹⁸O values show a bell-shaped evolution with a decline from -4.7 hto -7.0hin the first 2’000 m of the section, followed by an increase back to -4.5 hat

the top of the record (Fig. 4.4), and little dispersion with 1σ = 0.29 h of moving standard deviation (200 m window).

Despite the difference in absolute values acknowledged above, this bell-shaped evolution is similar in nature with the bell-shaped global δ¹⁸O record of Cramer et al. 2009 (Fig. 4.2, 4.5, A.8, and A.10), which shows values evolving from -0.5 h to -1.3 h in the early Eocene paired with global warming of the EECO (corresponding to an increase in global temperature from ca. 9.5 to 12 ^◦C, Zachos et al. (2001)) and followed by values increasing from -1.3 hto 0.5 hduring the subsequent global cooling (Fig. 4.2 and 4.5, ca. 12 to 6.5 ^◦C, Zachos et al. (2001)).

In particular therefore, the -1 hmagnitude negative shift that marks the EECO warming is similar in the data presented here and on the coeval global record of Cramer et al. 2009. However, a locally weighted linear regression shows that the measured δ¹⁸O reaches minimum values (below 97th percentile) between 49.4 and 49.0 Ma, i.e. ca 2.1 Ma later in the Pyrenean succession than in the deep-sea global record which occurs at 51.3±0.2 Ma (lower 97th percentile). In addition the rate of decline of δ¹⁸O before 51.3±0.2 Ma are 3 times higher in the Pyrenean record than in the global profile (Fig. 4.5). Aside from diagenetic effects, the δ¹⁸O signature of carbonates depend primarily on the temperature of calcification (related to water temperature in the environment), and on the δ¹⁸O of the surrounding water (δ¹⁸Ow) which itself is dependent on the mixing of global and local waters as well

as salinity and evaporation (Garzione et al. (2004), Pearson (2012)).

We rule out the possibility that warming continued in the Pyrenees until 49.2 Ma while cooling already started globally 2.1 Ma earlier. Such lag is not observed in coeval measured δ¹⁸O in the Barinatxe-Gorrondatxe section (Fig. 4.5, Payros et al. (2015)) ca. 200 km west of the study area, suggesting that our observation would imply a very local warming yet sustained for 2 Ma. Although there might have been some latitudinal diachrony in the onset and end of the EECO warming event (Payros et al. (2015)), the very globally homogeneous climate and reduced latitudinal gradient of the EECO period (Andreasson and Schmitz (2000)) excludes the perpetuation of a local temperature difference of several degrees for 2.1 Ma in such a spatially limited area. The critical period of the observed perturbation of the δ¹⁸O signal corresponds to a period of major tectonic change in the Pyrenees (Fig. 4.5). First, at 56 Ma, the convergence rate between Iberia and Europe records an important acceleration from a period considered as a tectonic quiescence with no convergence rate to 3.2 mm/yr (Macchiavelli et al. (2017)). This major

Pyrenees Eocene topographic evolution 69

EECO (Westerhold et al. 2018)

-2 -1 0 1

erosion induced exhumation period (Whitchurch et al., 2011)

mm/yr

1 3 5

Iberia-Europe convergence rates Southern Pyrenees shortening rates high shortening ratesslower shortening ratesquiescence

shift

Figure 4.5: Comparison between South Pyrenean, global and North Iberian continental margin δ¹⁸O records. A third order polynomial fit was added to the South Pyrenean record and the global δ¹⁸O records. δ¹⁸O global record from Cramer et al. (2009), north Iberian continental marginδ¹⁸O record from Payros et al. (2015). Pyrenean shortening rates from Grool et al. (2018), Iberia-Europe convergence rates from Macchiavelli et al.

(2017). EECO extension from Westerhold et al. (2018).

perturbation of the boundary conditions marks the onset of Pyrenean mountain building and is expressed by a 4-folds increase of shortening rates in the South Pyrenees, from 1 mm/yr to 4 mm/yr (Grool et al. (2018), Verg´es et al. (2002)), a ∼2-fold increase in the number of recorded active structures and a pulse of exhumation at ca. 50.9 Ma± 5 Ma (Beamud et al. (2011), Fitzgerald et al. (1999), Thomson et al. (2019), Whitchurch et al. (2011)). These major changes from at least 56 Ma onwards also correspond to a period of inferred major topographic growth evidenced by the onset of flexural subsidence (Verg´es et al. (1995)), the regional flooding of the basin at the base of the Eocene and the subsequent arrival of the first clastics (i.e. Roda, Castigaleu and Castissent fluvio-deltaic systems) (Garc´es et al. (2020), Puigdef`abregas and Souquet (1986), Verg´es et al. (2002)). In

the North Pyrenean domain, the Pyrenean exhumation supplied coarse-grained depositional systems such as the Poudingues de Palassou Formation (Vacherat et al. (2017)). We propose that the observed perturbation of the δ¹⁸O signal in the foreland basin is the signal of the major topographic development of the mountain range between ca. 54 and ca 49 Ma because of two main processes. First, the increase of the mean elevation of the Pyrenees during that period, one aspect of its topographic growth, would deliver increasingly negative δ¹⁸O waters to the basin because oxygen isotopic composition of precipitation decreases with increasing elevation through Rayleigh distillation process during orographic rainout (Caves Rugenstein and Chamberlain (2018), Poage and Chamberlain (2001), Rowley and Garzione (2007)). Huyghe et al. (2012) demonstrated with their morphological-hydrological model that such process could explain the negative offset betweenδ¹⁸O values measured on early to middle Eocene oyster shells in the South pyrenean and those recorded in coeval deposits of the Paris Basin. Our extensive dataset spanning a longer 12 Ma and denser (658 data points) record than previously presented by Huyghe et al. (2012) (5 Ma and 12 data points) suggests that the phenomenon identified by their study may have been at work already at 54 Ma and continued until 49 Ma. By constantly adding negative δ¹⁸O waters to the restricted foreland waters, and assuming a lack of mixing and equilibration in the restricted Tremp-Graus-Ainsa basin (also increasingly restricted with shortening) with the global ocean, the growth of the orogen steered the local δ¹⁸O composition of foreland carbonates towards more negative values at a faster pace than the global trend and forced a perpetuation of this trend during ca. 2 Ma longer than in the global oceanic reservoir. The δ¹⁸O record imposed an increasingly higher and increasingly negative water input superimposed on the global decrease and brought therefore the two δ¹⁸O curves to diverge.Second, nearshore and relatively enclosed environments are strongly influenced by freshwater input and evaporation effects. Since δ¹⁸O increases with salinity by approximately 0.2 h/psu (1 psu

= 1 g of NaCl per kg water, Maclachlan et al. (2007), Marshall (1992)) a higher proportion of freshwater input in the relatively confined South Pyrenean basin during Eocene times (Vacherat et al. (2017)) could potentially be the cause of the local δ¹⁸O values toward lower (more negative) values. An input of freshwater shifting theδ¹⁸O minimum from -5.5 hat 52 Ma to -6.5 hat 49 Ma would imply a change in salinity of ca. 5 psu during this period. This is roughly the difference between two contrasted marine environments such as mean proximal Bengal Bay (31 psu, Akhil et al. (2016)) and mean Atlantic salinity values (35 psu, Le Vine

Pyrenees Eocene topographic evolution 71 et al. (2007)). Past salinity is difficult to constrain (Huyghe et al. (2012)), but a decrease in salinity due to freshwater input alone would imply a relative salinity change equivalent to an environmental change from proximal to open ocean in the South Pyrenean basin. Developing a mixing model of freshwater input and oceanic boundary conditions to assess the resulting composition of basin waters would require constraining not only the precipitation regime but also the size of drainage areas and the physiographic configuration (areal and bathymetric) of the receiver basin. These constraints are beyond the scope of the present work but certainly highlight the need for further research and modeling for the prediction of the local distortion of isotopic proxies by tectonic topography. Nevertheless, the effect of increasing freshwater on the local decrease in salinity may have exacerbated the divergence between the local and global δ¹⁸O records.

After the Pyrenean δ¹⁸O record minimum, the oxygen isotope record returns to more positive values which are in more agreement with global δ¹⁸O values. We suggest that thisδ¹⁸O signal, reflects the post-EECO global decrease inδ¹⁸O values towards todays cooler climatic conditions (Cramer et al. (2003)), with a stable Pyrenean relief. Consistently, the shortening rate of the South Pyrenees slows down significantly (Grool et al. (2018), Vacherat et al. (2017), Verg´es et al. (1995)) close to the δ¹⁸O minimum in the South Pyrenean record (Fig. 4.5). Huyghe et al. (2012) suggested that the Pyrenees had reached an elevation of 2000±500 m between 49 and 41 Ma. Our high-resolution δ¹⁸O record supports this hypothesis, refining the timing at which the Pyrenees would have reached such elevation. Based on the difference in minimum and rate between the Pyrenean and global δ¹⁸O record, we suggest that the Pyrenees topography increased from at least 54 Ma and had reached this elevation by ca. 49 Ma. This is consistent with exhumation, convergence, shortening rates and topography estimations, which increase at the beginning of the Eocene (Fig. 4.5). Altogether, previous studies (e.g., Grool et al.

(2018), Macchiavelli et al. (2017), Thomson et al. (2019), Whitchurch et al. (2011)) and our results allow us to further the knowledge on the topographic evolution of the Pyrenees (Fig. 4.6) and highlight the importance of the Cretaceous to Paleocene times as a critical period for the full reconstruction of the topographic evolution of the Pyrenees. This study confirms that δ¹⁸O whole-rock data from foreland basins are fundamental data for understanding the tectono-climatic controls on the topographic evolution of mountain ranges, and provides tools for a better understanding of the interaction between tectonic and climatic signals in other active settings.

0

Paleocene Eocene Oligocene Miocene Plio.Plei.

Ypresian Lutetian BartonianPriabonian

isotopical and flexural model constraints on topography flexural model constraints on topography

Macchiavelli et al. 2017

Figure 4.6: Paleotopography evolution of the Pyrenees. Eocene to present maximum paleotopography estimates from Huyghe et al. (2012) and Curry et al. (2019). Erosion induces exhumation cooling periods from Whitchurch et al. (2011). Emerged Pyrenean mountain range timing from Vacherat et al. (2017) and end of deformation timing from Verg´es et al. (2002). Dashed grey lines indicate high uncertainty (Grool et al. (2018)).